Water-Rock Interaction: Proceedings of the Tenth International Symposium on Water-Rock Interaction, WRI-10, Villasimius, Italy, 10-15 July 2001

WATER-ROCK INTERACTION

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PROCEEDINGS OF THE TENTH INTERNATIONAL SYMPOSIUM ON WATER-ROCK INTERACTION WRI-10 / VILLASIMIUS / ITALY / 10-15 JULY 2001

Water-Rock Interaction Edited by

Rosa Cidu Department of Earth Sciences, University of Cagliari, Italy

Volume 1

4

/ EXTON(PA) / TOKYO A.A.BALKEMA PUBLISHERS LISSE/ ABINGDON

Cover: The coast of Nebida, Sardinia (Photo by F. Di Gregorio)

Copyright 02001 Swets & Zeitlinger B.V., Lisse, The Netherlands All rights reserved. No part of this publication may be reproduced, stored in a retrieval system, or transmitted in any form or by any means, electronic, mechanical, photocopying, recording or otherwise, without the prior written permission of the publisher. Published by: A.A.Balkema, a member of Swets & Zeitlinger Publishers www.balkema.nl and www.szp.swets.nl For the complete set of two wolumes, ISBN 90 2651 824 2 For Volume 1, ISBN 90 265 1 834 X For Volume 2, ISBN 90 2651 835 8 Printed in the Netherlands

Table of contents xv

Preface Organisation

XVII

Keynote lectures Water-rock interactions in mudrocks and similar low permeability material

3

A .H Bath, F.J.Pearson, A. Gautschi & H,N. Waber Interpreting kinetics of groundwater-mineral interaction using major element, < trace element, and isotopic tracers S.L.Brantley, M,Bau, S. Yau, B.Alexander & J. Chesley

13

Reducing ambiguity in isotopic studies using a multi-tracer approach Th.D.Bullen, A. F. White, C. W.Childs & J. Horita

19

Significance of geochemical signatures in sedimentary basin aquifer systems K M.Edmunds

29

Interactive processes due to subglacial volcanic activity - Local phenomena with global consequences H.Kristmannsdbttir

37

The inverse modeling of water-rock interaction G.Ottonello

47

Hydrothermal Water/RocMOrganic/MicrobeInteractions E.L.Shock

61

Exploring the sources of the salinity in the Middle East: an integrative hydrologic, geochemical and isotopic study of the Jordan River A. Vengosh, E.Farber, U.Shavit, R.Holtzman, M.Sega1, LGavrieli, ECO-Research Team & Th.D.Bullen

71

Geochemical cycles, global change and natural hazards The chemistry of rainwater in the Mt. Etna area (Italy): sources of major species A.Aiuppa, P.Bonfanti & K D 'Alessandro

83

Hydrothermal systems as indicators of paleoclimate: an example from the Great Basin, Western North America G.B.Arehart & S.R.Poulson

87

Temporal variations of 3He/4Heratios of dissolved helium in groundwaters of Mt Etna, Southern Italy K D 'Alessandro, F.Parello, B.Parisi, P.Allard & P.Jean-Baptiste

91

Magadi and Suguta: the contrasting hydrogeochemistry of two soda lake areas in the Kenya Rift Valley W.G.Darling

95

Geochemical precursors of the 2000 eruption of Mutnovsky Volcano, Kamchatka G.M.Gavrilenko & M. G.Gavrilenko

99

Helium geochemistry applied to crust-mantle interaction in the Apennines (Italy) F.Italiano, M.Martelli & P.M.Nuccio

103

Water-rock interactions during seismic an volcanic activity recorded at Mount Etna by continuous groundwater monitoring F'. Quattrocchi, G.Di Stefano, G.Galli, L.Pizzino, P.Scarlato, P.Allard, D.Andronico, D.Condarelli & T.Sgroi

107

V

The Ardea Basin fluid geochemistry, hydrogeology and structural patterns: new insights about the geothermal unrest activity of the Alban Hills quiescent volcano (Rome, Italy) and its geochemical hazard surveillance F. Quattrocchi, G.Galli, L.Pizzino, G.Capelli, D.De Rita, C.Faccenna, R.Funiciello, G.Giordano, D. Goletto, R.Mazza & C.Mancini

111

Response of an artesian well in southern Armenia to the 1400 km distant Izmit earthquake of August 17, 1999 H. Woith, R. Wang, C.Milkereit, JZschau, UMaiwald & A.Pekdeger

115

Discrete and continuous monitoring of groundwaters in the seismic area of the Umbria region (Italy) A. R.Zanzari, A.Martinelli, R. Cioni, M. Guidi, B.Raco, A.Scozzari, F. Quattrocchi, G.GaIIi & C.Mancini

119

Massflows of the subsurface hydrosphere: Global and regional cycles I? P.Zverev

123

Modelling water-rock interaction H2S-controlling reactions in clastic hydrocarbon reservoirs from the Norwegian Shelf and US Gulf Coast P.Aagaard, J.S.Jahren & S.N.Ehrenberg

129

Quantifying recharge of the Ghussein wells using chemical tracers Nizar S. Abu-Jaber

133

Scale versus detail in water-rock investigations 2: Field-scale models of fracture networks in mineral deposits B. R.Berger, R.B. Wanty & M L. Tuttle

137

The role of pressure solution in fracture healing: A multi-scale reaction-flow modeling approach 0.Bildstein & C.I.Steefe1

141

Reactions governing the chemistry of waters interacting with serpentinites J. Bruni, M. Canepa, F.Cipolli, L.Marini, G.Ottonello, M. Vetuschi Zuccolini, G.Chiodini, R. Cioni & A. Longinelli

145

Does plagioclase control the composition of groundwater in the crystalline basement? K.Bucher & IStober

149

Seawater-basalt interaction: field observations and modeling result 0.I? Chudaev, V.B.Kurnosov, 0.KAvchenko & N.A. Chepkaya

153

Arsenic sulphide precipitation in an active geothermal system: reaction path modelling J S . Cleverley, L. G.Benning, B. WMountain & M. C.Gorringe

157

Mineral growth in rocks: interacting stress and kinetics in vein growth, replacement, and water-rock interaction R. C.Fletcher & E. Merino

161

Large-scale hydrothermal dolomitization in the Southern Cantabrian Zone (NW Spain)

165

M. Gaspanirri, T.Bechstadt & M Boni Geochemical modelling of groundwater quality changes during Aquifer Storage and Recovery (ASR) in the dual porosity Chalk aquifer, England I. Gaus, P.Shand, I.N. Gale & J.Eastwood

169

Molecular dynamics simulation of the uranyl ion near quartz surfaces J.A. Greathouse, G.Bemis & R. T.Pabalan

173

VI

Relative importance of physical and geochemical processes affecting solute distributions in a clay aquitard G.A.Harrington, A.J.Love & A.L.Herczeg

177

Chemistry-transport coupled modelling of the Asp0 groundwater system (Sweden) since the last glaciation W.Kloppmann, D. Thie‘ry, C.Kervkvan, A.Bourguignon, P.Ne‘grel & J. Casanova

181

The Maqarin natural analogue study of a cement-buffered hyperalkaline groundwater plume: structural model and flow systems U K.Mader, M AdIer, V.Langer, P.Degnan, A. E. Milodowski, J.A. T. Smellie, E.Salameh, H.N Khoury, L. Y.Griffault & L. Trotignon

185

Water quality changes during aquifer storage recovery in limestone-silicate aquifer material JMirecki, M.D.Petkewich, K.J.Conlon, & B. G.Campbell

189

Analytical model for deep well injection of cold brine into a hot aquifer A.F.Moench & Y.K.Kharaku

193

Examination of the effect of uncertainty in thermodynamic and kinetic data on computer simulations of complex systems C.H.Moore

197

Evolution and self-organization of the water-rock system S.L.Shvartsev

20 1

Isotopic and chemical characteristics of old “ice age” groundwater, North Iceland A .E.Sveinbjornsddttir, S.Arn6rsson & J.Heinemeier

205

Supercritical water-rock interaction for development of deep-seated geothermal reservoirs N. Tsuchiya, N.Hirano, G.Bignull& K.Nakatsuka

209

Porewater geochemistry and modeling within Oligocene-Miocene clays of North Central Spain M.J. Turrero, J. PeAa, A.M. Ferndndez, P. G6mez & A. Garral6n

213

pH calculation through the use of alkalinity in geochemical modeling of hydrothermal systems M P. Verma & A. H. Truesdell

217

Scale versus detail in water-rock investigations 1: A process-oriented framework for studies of natural systems R.B. Wanty, B.R.Berger & M. L. Turtle

22 1

Thermodynamics, kinetics and experimental geochemistry Equation of state for aqueous non electrolytes N.N.Akin$ev

227

Soultz granite - saline water interactions at 175-200°C and 10-50 bar: experimental and thermo-kinetic modeling approaches M.Azaroua1, V.Plagnes & LMatsunaga

23 1

Solubility and reaction rates of oxides and hydroxides to high temperatures with in situ pH measurement P.Be‘nkzeth,D.A.Palmer, D.J. Wesolowski, C.Xiao & S.A. Wood

235

New insight on the chemical control of aqueous aluminum. Application for modelling water-rock interactions G.Berger & J. P. Toutain

239

Testing a clay/porewater interaction model through a laboratory experiment Ph. Blanc, E. Gaucher, B.Sanjuan, C.Crouzet, A.Seron & L. Grifsault

243

VII

Dissolution of synthetic zeolites at low temperature - preliminary results J. Cama, X Querol, C.Ayora, E.Sanz & J. Ganor

247

Dissolution rate of apophyllite. The effects of pH and implications for underground water storage L. C. Cave!, M. V.Fey & D. K.Nordstrom

25 1

Experimental Study on Mixture Corrosion Effects in Littoral Karst Area, coastal Liaodong Peninsula, China H. Chen, S.Zou & E.Bi

255

Reactivity of pyrite surfaces: Combining XPS and speciation in solution

259

M. Descostes, C.Beaucaire, H.Pitsch, F.Mercier & P. Zuddas Probing the electrical double-layer structure at the rutile-water interface with X-Ray standing waves P.Fenter. L.Cheng, S.Rihs, M.Machesky, M.J.Bedzyk & N. C.Sturchio

263

Enriched stable isotopes for determining the sorbed element fraction in soils in order to calculate sorption isotherms H. -E,Gabler & A.Bahr

267

To stir or not to stir - implications for silicate dissolution experiments JGanor & V.Metz

27 1

The bentonite-water interface and its role in the adsorption processes of metals T.Gavriloaiei

275

Limiting mechanisms of borosilicate glass alteration kinetics: Effect of glass composition S.Gin & C.Je!gou

279

Heavy-metal binding mechanisms in cement minerals C.A.Johnson, I.Baur & F.Ziegler

283

Thermodynamic elucidation of Eu anomalies in REE pattern in hydrothermal fluorite G.R.Kolonin & G.P.Shironosova

287

Elements transfers in compacted clayey materials under thermal gradient C.Latrille, M,Jullien & C.Pozo

29 1

The Chemical Durability of Yttria-Stabilized Zr02 pH and O2 Geothermal Sensors M.FManna, D.E. GrandstafJ; G.C.Ulmer & E.P. Vicenzi

295

Gibbs free energies of formation of uranyl silicates at 298.15 K W.F.McKenzie, L.Richard & S.Salah

299

Negative pressure and water-mineral interaction in the unsaturated zone of soils L.Mercury, P. Freyssinet & Y.Tardy

305

Solubility of Na, Al, and Si in aqueous fluid at 0.8-2.0 GPa and 1000-1300°C B. 0.Mysen & K. Wheeler

309

The surface chemistry of a gram-negative bacteria and its role in metal uptake B. T.Ngwenya & I. W.Sutherland

313

A test of aqueous speciation: Measured vs. calculated free fluoride ion activity D.K.Nordstrom

3 17

Fading of luminescense in feldspars - an autoradiographic method E. Oila, S.Pinnioja, M Siitari-Kauppi, V.Aaltonen & A.Lindberg

32 1

Sampling techniques and pH measurement methods for geochemical analysis of deep groundwaters H.Pitsch, C.Beaucaire, P.Meier & S. Grappin

325

VIII

Revised thermodynamic properties of malachite and azurite W.Preis & H. Gamsjager

329

Thermodynamic calculation of the distribution of organic sulfur compounds in crude oil as a function of temperature, pressure, and H2S figacity L.Richard & H. C.Helgeson

333

Measurement of quartz dissolution rates with a flow-through type autoclave reactor H.Sugita, I. Matsunaga, T.Yamaguchi & H. Tao

337

Experimental study of rock/water/C02 interaction at temperatures of 100-35OoC Y.Suto, L.Liu, T.Hashida, N. Tsuchiya & N. Yamasaki

34 1

The source of sodium in groundwater, Pannonian Basin, Hungary I. Varscinyi & L. 0.Kovcics

345

Silica solubility geothermometers for hydrothermal systems

349

M P. Verma Leaching kinetics of a quartz-chlorite schist and consequent changes in the rock structure T.Wells, P.Binning, G. Willgoose & A.Mews

353

Rate of mineral dissolution during granite-hydrothermal alteration P.Zuddas & F.Seimbille

357

Mineral surfaces and weathering Glauconite Dissolution Rates and the Chemical Evolution of Vadose Waters in the Homerstown Formation, Homerstown, New Jersey J.Betts & D.E. Grandstaff

363

Mineralogical evolution of bituminous mar1 adjacent to an alkaline water conducting feature at the Maqarin analogue site A. Cassagnabire, J. C.Parneix, S.Sammartino, L. Y Griffault, U Maeder & T.Milodowski

367

Oxidation of an argillaceous formation: mineralogical and geochemical evolution D. Charpentier, M. Cuthelineau, R.Mosser-Ruck & G.Bruno

371

Microscopic processes at the interface between metal sulphides and water G.De Giudici & P. Zuddas

375

Dissolution of calcite in CaC03-C02-H20 systems in porous media P.A.Diaz, KAlvarado & M.I. Rodriguez

379

Composition of charnockite weathering products in three climatic zones W.I S .Fernando, R.Kitagawa, B.P.Roser, Y.Hayasaka & Y.Takuhashi

383

Bascplica da Estrela stone decay: the role of rain-water C.A.M.Figueiredo, A.A. Mauricio & L.Aires-Baffos

387

A new model of rock weathering: design and validation on a small granitic catchment L. Franqois, A.Probst, Y.Godde'ris,JSchott, D.Rasse, D. Viville, 0.Pokrovsky & B.Dupre'

391

Characteristics of smectites from nickeliferous laterite in Australia A. Gaudin, Y.Noack, A.Decarreau & S.Petit

395

Surface area vs mass - which is most important during mineral weathering in soils? M.E.Hodson

399

Geochemistry of a profile at the weathering front in dolomite H.B.Ji, S.J. Wang, 2.Y.Ouyang, C.Q.Liu, C.XSun & XM.Liu

403

The use of U-isotopes on the study of a weathered cover in ParanB basin, Brazil J.R.Jime'nez-Rueda & D.M.Bonotto

407

IX

Water-Rock Interaction and the Water Chemistry of a Small Sierra Nevada Lake D.S.Kimba1 & R. W.Smith

41 1

The S.Antioco of Bisarcio Basilica (NE Sardinia, Italy): water-rock interaction in ignimbrite monument decay G.Macciotta, G.Bertorino, A. Caredda, S.Columbu, M Franceschelli, M Marchi, S.Rescic & R. Coroneo

415

Characterization of oxidation products onto pyrite: coupling of XPS and NMA F.Mercier, M. Descostes, C.Beaucaire, P. Trocellier & P.Zuddas

419

Local structure of uranium (V 1) sorbed on clinoptilolite and montmorillonite R.J.Reeder, M Nugent & R. T.Pabalan

423

Surface composition of enargite(CU3AsS4) A.Rossi. D.Atzei, B.Elsener, S.Da Pelo, F.Frau, P.Lattanzi, P.L. Wincott & D.J. Vaughan

427

Orthoclase surface structure and dissolution measured in situ by X-ray reflectivity and atomic force microscopy N.C.Sturchio, P.Fenter, L. Cheng & H. Teng

43 1

Disseminated calcite in a global suite of granitic rocks: Correlations with experimental solutes A.F. White, MS.Schulz, D. V. Vivit & Th.D.Bullen

43 5

Aqueous dissolution studies of synthetic and natural brannerites Y.Zhang, G.R. Lumpkin, B.S. Thomas, Z.Aly, R.A. Day, K.P.Hart & M Carter

439

Groundwater environments Shallow groundwater in the Sebou basin (Northern Morocco) T.Bahaj, M El Wartiti, M. Zaharaoui, R. Caboi & R. Cidu

445

Pore waters in Mesozoic mudrocks in southern England A.H.Bath

449

Groundwater in the urban area of Catania (Sicily, Italy). Geochemical features and human-induced alterations M.Battaglia & P.Bonfanti

453

Study on water quality in the area of Wadi Shueib, Jordan Valley, Jordan K.Becker, W.Ali & HHoetzl

457

Chemical evolution of ground waters in W-Iceland (SnEfeIlsnes) E.Bedbur, M.Petersen, H.Biallas, U. Wollschlager & S.Schmidt

46 1

Hydrogeochemistry in the Flumendosa river basin (Sardinia, Italy) R. Caboi, A. Cristini, M. Collu, F.Podda & L.Rundeddu

465

Hydrogeochemical characteristics of surface water and groundwater in areas underlain by black shales and slates of the Okchon zone, Korea H.-T.Chon & S.Y.Oh

469

New geochemical data of the high PCO2 waters of Primorye (Far East Russia) 0.V.Chudaev, V.A.Chudaeva, K.Sugimori, K.Nagao, B. Takuno, M Matsuo, A.Kuno & M.Kusakabe

473

Origin of fluorine within the Afyon-Isparta volcanic district, SW Turkey: is fluormica the key? HCoban, $.Caran & M.Gormiig

477

Groundwater composition of perched-water bodies at Azores volcanic islands J. V.Cruz & Z.M. Franqa

48 1

X

Groundwater circulation at Mt. Etna: evidences from "0, 2H and 3H contents W.D 'Alessandro, C.Federico, A.Aiuppa, M.Longo, F.Parello, P.Allard & P.Jean-Baptiste

485

Groundwater geochemistry in the Broken Hill region, Australia P. de Caritut, N.Lavitt & D.Kirste

489

The mineralised springs of the Marche and Abruzzi foredeep, central Italy: hydrochemical and tectonic features G.Desiderio, S.Rusi, T,Nanni & P. Vivalda

493

Water-rock interaction in a karstified limestone sequence, south Galala, Gulf of Suez, Egypt A A E l - F i b , M.N.Shaaban & M.A.Rashed

497

Hydrogeochemical characteristics of Hummar aquifer in Amman-Zarqa basin, Jordan A.R.EL-Naqa & K.M.Ibrahim

501

Geochemical characterization of groundwaters from the Hyblean aquifers, South-Eastern Sicily R.Favara, F. Grassa & M, Valenza

505

Monitoring of groundwater quality in Umbria (Central Italy) F.Frondini, G.Marchetti, A. Martinelli, L.Peruzzi & R. Crea

509

Salt water intrusion in the Pisa coastal plain (central Italy) F.Frondini, A.Zanzari & S. Giaquinto

513

Salinization in coastal plain of Grosseto: hydrochemical study E. G.Forcada, A.Bencini & G.Pranzini

517

Biogeochemical cycles of chloride, nitrogen, sulphate and iron in a phreatic aquifer system in The Netherlands J. Griffioen & Th.Keijzer

52 1

Magnesium concentration control in groundwaters in Iceland I. Gunnarson, S.Arndrsson & S.Jakobsson

525

Can major ion chemistry be used to estimate groundwater residence time in basaltic aquifers? A.L.Herczeg

529

Water chemistry at Snowshoe Mountain, Colorado: mixed processes in a common bedrock A.R.Hoch & M. M. Reddy

533

Evidence for brine circulation in a groundwater discharge zone

537

I?. M.Howes, C.Le Gal La Salle & A.L.Herczeg Origin of sodium-bicarbonate waters in the south-eastern part of the Great Artesian Basin: Influx of magmatic CO2 J.Jankowski & W McLean

541

Chemical evolution of groundwater in the Tularosa Basin in Southern New Mexico, USA Th.G.Kretzschmar, D.Schulze-Makuch & I.S. Torres-Alvarado

545

Investigation of the carbonate system in Aquifer Storage and Recovery: an isotopic approach C.Le Gal La Salle, J. Vanderzulm, J. Hutson, P.Dillon, P.Pavelic & R. Martin

549

Hydrogeochemical evolution of karst water system: A case study at Niangziguan Springs, northern China Y L i & Y. Wang

553

Exchange of solutes between primary and secondary porosity in a fractured rock aquifer induced by a change in land-use A.J.Love & A. L.Herczeg

557

XI

Redox chemistry of a river-recharged aquifer in the "Oderbruch" region in eastern Germany GMassmann, A.Pekdeger, C.Merz & M -Th.Schafmeister

56 1

The origin of Na-HC03 type groundwater in an eastern section of the Lower Namoi River catchment, New South Wales, Australia KMcLean, J.Jankowski & N.Lavitt

565

Mineralised waters and deep circulations in the French-Italian Alps J.P.Novel, G.M Zuppi, M.Dray, S.Fudral, G.Nicoud & P.Lacombe

5 69

Chemical and isotopic signatures of interstitial water in the French Chalk aquifer land water-rock interactions C.Plain, L.Dever, C.Marlin & E. Gibert

573

Water-rock interaction processes in the main thermal springs of Sardini (Italy) M Proto, C.Panichi, P.Zuddas & F.Podda

5 77

Arsenic and other redox-sensitive elements in groundwater from the Huhhot Basin, Inner Mongolia P.L.Smedley, M Zhang, G.Zhang & 2.Luo

58 1

Water-Rock reactions in a deep barite-fluorite underground mine, Black Forest, Germany I.Stober, KZhu & K.Bucher

585

Hydrologic controls on groundwater salinisation, Murray Basin, Australia I.P. Swane, T.R. Weaver, C.R.Lawrence & I. Cartwright

589

Sulfide-free and sulfide-bearing waters in the Northern Apennines, Italy L. Toscani & G. Venturelli

593

Elevation, landuse and water-rock interaction effects on groundwater quality S. Tweed, T.R. Weaver, G.P.Masur & I. Cartwright

597

Hydrochemical patterns of the Gavarres hydrological system and its surrounding aquifers (NE Spain) E. Vilanova & J.Mas-Pla

60 1

Hydrogeochemistry of shallow groundwaters from the nortkrn part of the Datong basin, China R. Wang, Y. Wang & H. Guo

605

Decoupling solute distributions from groundwater flow in low permeability media T.R. Weaver, S.K.Frape & J.A. Cherry

609

Sedimentary basins Fluid-sediment interaction and clay authigenesis along the flank of the Juan de Fuca Ridge M.B.Buatier, MSteinmann, C.Bertrand, A.M Karpo#& G.L.Fruh-Green

615

Hyperkarstic phenomena in the Iglesiente mining district (SW-Sardinia)

619

J.De Waele, P.Forti & G.Pema Diagenetic zeolite and clay minerals in Miocene Great Bahama Bank carbonate sediments (ODP Leg 166, Site 1007) A.M.Karpog S.M.Bernasconi, C.Destrigneville & P.Stille

623

Numerical study of the coupling effect between fluid diffusion and medium deformation for subsidence calculation over deep reservoirs G.Lecca, R.Deidda & G. Gambolati

627

Squeegee flow in Devonian carbonate aquifers in Alberta, Canada H. G.Mache1, B.E. Buschkuehle & K.Michae1

63 1

Flat lowland paleogeography of sediment-collecting basins: Evidence from formation waters E.Mazor

635

XI1

The influence of basement fluid upwelling and diagenesis on CaC03 stability in sediments from the eastern flank of the Juan de Fuca ridge C.Monnin, C.G. Wheat, M.M.Motti1 & S.Balleur

639

Fluid flow in the Cantabrian Zone (NW-Spain) - contributions to the diagenetic evolution JSchneider, T.Bechstadt, S.Zeeh & M Joachimski

643

Sulfate reduction rates and organic matter composition in sediments off Namibia C.J.Schubert, T.G.Ferdelman, B.B.Jmgensen & G.Klockgether

647

Origin of Ordovician organogenic dolomite concretions: Significance for the 6l80 of Lower Paleozoic SMOW J.Shah, R.Hesse & S.Islam

65 1

Illite crystallinity an expandability: XRD and HRTEM studies of Gaspe Peninsula mudstones and slates S.Shata & R.Hesse

655

H2S in North Sea oil fields: importance of thermochemical sulphate reduction in clastic reservoirs R.H. Worden & P.C.Smalley

659

Magmatic, metamorphic and minerogenetic processes Metasomatic reaction bands - a key to component mobility at metamorphic conditions

665

RAbart Petrology and alteration of basalts from the intraplate rises, Indian Ocean A. I/:Artamonos, V.B. Kurnosos & B.P.Zolotarev

669

Multiple fluid-flow events and mineralizations in SW Sardinia: an European perspective M.Boni, A.Iannace, I.M. Villa, L.Fedele & R.Bodnar

673

High-pressure melting and fluid flow during the Petermann Orogeny, central Australia I.S.Buick, D. Close, I.Scrimgeour, C.Edgoose, J.Miller, C.Harris & I. Cartwright

677

Natural zeolites from Cenozoic pyroclastic flows of Sardinia (Italy): evidence of different minerogenetic processes P. Cappelletti, G.Cerri, M.de 'Gennaro, A.Langella, S.Naitza, G.Padalino, R.Rizzo & M Palomba

68 1

Amphibole evolution in ultramafic amphibolites from NE Sardinia, Italy A.M Caredda, G.Cruciani, M Franceschelli & G.Carcangiu

685

Fluid geochemistry of Tieluping Ag ore and its implications for the CPMF model Y.J.Chen, Y.H.Sui &XL.Gao

689

Water-rock interaction in genesis of perlite at Monte Arci volcanic complex (West Sardinia, Italy) R. Cioni, G.Macciotta, M.Marchi, G.Padalino, R.Simeone & M Palomba

693

Trace element mobility in tourmalinite veins and surrounding metapelites from the Crummock Water aureole (Lake District, England) C.Corteel & N.J.Fortey

697

Gold ore-system in Arhcean greenstone structures of Middle-Dniper Area (Ukrainian Shield) Yu.Fomin, He.Lasarenko, Yu.Demikhos & Vl.Blazhko

70 1

Late Hercynian fluid circulation in the Charroux-Civray plutonic complex, NW Massif Central, France R.Freiberger, M -C. Boiron, M. Cathelineau & M. Cuney

705

Experimental study on clinoptilolite and mordenite crystallization M. R. Ghiara, C.Petti & R.Lonis

709

XI11

Ore fluid, of late Mesozoic porphyry-epithermal gold-copper system in East China R.Hua, X Li, J.Lu, P. Chen & X Liu

713

The characteristics and genesis of the kaolinite-bearing gold-rich Nurukawa Kuroko deposit, Aomori Prefecture, Japan D. Ishiyama, K.Hirose, TMizuta, 0.Matsubaya & Y.Ishikawa

717

Sea water-basalt interaction in the Kerguelen Plateau, Indian Ocean K B.Kurnosov, B.P.Zolotarev & A. KArtamonov

72 1

Magmatic versus hydrothermal processes in the formation of raw ceramic material deposits in southern Tuscany P.Lattanzi, M Benvenuti, P. Costagliola, C.Maineri, I. Mascaro, G.Tanelli, A.Dini & G.Ruggieri

725

Occurrence of halite in kaolin of NW Sardinia: genetic implications P.Mame1i

729

Interaction of twinning structure of the feldspars with water fluid - the most significant geological process in the Earth’s crust KS.Melnikov

733

Fossil geothermal systems in the continental rift zone of the Kiigiik Menderes within the Menderes Massif, Western Anatolia, Turkey N. Ozgzir

737

Gold in Sardinia: recent developments in exploration and exploitation J.Rayner & D.Manis

74 1

Hydrothermal mineralization of Zr and other “immobile elements”: field evidence and experimental constraints S.Salvi. B. Tagirov & B.Moine

745

Time-depth-temperature relations for igneous, metamorphic and hydrothermal processes: Visualized. through simplified-model numerical simulations H.Shigeno

749

Mass transfer, oxygen isotopic variation and gold precipitation in epithermal systems: a case study of the Hishikari deposit, southern Kyushu, Japan NShikazono, N.Yonehzwa & T.Karakizawa

753

Overprinted Cenozoic hydrothermal activities at the Toyoha Ag-Pb-Zn deposit, Japan TShimizu & A.Aoki

757

Natural decay series studies of the Kujieertai uranium deposit, NW China S.Zhanxue. L.Jinhui, L.Xueli & S. Weijun

76 1

Thermodynamic framework of the contact metamorphism around the Kakkonda granite in an active geothermal field, northeast Japan N. Takeno, H.Muraohz, T.Sawaki & M.Sasaki

765

Interaction of fluid inclusions with dislocations in quartz N.A. Tchepkaia & ZA.Kotelnikova

769

XIV

PREFACE

The 1Oth International Symposium on Water-Rock Interaction (WRI-10), sponsored by the International Association of Geochemistry and Cosmochemistry (IAGC), the Italian National Research Council (CNR), the University of Cagliari and the Societa Geochimica Italiana, was held in Villasimius, Italy, June 10-15, 200 1. Some 400 manuscripts were submitted by scientists from 45 countries for presentation in both oral and poster sessions of this symposium. Following reviews, 380 of these papers were accepted and included in these two volumes. The published papers describe the results of latest research on water-rock interactions in different geochemical environments ranging from surface and ground water systems and sedimentary basins to magmatic and geothermal systems. Many papers report the application of advanced methodologies, including isotopes, geochemical codes and analytical techniques. Then, there is an increasing interest in integrated studies of the water-rock systems. Although the environments covered are often interrelated, the papers were divided into 14 major topics listed below:

I . Geochemical cycles, global change and natural hazards 2. Modelling water-rock interaction 3. Thermodynamics, kinetics and experimental geochemistry 4. Mineral surfaces and weathering 5. Groundwater environments 6. Sedimentary basins 7. Magmatic, metamorphic and minerogenetic processes 8. Volcanic and geothermal processes 9. Trace element mobility 10. Pollution and remediation :general issues 11. Pollution and remediation :mining environments 12. Waste storage and disposal 13. Biogeochemical processes and organic complexation 14. Stable and radiogenic isotopes in WRI studies It can be noted that the number of papers for each topic may vary considerably. Though the traditional fields of previous WRI symposia, such as Modelling WRI / Thermodynamics / Weathering / Magmatic and Geothermal processes / Isotope studies, continue to attract many scientists, an increasing number of papers on Waste storage and disposal / Biogeochemical processes has been received. It is worthy of note that more than 100 papers deal with problems related with the quality of waters both in developed and developing countries, either in natural or contaminated environments. These papers have been distributed in the Groundwater Environments and Pollution and Remediation sessions, but their assignment was not always easy and overlapping may occur in some cases. The WRI-10 has benefited from the participation of eight invited Keynote Speakers who are the world’s foremost leaders in their fields, their papers highlight the recent advancements in WRI studies; the summary of their presentations are printed at the beginning of Volume 1. XV

All the submitted manuscripts were reviewed by at least one referee. Minor editorial modifications and correction of typographical errors were made by the Editor on the original manuscripts, but due to the large number of papers most of them were returned to authors with suggested modifications. A few papers needed extensive reorganisation and rewriting and were hopefully improved after consultations with Authors. These proceedings have been improved thanks to our friends and colleagues who devoted their time in careful reviewing all manuscripts submitted to WRI-10. Despite of the time constraints, they have been of considerable help, especially in recovering some of the original manuscripts; we are greatly indebted to them, it would have not been possible to complete the editing of theWRI- 10 proceedings without their kind collaboration. In addition to the Editor, reviewers of manuscripts were: H. Armannsson S. Arnorsson C. Beaucaire J.O. Bjarnason T.D. Bullen R. Caboi A.M. Caredda Raffaello Cioni P. Costagliola W.G. Darling G.B. De Giudici C. de Ronde A. Dini W.M. Edmunds S. Einarsson W. Evans L. Fanfani

I. Mascaro G. Massoth C.J. Milne C. Moore J. 6lafsson M. Olafsson H. Oskarsson N. Oskarsson G. Ottonello T. Paces C. Panichi H. Pitsch A. Reyes B. Robinson M. Rosen G. Ruggieri M.S. Schulz

K. Faure H. Franzson M. Franceschelli F. Frau G.O. Fridleifsson G. Gislason S.R. Gislason S. Hauksdbttir G. ivarsson Y. Kharaka D.G. Kinniburgh H. Kristmannsdottir P. Lattanzi G. Lyon C. Maineri L. Marini V. Marteinsson

P. Shand D. Sheppard R. Simeone P.L. Smedley A.E.Sveinbjomsdottir J. Thordsen 0. Vaselli D.V. Vivit R.B. Wanty J. Webster P. White C.P. Wood P. Zuddas Pierpaolo Zuddas

Authors have been encouraged to send their manuscripts by electronic submission and in fact most manuscripts were received in this form, so contributing to speed reviewing and editing. However, many manuscripts arrived in somewhat less than the required format. We wish a special thank to Laura Rundeddu, Stefania Da Pelo and Franco Frau for spending a great deal of time, and weekends, in carrying the tedious job of formatting and assembling the manuscripts. We are also indebted to Carla Ardau, Mario Lorrai, Riccardo Biddau, Claudia Dadea, and Francesca Podda for their valuable collaboration. Also thanks to Corsi & Congressi for addressing the myriad details of registration and logistics. Finally, we hope that our collective efforts have resulted in an improved Proceedings. The papers in these two volumes represent the latest results of investigations in the field of water-rock interactions and should be of interest to a large number of scientists working in hydrogeochemistry and geochemistry of natural and contaminated systems. The reader hopefully will find stimulating ideas that will contribute to expand earthscience education, as well as strengthen research activities for solving the environmental problems that face our society today.

Luca Fanfani Secretary General, WRI- 10

Rosa Cidu Editor, WRI- 10 Proceedings

XVI

ORGANISATION

INTERNATIONAL ASSOCIATION OF GEOCHEMISTRY AND COSMOCHEMISTRY (IAGC) EXECUTIVE COMMITTEE (IAGC) President: Eric M. Galimov, Russia Vice-President: John Ludden, France Secretary: Me1 Gascoyne, Canada Treasurer: David Long, USA Past-President: Gunter Faure, USA COUNCIL MEMBERS: Attila Demeny, Hungary John J. Gurney, South Africa Russel S. Harmon, USA Hochen Hoefs, Germany Marc Javoy, France Jan Kramers, Switzerland Gero Kurat, Austria N.V. Sobolev, Russia K.V. Subbarao, India Yishan Zeng, China

WORKING GROUP ON WATER-ROCK INTERACTION EXECUTIVE COMMITTEE Chairman: W.M. (Mike) Edmunds, UK MEMBERS: Tomas Paces, Czech Republic Yves Tardy, France Brian Hitchon, Canada Hitoshi Sakai, Japan Halld6r Armannsson, Iceland Yousif Kharaka, USA Oleg Chudaev, Russia Brian Robinson, New Zealand Luca Fanfani, Italy

XVII

WRI- 10 ORGANISING COMMITTEE SECRETARY GENERAL Prof. Luca Fanfani, Department of Earth Sciences, University of Cagliari SCIENTIFIC PROGRAMME Prof. Rafaele Caboi, University of Cagliari Prof. Pierfranco Lattanzi, University of Cagliari Prof. Antonio Longinelli, University of Parma Prof. Giulio Ottonello, University of Genova Dr. Costanzo Panichi, IIRG-CNR Pisa Prof. Paola Zuddas, University of Cagliari EXECUTIVE PROGRAMME Dr. Rosa Cidu, University of Cagliari Dr. Franco Frau, University of Cagliari FIELD TRIPS Prof. Pierfranco Lattanzi, University of Cagliari Prof. Mariano Valenza, University of Palerrno Prof. Giovanni Orsi, University of Naples ORGANIZING SECRETARIAT Corsi & Congressi, via Ghibli 8, 09126 Cagliari THE ORGANISING COMMITTEE ARE GRATEFUL TO THE MAJOR SPONSORS WHO ARE SUPPORTING THE WRI-10 SYMPOSIUM: Consiglio Nazionale delle Ricerche (CNR) Universita di Cagliari ENEL Distribuzione Sardegna Ente Sardo Industrie Turistiche (ESIT) Parco Scientific0 e Tecnologico della Sardegna Saras S.p.A Raffinerie Sarde

XVIII

Keynote lectures I invited speakers

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Water-Rock Interaction 2007, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Water-rock interactions in mudrocks and similar low permeability material A .H .Bath Intellisci, Loughborough, LE1 2 6SZ, England

F.J .Pearson Ground-Water Geochernistry, New Bern, NC 28560, USA

A .Gautschi NAGRA, CH-5430 Wettingen, Switzerland

H .N.Waber GGWW, Geol & Min-Pet Institute, University of Bern, CH-3012 Bern, Switzerland

ABSTRACT: Mudrocks have sub-micron pores in which water and solutes interact with mineral surfaces. Pore waters move very slowly, if at all, and mass transfers of solutes are oRen controlled by diffusion. Recent progress in investigating and understanding water-rock interactions in these systems is reviewed. Pore water sampling and analyses, mineral surface-water-soluteinteractions, and geochemical equilibrium modelling are some of the challenging aspects of these studies. A set of papers describing work at the Mont Terri Underground Rock Laboratory in Switzerland is introduced. Pore water compositions have been measured with an improving degree of reliability. Non-reactive solutes demonstrate that diffusion controls water and solute movements across the Opalinus Clay formation at Mont Terri. Progress has been made in understanding the reactions that control geochemical conditions with a combination of in situ data, experimentation and modelling. The findings have implications in various areas of geochemistry and hydrogeology 1 INTRODUCTIONAND SCOPE

and porosities from 0.4 to 35% of the particles being 4 is a strongly uni-directional process with little to no tendency for back-reaction. Fe(I1)-oxidation involves conversion of Fe(II)(OH),-aqueous species through a series of short-lived Fe(I1)-transition complexes and Fe(II1)-aqueous intermediates prior to precipitation as ferrihydrite (Millero, 1985). In the absence of other controlling factors, basic kinetic reaction theory predicts that the most reactive Fe(II)(OH),-species should be those that contain the lighter Fe isotopes due to their inherently weaker bonds (Urey, 1947). Therefore, the fact that the Fe in the ferrihydrite is consistently heavier than the

P

24

Figure 3 . Results of steady-state ferrihydrite precipitation experiments. a) Fe isotope fractionation factor a versus the proportion of the Fe(I1) aqueous species Fe(II)HC03+(aq) relative to Fe(II)(aq)total.Two different gas mixtures (5% CO2-95% air, 1% C02-99%air) were used to establish CO2 saturation. Error bars show maximum standard deviation based on replicate measurements. b) a values for experiments at pH=6.0 in which SO4 concentrations were varied. Arrow shows direction of decreasing Fe(I1) oxidation rate. coexisting aqueous Fe under both field and experimental conditions is unexpected, and requires one or more additional mechanisms such as equilibrium fractionation of Fe isotopes among coexisting Fe(I1)-aqueous species including transition complexes. The best explanation of the results of this paired field and laboratory study is that Fe comprising the Fe(I1)-aqueous species Fe(II)(OH),(aq) (including transition complexes), Fe(II)HC03+(aq) and Fe(II)S04 aq is isotopically heavy relative to that in the Fe(I1) '(as) pool. In addition, Fe(II)HC03+(aq)

0

constitutes a reactive Fe pool that apparently enhances Fe(I1) oxidation. In contrast, Fe(II)S04(aq) constitutes a non-reactive Fe pool that apparently supresses Fe(I1) oxidation. The results likewise suggest that a multi-tracer approach may allow successful application of Fe isotopes as a "biosignature", at least for systems dominated by ferrihydrite or the products of its recrystallization. For example, it is reasonable to assume that Fe in ferrihydrite produced abiotically in nature will be heavier than the aqueous Fe from which it forms. Along a reaction flowpath (e.g., along a stream bed), the Fe in the ferrihydrite should become progressively lighter downgradient as heavy Fe is preferentially removed from the water. In contrast, available data suggests that microbiallymediated iron reduction preferentially mobilizes light Fe from ferrihydrite substrates (Beard et al. 1999). Assuming that the mobilized ferrous iron ultimately re-oxidizes and forms a progressive accumulation of ferrihydrite or a similar mineral, Fe in each sequential deposit should be heavier with time as the microbes progressively deplete light Fe from the original Fe source. If the directional sense of either reaction flowpath or mineral accumulation

25

sequence can be determined by independent means, then the trend of Fe isotope compositions along that direction should differ for biotic and abiotic mechanisms. Potential determinants might include stratigraphic direction indicators or ratios of various trace metals that are incorporated differently during mineral precipitation. An alternative approach may be to consider together both the Fe and oxygen isoto e s stematics y would of Fe-oxyhydroxide minerals. The 6' 80pratio directly reflect the source of the oxygen, which could vary significantly between H2O and molecular 0 2 depending on whether the mineral formed biotically or abiotically (Mandernack et a1 1999). Given a sufficent number of well-constrained data, it is possible that biotic and abiotic mineral formation processes would describe different trends on a plot of 6l80 versus 656Fe. Although there currently is work in progress to assess the potential of this method, there are no data presently available to judge its merits.

5 USE OF CARBON ISOTOPES TO DISTINGUISH BIOGENIC AND ABIOGENIC METHANE The carbon (C) isotopes have proven to be especially useful in studies of water-rock interaction, and in contrast to the Fe isotope system the fractionation mechanisms in nature are relatively well understood. As noted by Hoefs (1987), the distribution of oxidized C-compounds in inorganic systems and reduced C-compounds in the biosphere is an ideal situation for the development of stable isotope fractionations. Naturally occurring variations of carbon isotope composition (reported as 6 1 3 C p ~ ~ ,

Regardless of fractionation mechanism, it is clear from these experiments that the use of 6l3CpDB on its own to confirm a biogenic origin for CH4 is now problematic at best. This presents another case in which a multi-tracer approach must be applied. Unfortunately, Horita & Berndt (1999) point out that other potential discriminants such as CH4/(CzHs + C3Hs) ratios of the gases produced in the hydrothermal experiments were similar to those in gases of microbial origin. Alternatively, it is possible that purely microbial processes result in different reaction intermediates that might fractionate the hydrogen isotopes differently than those produced during the abiotic reaction. If this proves to be the case, then different trends for biogenic and abiogenic CH4 formation might develop in plots of 6I3cpDBversus 6 2 (cf., ~ Witicar et al. 1986).

the per mil deviation of 13C/12Cof a sample from that of a standard belemnite collected from the Peedee Formation) span a range of greater than 100%0, with heavy carbonates and light methane (CH4) defining the extremes. The variations have been clearly documented to result from both equilibrium and non-equilibrium isotope effects. CHq in the shallow crust and biosphere is formed primarily as a product of either the digestion of organic compounds by microorganisms (i.e., microbial CH4) or the thermal decomposition of organic matter (i.e., thermogenic CH4). Because abiogenic CH4 formation is prohibitively slow in the absence of catalysts at low to moderate temperatures even under reducing conditions, most CH4 generated under such conditions has generally been presumed to be biogenic (Horita & Berndt 1999). On the other hand, there is some evidence supporting the production of abiogenic CH4 in the earth’s crust and upper mantle. The few reported cases generally describe CH4 associated with Fe-Ni rich mantlederived rocks (e.g., Rona et al. 1992, Abrajano et al. 1990, Kelley 1996). However, distinguishing abiogenic from biogenic CH4 is at best difficult. Lacking other reliable indicators a logical geochemical criterion for identifying abiogenic CH4 would be 6l3CpDB > -25%0, a value that is slightly greater than that of the heaviest demonstrably biogenic CH4 sample yet analyzed (Jenden et al. 1993). This is significant because 6l3CpDBof CH4 is widely used as a tool for identifying its origin and mechanisms of formation. Recognizing the significance of the spatial association of possible abiogenic CH4 with ultramafic rocks and in light of experiments that demonstrated that dissolved HC03- can be converted to CH4 in the presence of ultramafic rocks (e.g., Janecky & Seyfried, 1986), Horita and Berndt (1 999) performed controlled experiments to produce CH4 from dissolved HCO3- in the presence of Fe-Ni alloy under hydrothermal conditions. Their isotopic measurements of the experimental products revealed that the fractionation factor aCH4-CO2 ranged from 0.940 to 0.965, an essentially identical range to that resulting from microbial reduction of CO2 to CH4 at ambient temperature ( a ~ ~ 4 - ~=0 0.930 2 to 0.960; Botz et al. 1996). These a values are considerably less than those predicted for equilibrium fractionation between CH4 and dissolved HC03(i.e., 0.970 to 0.982 at the temperatures of the hydrothermal experiments). Therefore, Horita and Berndt (1999) attributed the light C isotope composition of the product CH4 to strong disequilibrium “kinetic carbon isotope fractionation”. Based on C mass balance in the experiments, they postulated the formation of a reaction “intermediate” such as formate ion that might provide control of the rate-limiting step of the observed isotope fractionation.

6 CONCLUSIONS These examples clearly demonstrate the potential risk of relying on a single isotope or other geochemical tracer alone for process identification and interpretation. In general, the use of a multitracer approach will considerably reduce ambiguity and is recommended. Because of the broad applicability of Sr, Fe and C isotopes to water-rock interaction issues, the possibilities for their misinterpretation are limitless and their correct application essentially requires a multi-tracer approach. Identifying effective multi-tracer methods will be easier in some cases than in others. For example, Sr isotope data can be constrained using a variety of other isotopic and geochemical parameters. It may prove more difficult to identify additional tracers that can be used to constrain Fe and C isotope data, particularly for the types of applications presented above. The multi-tracer approaches suggested above for the Fe and C isotope systems (Le., consider together aS6Fe and 6l80 of Fe-oxyhydroxide minerals, 613C and 62H of CH4) have reasonable chances for success, but will require considerable further field and laboratory studies. We feel strongly that isotopes are most powerful when used to constrain hypotheses based on other geologic, hydrologic and biogeochemical information. Regardless, by constantly assessing the validity of the isotope and other geochemical tracers in our toolbox, we invariably further our understanding of the way that isotopes work and realize new directions for their application.

26

REFERENCES Abrajano, T.A., N.C. Sturchio, B.M. Kennedy, G.L. Lyon, K. Muehlenbachs & J.K. Bohlke 1990. Geochemistry of reduced gas related to serpentinization of the Zambales Ophiolite. Appl. Geochem. 5 : 625-630. Anbar, A.D., J.E. Roe, J. Barling & K.H. Nealson 2000. Nonbiological fractionation of iron isotopes. Science 288: 126-128. Beard, B.L. & C.M. Johnson 1999. High precision iron isotope measurements of terrestrial and lunar materials. Geochim. Cosmochim. Actu 63: 1653-1660. Beard, B.L., C.M. Johnson, L.Cox, H. Sun, K. Nealson & C. Aguilar 1999. Iron isotope biosignatures. Science 285: 1889-1892. Beard, B.L., C.M. Johnson, J.L. Skulan & J. O'Leary 2000. Fe isotope fractionation in nature: when is it bugs, when is it not? EOS 81,48: F195. Blum, J.D., C.A. Gazis, A.D. Jacobson & C.P. Chamberlain 1998. Carbonate versus silicate weathering in the Raikhot watershed within the High Himalayan Crystalline series. Geology 164: 41 1-414. Blum, J.D. & Y. Erel 1997. Rb-Sr isotope systematics of a granitic soil chronosequence: the importance of biotite weathering. Geochim. Cosrnochim.Acta 61/15: 3 193-3204. Blum, J.D., Y. Erel & K. Brown 1994. s7Sr/s6Sr ratios of Sierra Nevada streamwaters: implications for relative mineral weathering rates. Geochim. Cosmochim. Acta 58: 5019-5025. Botz, R., H.D. Pokojski, M. Schmitt & M. Thomm 1996. Carbon isotope fractionation during bacterial methanogenesis by CO2 reduction. Org. Geochem 25: 255-262. Brantley, S.L., J.T. Chesley & L.L. Stillings 1998. Isotopic ratios and release rates of strontium measured from weathering of feldspars. Geochim. Cosmochim. Actu 62: 1493-1500. Bullen, T.D. & C. Kendall 1998. Tracing of weathering reactions and water flowpaths: a multi-tracer approach. In C. Kendall and J.J. McDonnell (eds.) Isotope Tracers in Catchment Hydrology: 5 19-576. Elsevier: Amsterdam. Bullen, T.D. & P.B. McMahon 1998. Using stable Fe isotopes to assess microbially-mediated Fe3+ reduction in a jet-fuel contaminated aquifer. Mineral Mag. 62A: 255-256. Bullen, T.D., A.F. White, D.V. Vivit & M.S. Schulz 1998. Granitoid weathering in the laboratory: chemical and Sr isotope perspectives on mineral dissolution rates. Proceedings of the PIh Intl. Symp. Water Rock Interaction: 383-386. Rotterdam: Balkema. Bullen, T.D., P.B. McMahon, K.W. Mandernack, D.A. Bazylinski, C.W. Childs and A.F. White 27

1999. Using Fe isotopes in biogeochemical studies: proceed, with caution! EOS 80,46: F479. Bullen, T.D., A.F. White, C.W. Childs, D.V. Vivit & M.S. Schulz. A demonstration of significant abiotic iron isotope fractionation in nature. Submitted to Geology. Childs, C.W., C.J. Downes & N. Wells 1982. Hydrous iron oxide minerals with short range order deposited in a spring/stream system, Tongariro National Park, New Zealand. Aust. J. Soil Res. 20: 119-129. Childs, C.W., N. Wells & C.J. Downes 1986. Kokowai Springs, Mount Egmont, New Zealand: chemistry and mineralogy of the ochre (ferrihydrite) deposit and analysis of waters. J. Royal Soc. New Zealand 16/1: 85-99. Clow, D.W., M.A. Mast, T.D. Bullen & J.T. Turk 1997. Strontium 87/strontium 86 as a tracer of mineral weathering reactions and calcium sources in an alpine/subalpine watershed, Loch Vale, Colorado. Water Resources Res. 33 : 1335- 135 1. Drever, J.I. & D.R. Hurcomb 1986. Neutralization of atmospheric acidity by chemical weathering in an alpine drainage basin in the North Cascade Mountains. Geology 114: 221-224. Edmond, J.M. 1992. Himalayan tectonics, weathering processes, and the strontium isotope record in marine limestones. Science 258: 15941597. Faure, G. 1986. Principles of Isotope Geology, Second Ed. 589 p. John Wiley & Sons: New York Garrels, R.M. & F.T. Mackenzie 1967. Origin of the chemical composition of some springs and lakes. In W. Stumm (ed.), Equilibrium Concepts in Natural Water Systems, Advanced Chemical Series 67: 222-242. American Chemical Society. Hoefs, J. 1987. Stable Isotope Geochemistry, Third Ed. 241 p. Springer Verlag: Berlin. Horita, J. & M.E. Berndt 1999. Abiogenic methane formation and isotopic fractionation under hydrothermal conditions. Science 285: 10551057. Janecky, D.R. & W.E. Seyfried, Jr. 1986. Hydrothermal serpentinization of peridotite within the oceanic crust: experimental investigations of mineralogy and major element chemistry. Geochim. Cosmochim. Acta 50: 13571378. Jenden, P.D., D.R. Hilton, I.R. Kaplan, I. Isaac & H.Craig 1993. Abiogenic hydrocarbons and mantle helium in oil and gas fields. In D. Howell (ed.) The Future of Energy Gases, U.S.G.S. Prof. Pap. 1570: 3 1-56. Johnson, T.M., M.J. Herbel, T.D. Bullen & P.T. Zawislanski 1999. Selenium isotope ratios as indicators of selenium sources and oxyanion reduction. Geochim. Cosmochim. Acta 63/18: 2775-2783.

Kelly, D.S. 1996.Methane-rich fluids in the oceanic crust. J Geophys. Res. 101:2943-2962. Kendall, C. 1998. Tracing nitrogen sources and cycling in catchments. In C. Kendall and J.J. McDonnell (eds.) Isotope Tracers in Catchment Hydrology: 5 19-576.Elsevier: Amsterdam. Krabbeiihoft, D.P., C.J. Bowser, C. Kendall & J.R. Gat 1994. Use of oxygen-18 and deuterium to assess the hydrology of ground-waterllake systems. In L.A. Baker (ed.) Environmental Chemistry of Lakes and Reservoirs: 67-90. American Chemical Society: Washington, D.C. Lovley, D.R., E.J.P. Phillips, & D.J. Lonergan 1991. Enzymatic versus nonenzymatic mechanisms for Fe(II1) reduction in aquatic sediments. Environ. Sci. Technol. 25: 1062-1067. Mandernack, K.W., D.A. Bazylinski, W.C. Shanks & T.DBullen 1999. Oxygen and iron isotope studies of magnetite produced by magnetotactic bacteria. Science 285: 1892-1896. Mandernack, K.W., T.D. Bullen, K. Taga & W.C. Shanks 2000.Biogeochemical influences on iron oxidation in a stream impacted by acid mine drainage as inferred from Fe and 0 isotope systematics. EOS 81,48:F178. Millero, F.J. 1985. The effect of ionic interactions on the oxidation of metals in natural waters. Geochim. Cosmochim. Acta 49:547-553. Polyakov, V.B., 1997.Equilibrium fractionation of the iron isotopes: estimation from Mossbauer spectroscopy data. Geochim. Cosmochim. Acta

61:4213-4217. Rona, P.A., H. Bougault, J.L. Charlou, P. Appriou, T.A. Nelsen, J.H. Trefry, G.L. Eberhart, A. Barone & H.D. Needham 1992. HydrothermaI circulation, serpentinization, and degassing at a rift valley-fracture zone intersection: MidAtlantic Ridge near 15%, 45'W. Geology 20:

783. Taylor, A.S., J.D. Blum, A.C. Lasaga & I.N. MacInnes 2000. Kinetics of dissolution and Sr release during biotite and phlogopite weathering. Geochim. Cosmochim. Acta 6417:1 191 - 1208. Urey, H.C. 1947. The thermodynamic properties of isotopic substances. J: Chem. Soc. (London): 562-

581. White, A.F., A.E. Blum, T.D. Bullen, D.V. Vivit, M.S. Schulz & J.F. Fitzpatrick 1999a.The effect of temperature on experimental and natural chemical weathering rates of granitoid rocks. Geochim. Cosmochim. Acta 63, 19/20: 3277-

3291. W t e , A.F., T.D. Bullen, D.V. Vivit, M.S. Schulz & D.W. Clow 1999b. The role of disseminated calcite in the chemical weathering of granitoid rocks. Geochim. Cosmochim. Acta 63, 13114:

1939-1953.

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Whiticar, M.J., E. Faber & M. Schoell 1986. Biogenic methane formation in marine and freshwater environments: CO2 reduction vs. acetate fermentation-isotopic evidence. Geochim. Cosmochim. Acta 50:693-709. Wiederhold, J.G. 2000.Iron isotope fi-actionation in a seasonally reduced soil. Project Report for M.S. program in Environmental Soil Science, Oregon State University.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Significance of geochemical signatures in sedimentary basin aquifer systems W.M.Edmunds British Geological Survey, Crowmarsh Giford, Wallingford, Oxon. OXI 0 8BB, UK

ABSTRACT: Large sedimentary basins contain geochemical and isotopic information relating to their history of emplacement and water-rock interaction during their evolution along flow pathways. Intercomparison of data from several different non-carbonate aquifers is made to infer palaeoclimatic history, salinity sources and to distinguish between mixing and reaction. Once understood, this information may then be used to further evaluate the flow processes and quality changes in large basins and impacts on water resources development.

1 INTRODUCTION

through the natural layering and groundwater abstraction leads to mixtures in age and quality. Deterioration in water quality occurs as water is drawn either from lower transmissivity strata, or is drawn down from the near-surface where, in semi-arid areas saline waters are commonplace. The aim of this paper is to review the scope of the diagnostic geochemical and isotopic information that exists in groundwaters in large basins. The purpose is to ensure that during evaluation of such systems, maximum benefit is obtained fiom the quality data (chemical and isotopic as well as dissolved gases). Several studies of large basins already exist where integrated scientific approaches have been adopted using a wide range of tools. These include the Milk River aquifer, Canada (see Froehlich et al. 1991 and associated papers), the Great Artesian Basin, Australia, Herczeg et al. (1991); (see also Radke et al. 2000 and papers refered to therein), East Midlands aquifer UK (see Edmunds & Smedley 2000 and papers referred to therein). Many other basins are well characterised but there is often a tendency to use either isotopic evaluation or chemical evaluation without full use of the other. Once an aquifer is fully understood using the multi-tracer approach then it may be possible to use quite simple tools such as C1 to monitor quality changes involving mixing and contaminant migration into the pristine systems. A summary is given in Table 1 of the tools that may be used in the geochemical evaluation of basin systems and the specific application of each. A distinction is drawn between inert and reactive tracers, both isotopic and chemical. Investigations of large sedimentary basins benefit more than other applications from the use of both inert and reactive tracers. In this paper the information that may be obtained is

Groundwater in large sedimentary basins forms the primary resource for water supply especially in arid and semi-arid regions. Most of these resources are of high quality and accumulating geochemical and isotopic evidence indicates that their replenishment occurred during the Late Pleistocene when climates were much wetter than today. In modern times recharge in semi-arid areas is small or negligible (reference) and groundwater has to be regarded as a nonrenewable resource. In humid regions groundwater in large basins is also under threat since large scale abstraction is rapidly depleting the reserves of pristine palaeowater; where modern water is being induced by abstraction into the aquifer this invariably is less pure, containing evidence of contamination by man. Although palaeowaters found in large sedimentary basins generally constitute a valuable resource of high quality water, long residence times in certain lithologies may produce harmful accumulations of some elements and may also give rise to salinity problems. The exploitation of groundwater from some large basins has only taken place over the past few decades. This has raised expectations in a generation or so of the availability of plentiful groundwater, yet in practice falling water levels testify to overdevelopment andor inadequate scientific understanding of the resource and its origins. In addition to quantity-related issues, quality issues exacerbate the situation. Thus the natural groundwater regime established over long time scales has usually developed chemical (and age) stratification in response to recharge over a range of climatic regimes and different geological controls. Borehole drilling cuts 29

Table I . Geochemical tools used for groundwaters in sedimentary basins: 1) Absolute age measurement; 2) Recharge estimation; 3) Relative age indicators; 4) Salinity diagnostics; 5 ) Facies changes; 6) Palaeoenvironmental indicators; 7) Redox indicators; 8) Pollution indicators; 9) Geothermometry; 10) Potability and health indicators.

1

2

3

4

5

6

7

8

9

10

GEOCHEMICAL/ ISOTOPIC TOOLS Iizert Tracers

. .

c1

Br (Br/Cl) ~1 37ci/35ci 'H 6I8O, tj2H Noble gas ratios Noble gas isotopes

.

0

.

0

.

.

0

e

. . .

Major ions/ratios Si Trace alkali metals Nutrients Metals (Mn, Fe, Cr, As, U.. .. s7~r/86~r

*

. m

*

0

0

0

.

.

.

.

.

.

0

*

0

0

I4c,6I3C Organics

.

*

Reactive Tracers

6I'B

.

.

*

. .. e

0

considered under the themes 1) palaeoclimate and environment, 2) natural geochemical processes, 3 ) physical phenomena, 4) human impacts.

evidence related to landscape elements at the time of recharge. 2.1 Radiocarbon

2 PALAEOCLIMATE AND ENVIRONMENT

Central to the use of groundwaters for reconstruction, as with all archive materials is the need for a reliable chronology. Over the timescales of interest (103->106 yr) the options for absolute dating are shown in Table 1. Of these, radiocarbon still offers the most useful option for reconstructing events spanning the PleistoceneEIolocene transition. However, many caveats apply to its use due to difficulty of knowing input conditions and the water-rock interaction involved, especially where carbonates are present along flow paths (Clark & Fritz 1997). In the predominantly non-carbonate aquifers of many large basins however age correction may be applied with caution to provide calibration of the flow sequence. Relative differences between radiocarbon activities along flow lines (expressed as pmc) may allow relative timescales to be established. Due to mixing and'or reaction, however, it may not be possible to resolve ages within a few thousand years.

Palaeoclimatic reconstruction of the late Pleistocene and especially the Pleistocene/Holocene transition is now being achieved using a number of high or moderate resolution archives, especially the wealth of data from polar and mid-latitude ice cores. Some of these data e.g. pollen and lake sediment geochemistry may be used to reconstruct the past hydrology (Gasse 2000). Groundwater presents another archive which makes a contribution to the multiproxy approach to palaeoclimatic investigation although the signals retained are generally only of low resolution, typically at the scale of &I1000 years at best (Stute & Schlosser 1999). Nevertheless the presence of water of a given age is direct evidence of major wet periods which cannot be obtained by other means. Evidence within these water bodies can then, under favourable conditions, be used to decipher palaeotemperatures, the contributing air masses and

30

which marked the LGM in north Africa. There is a wide range in the 6l8O values of the palaeowaters from Libya, which probably relate to oscillations in air mass sources, between the Mediterranean westerly flow and the African monsoon extending northwards. In Libya as in most parts of northern Africa clear evidence is found of superimposed groundwater derived from the major wet phases of the early and mid-Holocene (Edmunds & Wright 1979). The data from the Chad Basin also show a considerable range in isotopic composition of palaeowaters over a narrow time span (24 to 18 ka BP), although there is a near-absence in the confined aquifer of groundwaters of Holocene age. This is explained by the absence of Lake Chad and the drainage of the aquifer towards the depressions to the NE (via the Bahr el Ghazal channel) in the Late Pleistocene. The absence of Holocene recharge is then due to the creation of the lake (Megachad) over a much greater extent than today, which covered the discharge zone during the Holocene wet period (Edmunds et al. 1997). In the Aveiro aquifer Portugal, in contrast, there is uniform isotopic composition even a slight decrease in 6l80 across the Pleistocene transition and there is no recharge gap (Carreira et al. 1996). This indicates three points: i) that recharge has been continuous in the coastal basin; ii) the enrichment in isotopic values are related to the enrichment in isotopic content of the world's oceans during glacial times, and iii) that the air mass circulation at that latitude has remained constant.

Figure 1 . Plot of 6 ' * 0 vs age for groundwaters from Europe (UK and Portugal) and Africa (Libya and Nigeria).

2.2 Stable isotopes The oxygen-18 and deuterium changes along flow lines give insight into the past temperatures or to changes in air mass movements. In Figure 1 the 6l80 data are compared for a series of dated groundwaters from sedimentary basins in Europe and Africa. Age correction and interpretation is made possible in aquifers which are essentially sandstones with, at most, very minor carbonate and where water-carbonate interaction is minimal as indicated by 613C values. These four basins demonstrate a number of interesting aspects of the palaeohydrology contained in the flow sequences. In the East Midlands aquifer there is a change in 6'*0 of around -1.5%0 over the time span of the 14C dating. In addition there is an absence of ages between about 10 and 20 kyr BP, which corresponds to the ice cover or permafrost during the last glacial maximum (LGM). Despite this it is clear that the water-rock interaction in the aquifer was continuous over this time (Edmunds & Smedley 2000). Noble gas recharge temperatures also indicate a difference of around 6°C between the mean annual modern era and that of the LGM, suggesting that the l 8 0 data here record mainly a temperature effect. The same gap in recharge is seen at the time of the LGM in the data from eastern and southern Libya, where the data may be consistent with aridity,

2.3 Other inert tracers - Cl and NO3 The combined use of chloride and the stable isotopes of water (F"0, 62H), provide a powerful technique for studying past environments in groundwaters. Over continental areas groundwater solutes are dominantly of atmospheric origin and are concentrated in proportion to evaporation during recharge. The large freshwater reserves in basins of arid zones are indicative of wet conditions; with the lowered sea levels until 8-10 kyr BP, there was also the opportunity in coastal areas for fi-eshwater to displace salinity in the areas of former coastlines. The present day distribution of groundwater salinity therefore is mainly a legacy of the onset of periods of aridity since the LGM and especially during the past 4000 yr. Nitrate remains inert in the presence of dissolved oxygen (DO) and may retain the signature of the environmental conditions at the time of recharge. Dated groundwaters from Libya, Niger, Sudan, Mali and Algeria (Edmunds 1999) contain nitrate concentrations close to or above the 11.3 mg 1-' NO3-N drinking water limit (Fig. 2); in Africa these are considered to result naturally from leguminous vegetation. The frequency of occurrence of high nitrate groundwaters appears to suggest no major changes 31

Figure 2. Nitrate in groundwaters from predominantly non-carbonate aerobic aquifers in north Africa where radiocarbon analyses (as pmc) are also available. The gap in 14Cbetween 10-20 pmc is considered to correlate with drier climates at the time of the LGM.

in plant communities during the late Pleistocene and Holocene, but support evidence for a shift northwards of the Sahelian vegetation zones by some 500

ture of the sandstone, where most of the Br enrichment occurs early in the flow line. In the Great Artesian Basin the dilute waters derived from the NE of the basin and occurring along the flow path of some 1200 km all show similar slight Br enrichment and are considered to represent long term inputs of marine aerosol-derived salts (Herczeg et al. 1991); by contrast, the saline waters are Br-depleted and have an evaporite origin. In Mali the origins of two sets of groundwaters from the unconfined Azaouad Basin (Continental Intercalaire) may also be

km. 3 NATURAL GEOCHEMICAL PROCESSES 3.1 Development of salinity - rainfall or WRI? The build-up of groundwater salinity in basins occurs through the acquisition of solutes from input sources (principally rainfall), from progressive water-rock interaction or through mixing with formation waters. It is usually possible to distinguish rainfall salinity from geological sources using Br/C1 ratios; the other halogen elements may be also be used to follow the sources and evolution of salinity in age-calibrated aquifers. An example is given (Fig. 3) from the Continental Intercalaire in Algeria (Edmunds et al. inprep) along a section from the Saharan Atlas to the discharge area in the Chotts in Tunisia where the C1 increases progressively from around 200 to 800 mg I-'. The initial Br/C1 ratio is much below sea water and these values are maintained for some 400 m along the flow lines indicating that dissolution of halite predominates (any rainfall salinity signature is overprinted). Towards the discharge area of the chotts the Br/Cl ratios increase, which suggests increasing influence of marine formation waters, possibly the merging of different flow lines. The Br/Cl ratios for a number of non-carbonate basin aquifers containing mainly fresh, potable water are shown in Figure 4. At low salinity many of these (e.g. East Midlands) show slight Br/Cl enrichment relative to sea water; changes in Br/C1 are also detected between waters dated to the Holocene and late-Pleistocene which may have palaeoclimatic significance. In the Milk River aquifer (Fabryka-Martin et al. 1991), uniform Br enrichment occurs throughout the aquifer; this is related to the organic-rich na-

Figure 3. Evolution of Br/Cl (wt) ratios and other halogen elements along a line of section in the Continental Intercalaire aquifer (AlgeridTunisia). The reference values are also given for Br/Cl ratios in sea water and halite.

32

Figure 4. Br/Cl ratio for various aquifers in Europe, N America, Africa and Australia relative to sea water composition.

Figure 5. Water-rock interaction in three freshwater, noncarbonate aquifers as shown by NdCl ratios. Arrows refer to the lines of evolution along flow paths following infiltration.

separated using the Br/Cl ratios. Slightly enriched values in the Saharan (northern) set of groundwaters are attributed to rainfall origins whereas waters with greater enrichment are likely to be derived from the Niger river (containing organic material) when it extended to the north during the Holocene (Fontes et al. 1993).

teraction with the marine clays to give the high ratio which persists for most of the flow line. Addition of saline water with equimolar Na and C1 then accounts for the observed slight decrease in NdC1 along the flow path. Other ionic ratios and trace element enrichment may also act as indicators of basin groundwater evolution. In dilute groundwaters, incongruent reaction of carbonate minerals can lead to progressive enrichment in Mg/Ca (and Sr/Ca) ratios, SO4/C1 can be used to follow the reaction of gypsum, and WNa ratios to follow the reaction of K-feldspar; empirical trends can be tested by modelling. Trace elements, especially (Sr, Ba, Li, Rb, Cs) that are released during incongruent reaction (Fig. 6) can also be used i) to demonstrate that reactions are occurring and ii) to infer residence times as in the East Midlands aquifer (Edmunds & Smedley 2000).

3.2 Progressive water-rock interaction or mixing?

A major problem in hydrogeochemical interpretation is to discriminate between evolution arising from progressive downgradient water-rock interaction and mixing with already evolved water. To do this, inert and reactive tracers must be compared and the simple example of NdC1 may be used to illustrate. Data from three aquifers with defined residence times (Milk River, Great Artesian Basin and East Midlands) are compared (Fig. 5). In the East Midlands and the Great Artesian Basin an evolution from a low salinity water with NdC1 - 1 at outcrop can be followed. This represents inputs from rainfall having different amounts of evaporative concentration. In the East Midlands aquifer inputs from various contaminant C1 sources with low NdC1 from the modern era are also distinguishable. The increase in NdCl without C1 increase is diagnostic of reaction in this case a small increase of 10 mg 1-' Na over a timescale of some 40 kyr. The reaction of albite rather than ion exchange is confirmed by PHREEQC modelling; this freshwater diagenesis also involves the dissolution of K-feldspar (Van der Kemp et al. in press). In the Great Artesian Basin a similar conclusion was reached for an aquifer which has evolved over >1O6 yr. In the Milk River aquifer (Hendry et al. 1991) an increase in Na was recognised along the flow gradient attributed to the diffusion of both Na and C1 from the overlying Colorado Group shale. Here however, initial waters are distinctive by their high NdC1. This must indicate rapid water-rock in-

Figure 6 . Lithium showing progressive increase in low-Cl palaeowaters (pre-industrial era) as compared with mixing with low C1 modem (contaminant) waters.

33

3.3 Redox processes Groundwater in the phreatic sections and in parts of the confined sections of many large basin aquifers in Africa and elsewhere have evolved under anaerobic conditions. DO concentrations of several mg 1-I may persist in continental, non-carbonate aquifers for many thousands of years as reported for example from USA (Winograd & Robertson 1982; Rose & Long 1988), the Kalahari (Heaton et al. 1983) as well as in red bed aquifers from other areas (Edmunds & Smedley in review). These persistent aerobic conditions maintain very low concentrations of dissolved Fe, but may favour the mobility of those elements such as As, Se, MO, V, U, Cr which can form oxy-anion complexes such as MOO;, C1-03~and complexes with carbonate such as UOZ(CO~)~-. An example is given (Fig. 7) from the Continental Intercalaire aquifer in AlgeridTunisia (Edmunds et al. in prep) where a chronology of flow along the same line has been established (Guendouz et al. 1998). A redox boundary is inferred from the abrupt concentration changes for example in Fe and N03, neither DO nor redox potential (Eh) having been measured. The redox boundary is found some 300 km from the recharge area of the Saharan Atlas and some 200 km beyond the limit of modern or Holocene groundwaters (based on radiocarbon, stable isotopes, as well as noble gas data). The persistence

of DO over this distance demonstrates that the aquifer must be virtually free of organic material, pyrite or other electron donors. 4 PHYSICAL PHENOMENA The spatial distribution of chemical and isotopic indicators (as well as their variation with time) can also provide evidence of physical changes and can be used in the calibration of flow models. 4.1 Rates of recharge The use of C1 in unsaturated zone profiles to measure recharge is now well established (Allison et al. 1994). Records of the recharge variations with a resolution of decades to centuries may be stored as salinity for between 102-1O5 yr (Tyler et al. 1996). It follows that the salinity variations in the large basins may also reflect the recharge rates (see 3.1). Thus concentrations of C1 can be used as a qualitative guide to the minimum rates of recharge in basins. In basins such as East Midlands or the Great Artesian Basin little increase in C1 takes place over tens of km (thousands of years) and the variations in salinity (supported by Br/CI), may record a smoothed record of long term recharge rates and associated climate change. 4.2 Rates of movement and crossformational flow Flow velocities calculated using hydraulic parameters often give estimates at odds with rates of movement found using isotopic tracers. Hydraulic derived rates may exceed radiocarbon estimates by an order of magnitude as in the Milk River aquifer (Drimmie et al. 1991). Several factors may account separately or together for the discrepancy; i) parameters used may be derived from present day recharge estimates and piezometry which may have been different to those operating under natural gradients and with different climates in the past; ii) discharge may have occurred under aquifer-full conditions e.g. to rivers; iii) transmissivities may decrease with depth; iv) water is preferentially lost by leakage through aquitards.

4.3 Stratification, mixing and depth of circulation The degree of flow heterogeneity in an aquifer is usually expressed by the scatter of the geochemical data; conversely smooth trends along flow lines are likely to be an expression of homogeneous lithology and physical properties or well mixed systems. Samples of groundwater fiom large capacity boreholes in sedimentary basins are inevitably mixtures and information on stratification is needed where possible. Hydrogeophysical logging as a prelude to

Figure 7. Redox boundary expression by NO3, Fe, Cr and U in the Continental Intercalaire aquifer (AlgeridTunisia), along an 800 km flow line from the Atlas Mts to the discharge area in the chotts and oases of southern Tunisia.

34

tions prevailing, little attenuation capacity is provided and agricultural chemicals and other wastes will tend to accumulate rather than be degraded. Significant threats come however from irrigation practices where return waters will increase salinity, and from wastewater discharges, where transit times are likely to be rapid. Drawdown of saline waters from discharge areas (oases and sebkhat areas) as well as contamination from evaporites or saline formation waters in multilayered aquifers present hazards to the abstraction of fresher palaeowaters which are found in the most permeable horizons. Chemical and isotopic tracers are of particular diagnostic value for identifying salinity origins (Edmunds & Droubi 1998) and relevant techniques are indicated in Table 1 above. The long residence times of groundwaters in large sedimentary basins lead to the build up of some elements that may inhibit the groundwater use. Redox zonation may give rise to the accumulation of metals especially those forming oxy-anions in aerobic zones. In the aerobic section, of the East Midlands, for example, increases in concentrations of As, Se, Sb, Cr, MO, U towards the redox boundary are observed.

Figure S . Hydrogeochemical depth profile of groundwater from a newly drilled borehole in the East Midlands Triassic aquifer showing Holocene overlying late Pleistocene groundwaters above 200 m.

6 CONCLUSIONS AND APPLIED SIGNIFICANCE

depth sampling provides a powerful tool for resolving the three-dimensional properties of flow systems. In the East Midlands aquifer radio- and stable isotopic and chemical data demonstrate the extent of this stratification in a newly drilled borehole (Edmunds & Smedley 2000). The depth of origin of groundwaters in large basins is often not available due to loss of records, collapse of well screen for example and in this case the well head temperature can indicate the source horizon (or interval). Chemical geothermometers may also be used to verify the maximum depths of circulation. In sedimentary basins, silica geothermometry may be used although the alkali earth geothermometers are less appropriate (Truesdell 1984).

The exploitation of groundwaters from large basins is costly and it is emphasized here that a relatively small investment in geochemical ad isotopic analysis can be highly cost effective for the long term management process. At the exploration phase geochemical data can help to understand the origins of the groundwaters, their renewal rates and major facies changes in the water quality, such as redox and salinity boundaries. During evaluation, emphasis needs to be placed on maximising the information on depth stratification in age and in quality (using hydrogeophysical logging or pore waters) so that any subsequent changes in quality may be recognised. It is important that where possible large well-field schemes invest in one or more research boreholes where depth information may be obtained from core material; pore waters extracted from cores provide precise depth sampling of the water quality and the cores themselves provide reference material on the solid phase for geochemical modelling and accurate data on permeability and porosity for transport modelling. At the development phase monitoring programmes are most effective if supported by the detailed studies outlined so that the correct indicators of changes in water quality may be selected and interpreted.

5 HUMAN IMPACTS AND NATURAL HAZARDS Quality related risks associated with development of large basin groundwaters, in addition to overall palaeowater depletion, are threefold; i) anomalies in the natural baseline chemistry associated with waters of long residence time, ii) inducement of salinity changes, and iii) anthropogenic contamination. Anthropogenic contamination is a limited problem due to the size of basins and even for unconfined areas a degree of protection is afforded against diffuse or point source pollution due to low rates of recharge. However, with widespread aerobic condi35

ACKNOWLEDGMENTS

Guendouz, A., Moulla, A., Edmunds, W.M., Shand, P., Poole, J., Zouari, K. & A. Mamou 1998. Palaeoclimatic information contained in groundwaters of the Grand Erg Oriental, North Africa. In: Isotopic Techniques in the Study of Environmental Change, 555-57 1. Vienna: IAEA Heaton, T.H.E., Talma, A.S. & J.C. Vogel 1983. Origin and history of nitrate in confined groundwater in the western Kalahari. Journal of Hydrology, 62: 243-262. Hendry, J., Schwartz, F.W. & C. Robertson 1991. Hydrogeology and hydrochemistry of the Milk River aquifer system, Alberta, Canada, a review. Applied Geochemistry, 6: 369380. Herczeg, A.L., Torgersen, T., Chivas, A.R. & M.A. Habermehl 199 1. Geochemistry of ground waters from the Great Artesian Basin, Australia: Journal of Hydrology, 126: 225-245. Radke, B.M., Ferguson, J., Cresswell, R.G., Ransley, T.R. & M.A. Habennehl2000. Hydrochemistry and implied hydrodynamics of the Cadna-owie - Hooray Aquifer, Great Artesian Basin: 248pp. Kingston (Australia). Bureau of Rural Sciences. Rose, S. & A. Long 1988. Dissolved oxygen systematics in the Tucson Basin aquifer: Water Resources Research, 24: 127136. Smedley, P.L. & W.M. Edmunds in prep. Redox patterns and trace element behaviour in the east Midlands Triassic sandstone aquifer UK. Stute, M. & P. Schlosser 1999. Atmospheric noble gases. Chapter 11 (Pp. 349-377) In: P.G Cook and A.L. Herczeg (eds) Environmental Tracers in Subsurface. Hydrology: 349-377. Kluwer, Boston. Truesdell, A.H. 1984. Chemical geothermometers for geothermal exploration. Chapter 3 In: R.W. Henley, A.H.Truesdel1, Barton, P.B. (eds) Fluid-mineral equilibria in hydrothermal systems. El Paso: Society of Economic Geologists. Tyler, S.W., Chapman, J.B., Conrad, S.H., Hammermeister, D.P., Blout, D.O., Miller, J.J., Sully, M.J. & J.M. Ginnanni 1996. Soil-water flux in the southern Great basin, United States: temporal and spatial variations over the last 120,000 years. Water Resources Research, 32: 1481-1499. Van der Kemp, W.J.M., Appelo, C.A.J., Condesso de Melo, T., Gaus, I., Milne, C.J. & K. Walraevens. In press Hydrochemical modelling as a tool for understanding palaeowaters. In: W M Edmunds & C.J.Milne (eds) Palaeowaters in Coastal Europe:Evolution of Groundwater since the late Pleistocene. Special Publication Geological Society of London. Winograd, I J., & F.N. Robertson 1982. Deep oxygenated groundwater: anomaly or common occurrence? Science, 2 16:1227-1230.

This paper is published with the permission of the Executive Director, British Geological Survey, Natural Environment Research Council.

REFERENCES Allison, G.B., Gee, G.W, & S.W. Tyler 1994. Vadose-zone techniques for estimating groundwater recharge in arid and semiarid regions. Soil Science Society of America Journal, 58: 6-14. Carreira, P.M., Soares, A., Marques da Silva, M.M., Araguas, M.A. & K. Rozanski 1996. Application of environmental isotope methods in assessing groundwater dynamics of an intensively exploited coastal aquifer in Portugal. In: Isotopes in Water Resources Management, 2: 45-58, Vienna, IAEA. Clark, I. & P. Fritz 1997. Environmental isotopes in hydrogeology. Lewis, Boca Raton. Drimmie, R.J., Aravena, R., Wassenaar, L.I., Fritz, P., Hendry, M.J. & G. Hut 1991. Radiocarbon and stable isotopes in water and dissolved constituents, Milk River aquifer, Alberta, Canada. Applied Geochemistry, 6: 38 1-392. Edmunds, W. M., E. Fellman, I. Baba Goni, G. McNeill, & D. D. Harkness 1997. Groundwater, palaeoclimate and palaeorecharge in the southwest Chad basin, Borno State, Nigeria: Isotope Techniques in the Study of Past and Current Environmental Changes in the Hydrosphere and Atmosphere: 693-707. Vienna, IAEA. Edmunds, W.M. 1999. Groundwater nitrate as a palaeoenvironmental indicator. In: Geochemistry of the Earth’s Surface: 35-38. Rotterdam: Balkema. Edmunds, W.M. & P.L. Smedley 2000. Residence time indicators in groundwater: the East Midlands Triassic sandstone aquifer. Applied Geochemistry, 15: 737-752. Edmunds, W.M. & A. Droubi 1998. Groundwater salinity and environmental change. In: Isotope Techniques in the Study of Environmental Change: 503-5 18. Vienna: IAEA. Edmunds, W.M. et.al. in prep. Groundwater evolution in the Continental Intercalaire aquifer of southern Algeria and Tunisia: major, trace element and isotopic indicators. Edmunds, W. M., & E.P. Wright 1979. Groundwater recharge and palaeoclimate in the Sirte and Kufra basins, Libya. Journal of Hydrology, 40: 2 15-241. Fabryka-Martin, J., Whittemore, D.O., Davis, S.N., Kubik. P.W. & P. Sharma 1991. Geochemistry of halogens in the Milk River aquifer, Alberta, Canada. Fontes, J.Ch., Gasse, F. & J.N. Andrews 1993. Climate conditions of Holocene groundwater recharge in the Sahel zone of Africa. In: Isotope Techniques in the Study of Past and Current Environmental Changes in the Hydrosphere and the Atmosphere: 23 1-248. Vienna, IAEA. Frohlich, K., Ivanovich, M., Hendry, M.J., Andrews, J.N., Davis, S.N., Drimmie, R.J., Fabryka-Martin, J., Florkowski, T., Fritz, P., Lehmann, B., Loosli, H.H., & E. Nolte 1991. Application of isotopic methods to dating of very old groundwaters: Milk River aquifer, Alberta, Canada. Applied Geochemistry, 6: 465-472. Gasse, F. 2000. Hydrological changes in the African tropics since the Last Glacial Maximum. Quaternary Science Reviews, 19: 189-211.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Interactive processes due to subglacial volcanic activity Local phenomena with global consequences H.Kristmannsd6ttir Orkustofiun, Grensasvegur 9,108 Reykjavik

ABSTRACT: Vast areas of the inland of Iceland are covered by glaciers and they also coincide with the zones of active rifling and volcanism. Subglacial geothermal activity and volcanic eruptions initiate interactive processes between acidified meltwater and basaltic tephra on one hand and subsequent reactions with atmospheric air in the runoff rivers. The effect of interactive processes on glacial water chemistry before or at early stages of the volcanic activity can also be made use of in the design of automatic warning systems for sudden floods from the glaciers due to volcanic activity or increased geothermal activity due to emplacement of magma at shallow levels in the earth crust or prior to volcanic eruptions. Even though such phenomena are very localized and almost confined to Iceland the consequences may be noticeable on a global scale as influx of such floodwater into the oceans may result in CO;! removal from the atmosphere. Further more the interactive processes are similar to processes observed and expected to take place in groundwater and surface waters around active volcanoes.

1 INTRODUCTION

subglacial volcanic activity are somewhat similar to the effects of subaerial volcanic eruptions on local groundwater and surface water chemistry. That volcanic eruptions affect local surface waters and local springs is well known (Kjartansson 1957, Klein 1981, Tilling & Jones 1996). Such effects are especially prominent when chemicals are accumulated in snow (Gislason et al. 1992) and then suddenly released in the first thaw. One of the differences between the effects of a subglacial and a subaerial volcanic eruption is that great amounts of volatiles are released to the atmosphere in a subaerial eruption whereas in a subglacial eruption most of the volatiles are confined to the meltwater formed during the eruption. In the case of intrusive activity the composition of gas emitted from a magma will be considerably modified if it reacts within a water dominated geothermal system prior to emission. The conditions in Iceland are unique with big glaciers covering active volcanoes. Nowhere else in the world similar environments are known even though small glaciers and ice sheets are connected to volcanoes elsewhere like in Alaska (Trabant et al. 1994) and water volcanic interaction occurs in crater lakes in New Zealand (Hurst & Dibble 1981, Hurst et al. 1991) and a few other places (Casadewall et al. 1984, Kusakabe et al. 1989). The interactive processes observed between the volcanic products, glaciers and atmosphere in this unique environment, are thus quite local phenomena.

Iceland is very active volcanically due to its location on the Mid Atlantic Ridge. Spreading takes place in the direction away from the central zone of active rifting (Fig. 1) where the main volcanic centers are located as well. Several fairly large glaciers still cover the highest mountains in Iceland and many of the volcanic centers are located below them as well as many of the high-temperature geothermal areas in the country (Szmundsson 1979) with subsurface temperatures exceeding 200 "C at 1 km depth. The high-temperature geothermal fields are believed to draw their heat from cooling magma bodies so the effects of a volcanic event and of an active geothermal area on glacier river water may be somewhat similar (Bjornsson et al. 1982). Due to volcanic/geothermal and glacier interaction the danger of major floods is an impending reality in Iceland. Prior to a major flood due to volcanic activity or increased geothermal activity changes may be detected in the chemistry of the glacial river water (Kristmannsdottir & Bjornsson 1984, Gislason et. al. 1998, Kristmannsdottir et al. 1999, 2000a, b). The reason may be that during a volcanic event shallow sills or magma chambers may be intruded near to the surface causing increased geothermal activity and subsequent melting of glacier ice. It is known that prior to a volcanic event such intrusive activity may affect groundwater chemistry (Sigvaldason 1964, Olafsson & Kristmannsdottir 1989). The effects of 37

Figure 1. A map of Iceland showing the location of the biggest glaciers, the active zones of rifiing and volcanism and the main volcanic centers in the country. Based on Saemundsson (1 979).

In spite of that they are of great general importance both locally and probably also globally. Locally they may greatly affect the physical and biological environment at least temporarily. Globally they may affect the growth of biota in the oceans and induce a short term decrease in atmospheric CO:! (Gislason et al. 1998). Changes in glacial river water due to volcanic activity andor increased geothermal activity (Kristmannsdottir et al. 2000a) are well documented, but may still be difficult to detect at an early stage, and to differentiate between the effects from a volcanic event and a drainage from a geothermal field or subglacial lake. The main reason is that normally the seasonal changes in chemistry of the river are not well documented, primarily due to lack of funds for research. After a subglacial volcanic eruption in Vatnajokull in 1996 (Gudmundsson et al. 1997) and a subsequent catastrophic flood from the Grimsvotn subglacial lake a research project comprising additional studies of changes in chemistry of water in glacial rivers draining the Vatnajokull and Mfrdalsjokull glaciers was initiated. The main aim of the project was to develop an automatic warning system for volcanic eruptions and extreme floods (Kristmannsdottir et al. 2000a, b, Elefsen et al. 2000). As a consequence much more comprehensive data has now been acquired about the water chemistry of glacial rivers in Iceland, even though more detailed studies are needed.

2 BACKGROUND The chemical weathering rate in Iceland is very high (Gislason et al. 1996), mainly due rapid mechanical weathering and the common occurrence of very soluble fresh and glassy basalts. The origin of the chemical constituents in glacial river water derive from at least three different sources.

Figure 2. Conductivity against flow rate measured monthly over one year for the river Skeidara, on Skeidarkrsandur south of the Vatnajokull glacier. Its water is mostly glacial meltwater and the flowpath from the mouth of the glacier to the sea is less than 10 km. Occasionally, in jokulhlaups, there is geothermal/volcanic input into the massflow of the river (Kristmannsdottir et al. 2000a, b).

38

Figure 3. Total dissolved solids (TDS) in water samples from the river Skeidara draining Vatnajokull glacier to the south as well as the Grimsvotn subglacial lake, during the years 1962-1997 arranged by months (Palsson & Vigfiisson 1996, and Orkustohun data).

Firstly there is meltwater which has been drained far away and flowed a long way below the glacier and has leached chemical substances from the ice and reacted with the bedrock and sediments for a considerable period of time and therefore contains relatively high concentrations of dissolved solids. Secondly there is meltwater from the spring and summer thaws, which has not had much time for leaching and reaction with rocks and ice and contains considerably less dissolved solids. Thirdly there is meltwater and run-off from subglacial geothermal areas which has a much higher concentration of dissolved solids and quite different chemical characteristics from the other type of glacial meltwaters. Mixing of the first two components will often give quite a regular relation between the concentration of a major chemical component and discharge. Influx of the geothermal component on the other hand often occurs in irregular events interrupting a regular relation between flow and concentration. In Figure 2 the relation between conductivity as a measure of total concentration of dissolved solids and flow rate for the River Skeidara, draining Vatnajokull and where jokulhlaups from the subglacial lake and volcanic centerhigh-temperature geothermal area at Grimsvotn occur is shown. The river Skeidara is located in Skeidarhsandur to the south of the Vatnajokull glacier. Its massflow is mostly composed of glacial meltwater and its flowpath from the mouth of the glacier to the sea is less than 10 km. Occasionally, in jokulhlaups from Grimsvotn, there is a high geothermal/volcanic input into the mass of the river (Kristmannsdottir et al. 2000a, b) and there

may also at times be some leakage from the lake especially during volcanic activity (Bjornsson & Kristmannsdottir 1984, Gislason et al. 1997). The Grimsvotn subglacial lake is thus connected to one of the most active central volcanoes in Iceland and the biggest high-temperature geothermal field in the country. There is continuous melting of ice and release of geothermal fluids, mostly steam and gases into the system and jokulhlaups occur at relatively regular intervals when the water level has reached a certain limiting height (Bjornsson 1974, 1992, Bjornsson & Gudmundsson 1993, Gudmundsson et al. 1995). Occasionally a jokullhlaup is either triggered by or follows a volcanic eruption in either the Grimsvotn volcanic system or one of the other volcanoes below the Vatnajokull icecap (Gudmundsson & Bjornsson 1991, Gudmundsson et al. 1997). The volatile release of the Grimsvotn volcano has been estimated to be of the same order of magnit,ude as many major active ,volcanoes worldwide (Agustsd6ttir et al. 1992, Adstsdottir & Brantley 1994). The jokulhlaups from Grimsvotn have occurred at irregular intervals, 1-12 years during the last century. Most of the points in Figure 2 for conductivity against flow rate over one year fall on a rather regular curve, whereas a few samples from floods fall well outside and show high conductivity for low flow rates. The figure shows the conditions at fairly normal times with no major jokulhlaups, but apparently a small "leakage" from a geothermally influenced water component. Even though geochemical monitoring has neither been regular nor continuous in Skeidara the total concentration of dissolved sol-

39

charge (Fig. 5 ) is almost a straight line with conductivity steadily decreasing with increased flow rate.

ids (TDS) has been measured intermittently for a long time as well as discharge and sediment load. In Figure 3 TDS (arranged by the months) from Skeidara during the years 1962 to 1997 is shown, reflecting quite clearly both great increases of TDS during jokulhlaups, some seasonal changes in TDS and periods of "leakage" of geothermal water fiom Grimsvotn. In Figure 4 is shown the relation between conductivity and flow rate in the river Mulakvisl draining the glacier Myrdalsjokull (Fig. l), also an area very active volcanically. Below that glacier one of the largest volcanoes in Iceland is located (Thorarinsson 1959, 1967) from which giant floods have occurred approximately twice a century. The figure shows on one hand a rather similar curve as the previous one for the Skeidara river and on the other hand a more or less constant conductivity. In this river there is a high proportion of geothermal influx and floods may occur. The area has been fairly active seismically during the last few years and a jokulhlaup fiom a supposed volcanic or intrusive event did occur in 1999, mostly into another river draining the glacier Myrdalsjokull.

0

h

p

5 "I

-

0

0

0 0

0

0

0

20

O 0

40

0

60

80

80

100

120

The changes observed in the glacial river water will be somewhat different if there is a volcanic eruption influencing the water chemistry than by geothermal activity, even though some changes are rather similar. Many of the geothermal fields below the Icelandic glaciers are closely connected to volcanic centers and the difference between volcanic and geothermal effects is not always very distinct. Two rather small cauldrons in Vatnajokull, not far from the Grimsvotn subglacial lake, give rise to frequent jokulhlaups in the river Skafta draining to the southwest of Vatnajokull, at intervals of one to two years (Zophoniasson & Palsson 1996). Einarsson & Brandsdbttir (1997) have suggested that small volcanic eruptions quite fiequently occur at the cauldrons and may follow many of the jokulhlaups.

200

PO

60

Figure 5. Conductivity against flow rate rate measured monthly over one year for one of the rivers north of the Vatnajokull glacier, Skjalfandafljot. Its massflow is mostly spring water in origin with a minor input of glacial meltwater. No constant or seasonal geothermal component is part of the discharge.

3w -0

100

40

Flow rate (m'ls)

Mulakvisl

3 --

20

I00

Flow rate (m'/s)

Figure 4.Conductivity against flow rate measured monthly over one year for the river Mulakvisl, draining the M9dalsjokull glacier (Fig. 1) and there is a considerable geothermal influx (Kristmannsdottir et al. 2000a, b) into the river.

3 EFFECTS OF INCREASED GEOTHERMAL ACTIVITY Besides being much more highly mineralized (Arnorsson et al. 1982) geothermal fluids (Kristmannsdottir et al. 1999) carry some substances not normally found in glacial river water. Other components will be found in greatly different relative concentrations from those of the normal glacial river water. The subglacial high-temperature geothermal fields will emit steam melting the ice, and some of its components are subsequently dissolved into the meltwater. Some deep water component may also enter the system (Bjornsson & Kristmannsd6ttir 1984). The gases accompanying geothermal steam (Armannsson et al. 1982, Arnorsson 1990, Giggenbach 1992, Kristmannsdottir et al. 1999) will be mainly C02, HzS, CH4 and Hl. The relative concentrations vary, but are most commonly, like those

Figure 5 shows the relation between conductivity and flow rate for a river falling towards the north from the Vatnajokull glacier, Skjalfandafljot. The massflow of the river is mostly spring water originating from the vicinity of the roughly 200 km long flowpath towards the sea shore with a small input of glacial melt water. No constant or seasonal geothermal component is part of the discharge and no major floods from volcanic or geothermal events have been recorded in the river since early last century (Thorarinsson 1950). Due to the location of the river floods may be expected in the event of subglacial eruptions in the northern part of the glacier Vatnajokull (Fig. 1). The relation of conductivity to dis-

40

the atmosphere and organic soil dilutes the 14C concentration of the water, and thus the apparent 14Cage will probably be relatively high, and also yield a relatively high 613C value (Sveinbjornssdottir et al. 1998).

shown in Table 1. The CO2 concentration will be far the highest and often there is a considerable H2S concentration. Mercury concentration may be relatively high in the gas, but it will be quickly oxidized and vented out of the water (Edner et al. 1991, Varekamp & Buseck 1983). Arsenic may also be in rather high concentration in steam (Armannsson & Kristmannsdottir 1992) as well as ammonium and boron. Subsequently to the dissolution of the acid gases the meltwater will become quite acid enhancing reaction with the sediments within the subglacial lakes and those carried by the subglacial rivers. The sediments in question will mostly consist of fresh basaltic hyaloclastites (Palsson & Vigfusson 1996) which are easily leached by the acidified waters. As studies of palagonitization have shown (Jakobsson 1972) the alkali metals will be leached out first, followed by calcium, aluminium, silica and magnesium. Where the water accumulates below the glacier for considerable periods as in the Grimsvotn subglacial lake (Bjornsson & Kristmannsdottir 1984, Kristmannsdottir et al. 1999, Gislason et al. 1997) extensive reactions will be expected to take place. Metals will be leached out of the basalt tephra in contact with the acidified meltwater. The water will also react with atmospheric air enclosed in the glacial ice. Part of the reactions will happen at somewhat elevated temperatures and there is expected to be some convection within the system, but probably some parts of the lake will become stratified and stagnant. The result of these reactions will depend on the input of geothermal fluid, the delay time in the subglacial lake for reactions and evolution of the system, the temperature and convection in the lake. The net outcome would be expected to be some kind of bicarbonate-sulfate waters. As long as the water is enclosed in a subglacial lake and also when it flows below the glacier or within subglacial channels the accessibility to atmospheric air is limited and it is expected to be undersaturated with respect to oxygen when it emerges at the outlets of the glacier. From then on it can react freely with atmospheric air and the more the longer the subaerial flowpath of the river. The water will however be expected to be supersaturated with other volatile substances like CO2 and H2S, which will then be vented out along the subaerial flowpath. Mercury will also be quickly vented out and oxidized. Changes in the stable isotope ratios 6D and 6'*0 will also be observed as the different water components have different origins. The carbon isotope ratio of 13Cand the activity of I4C will also show changes. The impact of geothermal fluids on the stable 6D and 6I8O in the river water is ex ected to be rather insignificant and the 6D and 6l 0 ratios probably depend largely on the altitude where the water has fallen as rain and may also depend on the age of the water. Addition of carbon from sources other than

4 EFFECTS OF VOLCANIC ACTIVITY In a subglacial volcanic eruption the meltwater will come into direct contact with the magma on the glacier floor and therefore the chemical changes will be more significant and somewhat different from those resulting from the input of purely geothermal steam and water. Among the volcanic gases (Bames 1984, Muffler et al. 1992, Gerlach & Casadewall 1986, White & Waring 1963) there will be CO, Sol, COS, S2, HCl and HF besides the ones (C02, H2S, CH4 and H2 etc.) present in geothermal steam. Higher concentrations of mercury will be expected to accompany a direct influx of magma (Cox 1983) into meltwater than high-temperature geothermal activity. The effects on 6D and 6"O ratios in the water will probably not be significant as the mass fraction of magmatic H20 will be very low in relation to that derived from glacier ice. The isotopic ratios will therefore mostly depend on the location of the eruption and the age of the melted ice. The CO2 gas from a magmatic source will contain no I4C and as the main source of carbon will be from the magma the apparent 14C age of the water will be high and a relatively high 613C value will be observed. The acidic gases HCl, HF and SO2 will react instantly with water and the resulting meltwater will become extremely acid and have a much lower pH than meltwaters originating from increased geothermal activity. The concentrations of sulfur species, sulfate (SO4), thiosulfate (S2O3) and possibly hydrogen sulfide (H2S) (Casadewall et al. 1984) will be much higher as well as concentrations of chloride, fluoride and ammonia. Heavy metal concentration may increase considerably due to a sharp drop in pH and subsequent reaction with tephra. 5 OBSERVATIONS AND ANALYTICAL DATA The results of reactions between meltwater, geothermal fluids and volcanic emanations have been studied in water samples collected during jokulhlaups from the subglacial lake Grimsvotn (Fig. 1) in Vatnajokull (Rist 1955, Sigvaldason 1965, Steinthorsson & Oskarsson 1983, Kristmannsdottir & Bjornsson 1984, Bjornsson & Kristmannsdottir 1984, Kristinsson et al. 1986) and from the nearby Skafik cauldrons. Samples from rivers draining the glacier Mfidalsjokull showing more or less constant influence of geothermal activity have also been

r

41

Table 1. Chemical composition of: (1) A typical high-temperature geothermal water from Krafla geothermal field (Kristmannsdottir et al. 1999), (2) the range of composition of water during jokulhlaups in Skeidara 1954-1999 (Kristmannsdottir et al. 2000a, Palsson et al. 1999), (3) typical composition of water in Skeidara at the peak of summer (Orkustofnun unpublished data), (4) Sample CO]lected at the eruption site in Vatnajokull 1996 two months after the end of eruption (Gislason et al. 1997) and (5) chemical composition of a spring in Grimsvotn in 1984 a year after an eruption in 1983 (Orkustofnun unpublished data). Concentrations in mgll.

Location

High-temp. geothermal field (1)

Skeidara jokulhlaup (2)

Skeidara summer (3)

Vatnajokull eruption site (4)

Grimsvotn spring 1984 (5)

Temperature "C

300

0

5

2

80

pH I "C

6.6120

6.0-7.5120

7.6123

7.39119

7.0122

Hydr sulf. (H2S) Tot.cXb.(C02)

478

0-0.3

0

48600

340-680

26

330-460

30

TDS Silica (SiOz)

313

44-67

Sodium (Na)

83

10-90

Potassium (K)

14.2

Calcium (Ca)

74

558 950

76

325

4.8

56.2

145

0.6-19

0.25

3.9

30

1.4

28-77

6.6

33.3

47

Magnesium (Mg)

0.05

8-16

1.4

1.9

20

Sulfate (SO,)

55.5

13-1 10

2.6

148

38.5

Chloride (Cl)

38

4-43

1.5

5

18.5

Fluoride (F)

0.3

0.16-0.66

0.06

1.5

1.5

Aluminium (Al)

1.o

1 the hypothesis is supported by the datum. For 1 > B > 101I2 there is some evidence against the hypothesis (but no more than a bare mention) while for 10-”2> B > 10-’ the evidence is substantial. Being the Bayes factor of sample BIS 13 higher than 1, the hypothesis that BIS13 belong to the same population of the theoretical trend is supported. For the other two mature N a - H C 0 3 waters however (samples SC 102 and SC 103; actually pertaining to a neighboring, but geologically interconnected, catchments) the computed Bayes factors are respectively 0.198 and 0.762 (Table 3). We will nevertheless see later on that a forward model computation extended beyond the upper limit of the inverse model accounts for the chemistry of these samples rising substantially their Bayes factors (italics in Table 3). Finally we may note that, according to Bayes factor analysis, the evidence that the immature waters BIS72, BIS75 do not belong to the same population is decisive (B < 1o-2).

(39) where u = ( x o - x t ) / o x ; v = ( y o - y r ) / o y In Figure 8 we see the elliptical contours of equal probability density calculated for the mature NaHC03 water BIS 13 (mean of duplicated analyses).

5.2 Step 2: local resolution of reaction surfaces With the continuum limit satisfacted, the vector of solute concentrations was readily obtained by Marini et al. (2000) by the partial derivatives in dc of the

Figure 8. Contours of equal probability density for sample BIS13. The point of tangency with the theoretical trend defines the Bayes factor of the datum.

56

equations given in Table 1 (see also Sciuto & Ottonello 1995b):

The compositional path depicted by the continuous functions in 6 was ascribed to 4 main (kinetically controlled) reactions (4 1-1,2,3,4) involving feldspars and carbonates and assuming monomeric aluminum and dissolved silica content as buffered by two instantaneous equilibria (4 1-5,6).

(length Nr; comprising both equilibrium and kinetically-controlled reactions). Since experimental reaction rates are usually expressed in a time scale, conversion of experimental data from time-mode to progress mode is required. However, being expressed in fractional form, this conversion implies the simple knowledge of final and initial "ages" of the reacting waters. The following t -+ I; conversion (s time scale), based on the measured 3H activities, was adopted:

NaAISi,O, + 4 H + a N a + + A 1 3 + + 3 S i 0 2 ( , ,+2H,O , (41-1)

t(s) = 3.1536 10+7exp (-1.6282 + 8.5693 6). (46)

dn --

d4-

= ff(6) x 10~") x In 10

KAlSi,O,

(40)

*

+ 4H+ (rs K' + A13++ 3Si02,,,, + 2H,O (41-2)

Ca2++ CO:c a ~ gCO,), ( tjCa2++ ~ g , + +2 ~ 0 : -

CaCO,




+ 6H'

(rs

2A13++ 2SiO,,,,,

+ 5H,O (41-6)

Solving at constant P,T conditions also AI3+ and Si02,,,, are implicitly constant and the system of equations (41) was thus rewritten, in terms of ordinary differential equations (ODE) as:

through a Monte Carlo exploration of the model space followed to Simplex and Steepest Descent approaches. Although the chi-squares method is common practice in geochemical exploration it is worth to stress the differences between this method and the least squares procedure in terms of informational density content. To this purpose I will denote the model parameters obtained from the first inverse modeling procedure as dohs of the new procedure, with Gaussian uncertainties with respect to the true value described by the covariance operator Cd. As done in the first step, I will assume that I do not have any a priori information on the new model parameters. Finally, I will assume that the theoretical relationship d = g(m) holds only approximately, with Gaussian uncertainties described by the covariance operator CT. The probability density representing the a posteriori information in the new model space is (Tarantola 1987):

In matrix notation, the problem may be expressed as follows: 1 0 0 0 0 0

0 1 0 0 0 0 0 0 1 0 0 0

dNa'/d< dK+/d< dMg2'/d
1 leads to the activity paths superimposed in Figures 3 to 6 ( [XI symbols). We may appreciate thus that the model behaves correctly (from a thermodynamic point of view) attaining (or at least approaching) the limits of complete equilibrium (invariant andor univariant points). Concerning more properly the analysis of the informational content of the log [Ca2']/[H'I2 versus log [Mg2']/[H'I2 activity plot, although the slope and intercept of the linear relation are quite similar to those of the first inverse modeling step for 6 11 (Fig. 3), the abrupt changes in the activityactivity paths observed for >1 are sufficient to rise substantially the Bayes factors of the mature waters (Table 3) confirming in this way the complete geognostic significance o f the model. Being the extended forward model parameters non linearly correlated, the computation of the a posteriori probability density cannot be performed.


< ) tope ratio data that are impossible or extremely costly by thermal ionization methods. If the future is as bright as some believe, it may soon be possible to combine metal isotope ratios, light element stable isotope ratios, organic biomarkers, and trace element associations into a search engine for the neglected fossil record of ancient hydrothermal ecosystems. A record that parallels the well-established microbial fossil record from photosynthetic environments may reveal much about early ecosystems on the Earth and the type of evidence that could be sought in extraterrestrial samples, especially from planets where photosynthesis may not be an option (Shock, 1997).

5 THE FOSSIL RECORD OF HYDROTHERMAL ECOSYSTEMS Microfossils, biomarkers and other lines of evidence have pushed the geological record of life back at least 3.5 billion years. Almost all of that record comes from rocks selected for their likelihood of containing a record of photosynthetic life. This bias has excluded from study a huge inventory of hydrothermally altered and/or deposited samples that may contain evidence of chemosynthetic life that is distinct from the well-established fossil record of microorganisms. Meanwhile a suite of new analytical methods makes it possible to explore these samples for novel signatures of life. As a consequence, there is an enormous opportunity to devise new geochemical tracers for evidence of ancient life in hydrothermal deposits. Various microorganisms have been shown to accumulate metals within their cells, on cell walls or in associated biofilms (Schultze-Lam et al., 1996; Konhauser and Ferris, 1996; Farag et al., 1998; Fortin and Ferris, 1998; Ferris et al., 1999; Langley and Beveridge, 1999; Parmar et al., 2000; Southam et al., 2000). In some hydrothermal ecosystems chemosynthetic microbes contribute to these biofilms. As an example, biofilms that contain members of both the Aquificales and “Korarcheota” are present at Calcite Springs and elsewhere at Yellowstone (Reysenbach, et al., 1994; 2000). Our analyses of water and biofilm from Calcite Springs reveal the dramatic increases in metal concentrations that microbes and biofilms can achieve. Iron, cobalt, copper, chromium, vanadium, aluminum and molybdenum are all enriched by more that 100-fold in the biofilm. On the other hand, several elements do not seem to be enriched by much in the biofilm

6 THE EMERGENCE OF LIFE Evidence that the deepest lineages in the universal phylogenetic tree are hosted by hydrothermal ecosystems leads to the hypothesis that life emerged at high temperatures and has evolved to adapt to lower temperatures including those that we find “normal” (Pace, 1991; 1997; Shock et al., 1995; 1998; Shock, 1996; 1997). Testing this hypothesis involves interdisciplinary research combining molecular biology of hyperthermophiles, biogeochemistry of hydrothermal systems, and theoretical and experimental models of water/rock/organic/microbe interactions. Progress in theoretical models has revealed the potential for organic synthesis in hydrothermal systems (Shock and Schulte, 1998; Amend and Shock, 1998), and recent experimental results are promising (McCollom et al., 1999; 2001; Cody et al., 2000). Organic synthesis is necessary but not sufficient for constructing an emergence of life model. After all, organic compounds are common throughout much of the solar system (carbonaceous chondrites, interplanetary dust particles, icy satellites, comets, Kuiper belt objects), but appear to have abiotic origins. In fact, it may be that the emphasis placed on abiotic organic synthesis has mislead the search for the origin of life for many decades. In66

stead, it may be more productive to shift the emphasis to habitat. It is a familiar idea that organisms require certain habitats, and those are the conditions in which we expect to find them. On the other hand, researchers have seldom considered suitable habitats for the emergence of life. (Russel and Hall, 1997; Shock et al., 2000) One of the more compelling reasons for considering hydrothermal systems as such a habitat is that they supply so much that is conducive to life. Not only are there abundant forms of geochemical energy, reactive mineral surfaces, and a thermodynamic drive that stabilizes organic compounds, but there are steep potential gradients of many types. Furthermore, thermophilic microorganisms thrive at these conditions, and apparently they always have. Hydrothermal ecosystems appear to be the oldest ecosystems on the planet. Like life, they are simultaneously dynamic and persistent.

Amend, J.P. & Shock, E.L. 1998. Energetics of amino acid synthesis in hydrothermal ecosystems. Science 28 1 : 16591662. Amend, J.P. & Shock, E.L. 2001. Energetics of overall metabolic reactions in thermophilic and hyperthermophilic Archaea and Bacteria. FEMS Microbiology Rev. (in press). Amend, J.P., Amemd, A.C., & Valenza, M. 1998. Determination of volatile fatty acids in the hot springs of Vulcano, Aeolien Islands, Italy. Org. Geochem. 28: 699-705. Anbar, A.D., Roe, J.E., Barling, J. & Nealson, K.H. 2000. Nonbiological fractionation of iron isotopes. Science 288: 126-128. Ball, J.W., Nordstrom, D.K., Cunningham, K.M., Schoonen, M.A.A., Xu, Y.,& DeMonge, J.M. 1998 Water-chemistry nd on-site-sulfur-speciationdata for selected springs in Yellowstone National Park, Wyoming, 1994- 1995. U S Geological Surver Open File Report 98-574. Barns, S. M.; Fundyga, R. E.; Jeffries, M. W.; & Pace, N. R. 1994 Remarkable archaeal diversity detected in a Yellowstone National Park hot spring environment. Proc. Natl. Acad. Sci. U. S. A. 9 l(5): 1609- 13. Barns, S.M., Delwiche, C.F., Palmer, J.D. & Pace, N.R. 1996. Perspectives on archaeal diversity, thermophily and monophyly from environmental rRNA sequences. Microbiologjj 93: 9188-9193. Beard, B.L., Johnson, C.M., Cox, L., Sun, H., Nealson, K.H. 6r Aguilar, C. 1999. Iron isotope biosignatures. Science 285: 1889-1892. Brewer, W.H. 1866. On the presence of living species in hot and saline waters in California. Amer. Jour. Sci. 41: 391394. Brock, T.D. 1978. Thermophilic Microorganisms and Life at High Temperatures. Springer-Verlag, New York pp.465. Bult, C.J., White, O., Olsen, G.J., Zhou, L., Fleischmann, R.D., Sutton, G.G., Blake, J.A., FitzGerald, L.M., Clayton, R.A., Gocayne, J.D., Kerlavage, A.R., Dougherty, B.A., Tomb, J.F., Adams, M.D., Reich, C.I., Overbeek, R., Kirkness, E.F., Weinstock, K.G., Merrick, J.M., Glodek, A . , Scott, J.L., Geoghagen, N.S., & Venter, J.C. 1996. Complete genome sequence of the methanogenic archaeon, Methanococcus jannaschii. Science 273: 1058-73. Chen, J.M., & Hao, O.J. 1998. Microbial chromium (VI) reduction. Critical Rev. Env. Sci. Tech. 28(3): 2 19-25I . Chen, Y., Bonzongo, J-C,., J., Lyons, W.B., & Miller, G.C. 1997. Inhibition of mercury methylation iii anohic f r c d water sediment by group VI anions. E17v. Tosicdo,yi. C ’ h c , / ~ / 16(8): 1568-1574. Christensen, H., Hansen, M. & Sorensen, .I. 1999. Cotii1tii1; and size classification of active soil bacteria by fluorescence in situ hybridization with an rRNA oligonucleotide probe. Appl. Environ. Microbiol. 65(4): 1753-1761. Clement, B.G., Kehl, L.E., DeBord, K.L. & Kitts, C.L. 1998. Terminal restriction fragment patterns (TRFPs) a rapid, PCR-based method for the comparison of complex bacterial communities. Journ. Microbiology. Methods 3 1: 135-142. Clifton, C.G., Walters, C.C., & Simoneit, B.R.T. 1990. Hydrothermal petroleums from Yellowstone National Park, Wyoming, U.S.A. Appl. Geochem. 5: 169- 19 1.

ACKNOWLEDGEMENTS

I wish to thank many colleagues who have worked in my lab (GEOPIG) over the past fourteen years for many enlightening discussions, enthusiastic trips to various outcrops, and willingness to pioneer efforts in things that none of us had tried. Special thanks go to Tom McCollom, Mitch Schulte, Jan Amend, D’Arcy Meyer, Misha Zolotov, Gavin Chan, Andrey Plyasunov, Karyn Rogers and Melanie Summit who have all contributed one way or another to the ideas and insights presented here. Thanks also to Bill McKinnon, Bob Criss, Bob Osburn, Anna-Louise Reysenbach, Mike Russell, Jody Deming, John Baross, Mike Adams, Roy Daniels and Ha1 Helgeson for support and encouragement. REFERENCES Amann, R.I., Ludwig, W. & Schleifer, K.H. 1995. Phylogenetic identification and in situ detection of individual microbial cells without cultivation. Microbiol. Rev. 59: 143169. Amann, R. & Ludwig, W. 2000. Ribosomal RNA-targeted nucleic acid probes for studies in microbial ecology. FEMS Microbiology Rev. 24(5): 555-565. Amend, J.P. & Helgeson, H.C. 1997a. Group additivity equations of state for calculating the standard molal thermodynamic properties of aqueous organic species at elevated temperatures and pressures. Geochim. Cosmochim. Acta 6 1 : 1 1-46. Amend, J.P. & Helgeson, H.C. 1997b. Calculation of the standard molal thermodynamic properties of aueous biomolecules at elevated temperatures and pressures. Part 1. La-amino acids. J. Chem. Soc. Faraday Trcns. 93: 19271941.

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Coates, J.D., Michaelidou, U., Bruce, R.A. O’Connor, S.M. & Crespi, J.N. 1999 Ubiquity and diversity of dissimilatory (per)chlorate-reducing bacteria Appl. Environ. Microbiol. 65: 5234-5241. Cody, G. D.; Boctor, N.Z.; Filley, T. R.; Hazen, R. M.; Scott, J. H.; Sharma, A.; & Yoder, H. S., Jr. 2000 Primordial carbonylated iron-sulfur compounds and the synthesis of pyruvate. Science 289: 1337-1340. Copeland, J.J. 1936. Yellowstone thermal Myxophyceae. Ann. N. Y. Acad. Science 36: 1-229. Crump, B.C. & Baross, J.A. 2000. Archaeaplankton in the Columbia River, its estuary and the adjacent coastal ocean, USA. FEMS Microbiology Ecology 3 1 : 23 1-239. Dale, J.D., Shock, E.L., Macleod, G., Aplin, A.C. & Larter, S.R. 1997. Standard partial molal properties of aqueous alkylphenols at high pressures and temperatures. Geochim. Cosmochim. Acta 61 : 40 17-4024. DeLong, E.F., Taylor, L.T., Marsh, T.L. & Preston, C.M. 1999. Visualization and enumeration of marine planktonic archaea and bacteria by using polyribonucleotide probes and fluorescent in situ hybridization Appl. Environ. Microbiol. 65( 12): 5554-5563. Delaney, J. R.; Kelley, D. S.; Lilley, M. D.; Butterfield, D. A.; Baross, J. A.; Wilcock, W. S. D.; Embley, R. W.; & Summit, M. 1998 The quantum event of oceanic crustal accretion: impacts of diking at mid-ocean ridges. Science 281: 222-230. Deming, J. W.; & Baross, J. A.. 1993 Deep-sea smokers: windows to a subsurface biosphere? Geochinz. Cosmochim. Acta 57: 3219-3230. Dojka, M A . , Harris, J.K. & Pace, N.R. 2000. Expanding the known diversity and environmental distribution of an uncultured phylogenetic division of Bacteria. Appl. Environ. Microbiol. 66(4): 1617-1621. Dojka, M.A., Hugenholtz, P., Haack, S.K. & Pace, N.R. 1998. Microbial Diversity in a Hydrocarbon- and ChlorinatedSolvent-Contaminated Aquifer Undergoing Intrinsic Bioremediation. Appl. Environ. Microbiol. 64( 10): 38693877. Edgcomb, V.P., McDonald, J.H., Devereux, R. & Smith, D.W. 1999. Estimation of bacterial cell numbers in humic acidrich salt marsh sediments with probes directed to 16s ribosomal DNA. Appl. Environ. Microbiol. 65(4): 1 5 16- 1523. Farag, A.M., Woodward, D.F., Goldstein, J.N., Brumbaugh, W. & Meyer, J.S. 1998. Concentrations of metals associated with minin waste in sediments, biofilm, benthic macroinvertebrates, and fish from the Coeur d’Alene River basin, Idaho. Arch. Environ. Contam. Toxicology 34(2): 1 19-127. Ferris, F.G., Konhauser, K.O., Lyven, B. & Pedersen, K. 1999. Accumulation of metals by bacteriogenic iron oxides in a subterranean environment. Geornicrobiology 16(2): 18 1 192. Fortin, D. & Ferris, F.G. 1998. Precipitation of iron, silica, and sulfate on bacterial cell surfaces. Geomicrobiology 15(4): 309-324. Fournier, J-C., Soulas, G., & Parekh, N.R. 1997. Main microbial mechanisms of pesticide degradation of soils. In J. Tarradelles, G. Bitton & D. Rossel (eds.) 86-1 15 Soil Ecotoxicology Boca Raton:CRC Press.

Galtier, N., Tourasse, N. & Guy, M. 1999 A nonhyperthermophilic common ancestor to extant life forms. Science, 283: 220-22 1. Giovannoni, S.J. & Cary, S.C., 1993. Probing marine systems with robosomal RNAs. Oceanog. 6: 95-104. Haas, J.R. & Shock, E.L. 1999. Halocarbons in the environment: Estimates of thermodynamic properties for aqueous chloroethylene species and their stabilities in natural settings. Geochinz. Cosomochim. Acta 63: 3429-3441. Hanson, R.S., Tsien, H.C., Tsuji, K., Brusseau, G.A., & Wackett, L.P. 1990. Biodegradtion of low-molecular-weight halogenated hydrocarbons by methanotrophic bacterica. FEMS Microbiol. Rev. 87: 273-278. Helgeson, H.C., Delany, J.M., Nesbitt, H.W. & Bird, D.K. 1978 Summary and critique of the thermodynamic propcrties of rock-forming minerals, Amer. Jour. Sci., 278-A: 1229. Helgeson, H.C., Knox, A.M., Owens, C.E. & Shock, E.L. 1993. Petroleum, oil field waters and authigenic mineral assemblages: Are they in metastable equilibrium in hydrocarbon reservoirs? Geochim. Cosmochim. Acta 57: 3295-3339. Helgeson, H.C., Owens, C.E., Knox, A.M, & Richard, L. 1998 Calculation of the standard molal thermodynamic properties of crystalline, liquid, and gas organic molecules at high temperatures and pressures. Geochim. Cosmochim. Acta 62: 985-1081. Henson, M.J., Yates, M.V., & Cochran, J.W. 1989. Metabolism of chlorinated methanes, ethanes, and ethylenes by a mixed bacterial culture growing on methane. Jour. Industrial Microbiol4: 29-35. Hobman, J.L., Wildon, J.R., & Brown, N.L. 2000 Microbial mercury reduction In D.R. Lovley (ed) Environmental Microbe-Metal Interactions: 177-197.Washington: ASM Press Hugenholtz, P., Pitulle, C., Hershberger, K.L., & Pace, N.R. 1998a. Novel division level bacterial diversity in a Yellowstone hot spring. J. Bacteriol. 180: 366-376. Hugenholtz, P., Goegel, B.M. & Pace, N.R. 1998b. Impact of culture-independent studies on the emerging phylogenetic view of bacterial diversity. Jour. Bacteriology 180( 18): 4765-4774. Hugenholtz, P., Tyson, G.W., Webb, R.I., Wagner, A.M. & Blackall, L.L. 2001 Investigation of candidate division TM7, a recently recognized major lineage of the domain Bacteria with no known pure-culture representatives. Appl. Environ. Microbiol. 67( 1): 4 1 1-4 19. Jakosky, B.M. & Shock, E.L. 1998. The biological potential of Mars, the early Earth and Europa. Jour. Geophys. Res. 103: 19359-19364. Johnson, J.W., Oelkers, E.H., & Helgeson, H.C. 1992. SUPCRT92: A software package for calculating the standard molal properties of minerals, gases, aqueous species, and reactions from 1 to 5000 bar and 0 to 1000°C. Comp. Geosci. 18: 899-947. Kaye, J.Z. & Baross, J.A. 2000. High incidence of halotolerant bacteria in Pacific hydrothermal-vent and pelagic environments. FEMS Microbiology Ecology 32: 249-260. Kertesz, M.A., Cook, A.M., & Leisinger, T. 1994. Microbial met a bo I ism of s u I fur- and phosphor u s-co ii t a i n i ii 2 x c no 13i otics. FEMS Microbiol. Rev. 15: 195-2 15.

68

Konhauser, K.O.. & Ferris, F.G. 1996. Diversity of iron and silica precipitation by microbial mats in hydrothermal waters, Iceland: Implications for Precambrian iron formations. GeofogV 24(4): 323-326. Langley, S. & Beveridge, T.J. 1999. Metal binding by Pseudomonas aeruginosa PAOl is influenced by growth of the cells as a biofilm. Can. Jour. Microbiology 45(7): 616-622. Lee, N., Nielsen, P.H., Andreasen, K.H., Juretschko, S., Nielsen, J.P., Schleifer, K-H. & Wagner, M. 1999. Combination of fluorescent in situ hybridization and microautoradiography-a new tool for sturcture-function analyses in microbial ecology. Appl. Environ. Microbiol. 65(3): 1289-1297. Liu, W-T., Marsh, T.L., Cheng, H. & Forney, L.J. 1997. Characterization of microbial diversity by determing terminal restriction fragment length polymorphisms of genes encoding 16s rRNA. Appl. Environ. Microbiol. 63(1 I): 4516-4522. Lovley, D.R. 1993. Dissimilatory metal reduction. Annu. Rev. Microbiol. 47: 263-290. Lovley, D.R. 2000. Fe(II1) and Mn(1V) reduction. In D.R. Lovley (ed) Environmental Microbe-Metal Interactions: 332.Washington: ASM Press MacRae, I.C. 1989. Microbial metabolism of pesticides and structurally related compounds. Rev. Env. Contamination Toxicology 109: 1-87. Massana, R., DeLong, E.F. & Pedros-Alio, C. 2000 A few cosmopolitan phylotypes dominate planktonic archaeal assemblages in widely different oceanic provinces. Appl. Environ. Microbiol66(5): 1777-1787. McCollom, T.M. & Shock, E.L. 1997. Geochemical constrainsts on chemolithoautotrophic metabolism by microorganisms in seafloor hydrothermal systems. Geochim. Cosmochim. Acta 61: 4375-4391. McCollom, T.M., Ritter, G. & Simoneit, B.R.T. 1999 Lipid synthesis under hydrothermal conditions by FischerTropsch-type reactions Orig. Life Evol. Biosph. 29: 153166. McCollom, T.M., Seewald, J.S., & Simoneit, B.R.T. 2001. Reactivity of monocyclic aromatic compounds under hydrothermal conditions. Geochim. Cosmochim. Acta 65: 455468. Murray, A.E., Hollibaugh, J.T. & Orrego, C. 1996. Phylogenetic compositions of bacterioplankton from two California estuaries compared by denaturing gradient gel electrophoresis of 1 6 s rDNA fragments Appl. Environ. Microbiol. 62(7): 2676-2680. Muyzer, G.,de Waal, E.C. & Uitterlinden, A.G. 1993. Profiling of complex microbial populations by denaturing gradient gel electrophoresis analysis of polymerase chain reaction-amplified genes coding for 16s rRNA Appl. Environ. Microbiol. 59(3): 695-700. Oremland, R.S., Dowdle, P.R., Hoeft, S., Sharp, J.O., Schaefer, J.K., Miller, L.G., Blum, J.S., Smith, R.L., Bloom N.S. & Wallschlaeger, D. 2000. Bacterial dissimilatory reduction of arsenate and sulfate in meromictic Mono Lake, California. Geochim. Cosmochim. Acta 64( 18): 3073-3084. Ouverney, C.C. & Fuhrman, J.A. 1999. Combined microautoradiography- 16s rRNA probe technique for determination of radioisotope uptake by specific microbial cell types in situ. Appl. Environ. Microbiol. 65(4): 1746-1752.

Pace. N.R. 1991 Origin of life: Facing up to the physical setting. Cell 65: 53 1-533. Pace, N.R. 1997. A molecular view of microbial diversity and the biosphere. Science 276: 734-740. Pace, N. R. 2001. The universal nature 01’ biochcmi\ii-) Proc Natl. Acad. Sci. U.S.A. 98(3): 805-808. Parmar, N., Warren, L.A., Roden, E.E. & Ferris, F.G. 7000. Solid phase capture of strontium by the iron-reducing bacteria Shewanella alga strain BrY. Chem. Geol. 169: 28 1-288. Pennisi, E. 1998 Genome data shake the tree of life. Science 280: 672-674. Plyasunov, A.V. & Shock, E.L. 2000. Standard state Gibbs energies of hydration of hydrocarbons at elevated temperatures as evaluated from experimental phase equilibria studies. Geochim. Cosmochim. Acta 64: 281 1-2833. Ramsing, N.B., Kuhl, M. & Jorgensen, B.B. 1993. Distribution of sulfate-reducing bacteria, 0 2 , and H2S in photosynthetic biofilms determined by oligonucleotide probes and microelectrodes. Appl. Environ. Microbiol59( 1 I): 3840-3849. Ramsing, N.B., Ferris, M.J. & Ward, D.M. 2000. Highly ordered vertical structure on Synechococcus populations within the one-milimeter-thick photic zone of a hot spring cyanobacterial mat. Appl. Environ. Microbiol. 66(3): 10381049. Reysenbach, A-L., Wickham, G.S. & Pace, N.R. 1994. Phylogenetic analysis of the hyperthermophilic pink tilanient community in Octopus Spring, Yellowstone National Park. Appl. Environ. Microbiol. 60: 2 1 13-2 1 19. Reysenbach, A-L., Ehringer, M. & Hershberger, K. 2000. Microbial diversity at 83°C in Calcite Springs, Yellowstone National Park: another environment where the Aquzficales and “Korarchaeota” coexist. Extremophiles 4: 61 -67. Richard, L., & Helgeson, H.C. 1998 Calculation of the thermodynamic properties at elevated temperatures and pressures of saturated and aromatic high molecular weight solid and liquid hydrocarbons in kerogen, bitumen, petroleum, and other organic matter of biogeochemical interest. Geochim. Cosmochim. Acta 62: 3591-3636. Rondon, M.R., August, P.R., Bettermann, A.D., Brady, S.E., Grossman, T.H., Liles, M.R., Loiazono, K.A., Lynch, B.A., MacNeil, LA., Minor, C., Tion, C.L., Gilman, M., Osburne, M.S., Clardy, J., Nandelsman, J. & Goodrnan, R.M. 2000. Cloning the soil metagenome: a strategy for accessing the genetic and functional diversity of uncultured microorganisms. Appl. Environ. Microbiol. 66(6): 2541 -2547. Russell, M.J. & Hall, A.J. 1997. The emergence of life from iron monosulphide bubbles at a hydrothermal redox front. Jour. Geol. Soc., 154: 377-402. Schramm, A., Larsen, L.H., Revsbech, N.P., Ramsing, N.B., Amann, R. & Schleifer, K-H. 1996. Structure and function of a nitrifying biofilm as determined by in situ hybridization and the use of microelectrodes. Appl. Env. Microbiol. 62( 12): 4641-4647. Schultze-Lam, S., Fortin, D., Davis, B.S. & Beveridge, T.J. 1996. Mineralization of bacterial surfaces. Chem. Geof. 132: 171-181. Setchell, W.A. 1903. The upper temperature limits of life. Science 17: 934-937.

69

Shock, E.L. 1988. Organic acid metastability in sedimentary basins. Geology 16: 886-890. Shock, E.L. 1989. Corrections to “Organic acid metastability in sedimentary basins.” Geology 17: 572-573. Shock, E.L. 1994. Application of thermodynamic calculations to geochemical processes involving organic acids. In: The Role of Organic Acids in Geological Processes (eds. M. Lewan & E. Pittman) Springer-Verlag pp. 270-3 18. Shock, E.L. 1995. Organic acids in hydrothermal solutions: Standard molal thermodynamic properties of carboxylic acids, and estimates of dissociation constants at high temperatures and pressures. Amer. Jour. Sci. 295: 496-580. Shock, E.L. 1996. Hydrothermal systems as environments for the emergence of life. In G. R. Bock & J. Goode Evolution of Hydrothermal Ecosystems on Earth (and Mars?) 40-60 Chichester:Wiley. Shock, E.L. 1997. High temperature life without photosynthesis as a model for Mars. Jour. Geophys. Res. 102: 23,68723,694. Shock, E.L. & Helgeson, H.C. 1990. Calculation of the thermodynamic and transport properties of aqueous species at high pressures and temperatures: Standard partial molal properties of organic species. Geochim. Cosmochim. Acta 54: 915-945. Shock, E.L. & Schulte, M.D. 1998. Organic synthesis during fluid mixing in hydrothermal systems. Jour. Geophys. Res. 103: 28513-28527. Shock, E.L., Amend, J.P. & Zolotov, M.Yu. 2000. The early Earth vs. the origin of life. In R. Canup and K. Righter (eds.), The Origin of the Earth and Moon. 527-543 Tucson:University of Arizona. Shock, E.L., McCollom, T. & Schulte, M.D. 1995. Geochemical constraints on chemolithoautotrophic reactions in hydrothermal systems. Origins of Life and Evolution of the Biosphere 25: 141-159. Shock, E.L., McCollom, T. & Schulte, M.D. 1998. The emergence of metabolism from within hydrothermal systems. In: J. Wiegel & M. Adams (eds.) Thermophiles: the keys to molecular evolution and the origin of life? 59-76London: Taylor & Francis. Shock, E.L., Sassani, D.C., Willis, M. & Sverjensky, D.A. 1997. Inorganic species in geologic fluids: Correlations among standard molal thermodynamic properties of aqueous ions and hydroxide complexes. Geochim. Cosmochim. Acta 6 1 : 907-950. Shock E.L., Amend, J.P, Meyer, D., Rogers, K., Osburn, G.R. Chan, G. & Fischer, T. Geochemistry of hydrothermal ecosystems at Yellowstone National Park. (in prep). Southam, G., & Beveridge, T.J. 1994. The in vitro formation of placer gold by bacteria. Geochim. Cosmochim. Acta 58(20): 4527-4530. Southam, G., Fyfe, W.S. & Beveridge, T.J. 2000. Inimobilization of free ionic gold and L-asparagine-complexed ionic gold by Sporosarcina ureae: the importance of organo-gold complexes in gold mobility. Miner. Metnll. Process 17(2): 129- 132. Stolz, J.F. & Oremland, R.S. 1999. Bacterial respiration of arsenic and selenium. FEMS Microbiol. Rev. 23: 61 5-627.

Summit M. and Baross, J.A. 2001 A novel microbial habitat in the mid-ocean ridge subseafloor. Proc. Natl. Acad. Sci. U.S.A. 98 (in press). Suzuki, M.T., Taylor, L.T. & DeLong, E.F. 2000. Quantitative analysis of small-subunit rRNA genes in mixed microbial populations via 5 ’ -nuclease assays. Appl. Environ. Microbiol. 66(11): 4605-4614. Takai, K. & Horikoshi, K. 2000. Rapid detection and quantification of members of the archaeal community by quantitative PCR using fluorogenic probes Appl. Environ. Microbiol. 66( 1 I): 5066-5072. Tanner, M.A., Goebel, B.M., Dojka, M.A. & Pace, N.R. 1998. Specific ribosomal DNA sequences from diverse environmental settings correlate with experimental contaminants. Appl. Environ. Microbiol. 64(8): 3 I 10-3 1 13. Teske, A., Brinkhoff, T., Muyzer, G., Moser, D.P., Rerhmeicr. J. & Jannasch, H.W. 2000. Diversity of ttiiosulfatcoxidizing bacteria from marine sedimeiits and hydrothermal Vents. Appl. Env. Microbiol. 66(8): 3 125-3 133. TIGR, 2001 The Institute for Gemonic Research Databases,

btt&h!FELi gr,s rg/t_db/. Wang, Y.-T. 2000 Microbial reduction of chromate. In D.R. Lovley (ed) Environmental Microbe-Metal Interactions: 225-235.Washington: ASM Press. Ward, D.M., Bateson, M.M., Weller, R., & Ruff-Roberts, A.L., 1992. Ribosomal RNA analysis in microorganisms as they occur in nature. Adv. Micro. Ecol. 12: 2 19-286. Whitman, W.B., Coleman, D.C. & Wiebe, W.J. 1998. Prokaryotes: The unseen majority. Proc. Natl. Acad. Sci. U. S. A. 95: 6578-6583. Woese, C. R.. 2000 Interpreting the universal phylogenetic tree. Proc. Natl. Acad. Sci. U. S. A . 97( 15): 8392-8396. Woese, C.R., Kandler, O., & Wheelis M.L. 1990 Towards a natural system of organisms: Proposal for the domains Archaea, Bacteria, and Eucarya. Proc. Natl. Acad. Sci. U. S. A 87: 4576-4579. ZoBell, C.E. 1958. Ecology of sulfate reducing h a c [ e r i a / ’ / . [ I ducers Monthly Penn. Oil Prod. A.ss.22: 12-29.

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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Exploring the sources of the salinity in the Middle East: an integrative hydrologic, geochemical and isotopic study of the Jordan River Avner Vengosh & Efrat Farber Department of Geological and Environmental Sciences, Ben Gurion University,Beer Sheva, Israel

Uri Shavit, Ran Holtzman & Michal Segal Faculty of Agricultural Engineering, Technion,Israel Institute of Technology,Haifia, Israel

Ittai Gavrieli Geological Survey of Israel, Jerusalem, Israel

ECO-Research Team Amman, Jordan

Thomas Bullen Water Resources Division, U S . Geological Survey, Menlo Park, California, USA

ABSTRACT: Continued stress on limited water resources in the Middle East has led to significant drawdown of water levels and rapid degradation of water quality. Increasing salinity clearly endangers future exploitation of regional water resources. Here we describe an ongoing study that involves researchers from Israel, Jordan and the Palestinian Authority, in which we examine the sources of salinity along the Lower Jordan River. In the past, the river received considerable fresh water from Lake Tiberias, the Yarmouk River and local runoffs. Currently, a much smaller flow rate of mostly poor quality fluids enters the river, leading directly to the degradation of water quality. Using a variety of diagnostic geochemical tracers for elucidating salinity sources, we have identified three principal zones along the river in which salinity is dramatically modified due to both surface water and groundwater inflows. The influence of groundwater on water quality of the river was previously unrecognized, and adds an additional constraint on future management policy. Future reductions of surface fresh water inflows will further increase the influence of the saline groundwater.

1 INTRODUCTION

Vengosh & Rosenthal 1993, Salameh 1996,Vengosh et al. 1999, Marie & Vengosh 2001). In addition, human activity produces low-quality fluids that enter the aquifer and further degrade water quality. These fluids include those derived from sewage, wastewater irrigation, landfills, and agricultural return flows. The superposition of natural and anthropogenic contaminants that affect water quality provides a scientific challenge and requires unconventional interpretations for the origin of the salinity. In this study we focus on the Jordan River, which is one of the most symbolic resources in the Middle East and the center of the peace treaty between Israel and Jordan. We show that integration of a wide range of geochemical and isotopic techniques is a useful methodology for delineating the different sources of the salinity. This scientific evaluation is essential for modeling future scenarios of river management and understanding the consequences of human activities on water quality. Since water sharing is one of the key components of the peace treaty between Israel and Jordan, its implementation and compliance can only be achieved through a deep understanding of the different processes affecting water quality in the Jordan River.

Water is one of the most valuable natural resources in the Middle East. The combination of population growth, economic and agriculture development, and arid climate with insufficient precipitation results in over exploitation of the water resources in the region. The continued stress leads to rapid degradation of the quality of fresh water resources due to salinisation and contamination process (e.g., Vengosh & Rosenthal 1994, Salameh 1996). The lack of sufficient water combined with rapid water quality deterioration presents a serious challenge to the people in the region. In order to be able to manage and share the water resources in conditions of accelerating degradation, it is crucial to understand the origin and mechanisms of the contamination process. The salinity that threatens the fresh water resources is derived from different sources, both natural and anthropogenic. In general, overexploitation of fresh water resources results in a rapid decrease of water level, which then triggers lateral and under flow of saline waters from adjacent aquifers. Consequently, the overexploited aquifer become saline due to mixing with saline waters (e.g., 71

accumulates local runoff and flows through the Upper Jordan River into Lake Tiberias (210 m bsl). The Lower Jordan River starts at Alumot, downstream from Lake Tiberias, and ends at the Dead Sea in the south (410 m bsl). The river symbolizes the history of the region. Starting with the Israelites crossing the river and continuing with the Prophets, Elijah, Elisha, John the Baptist and Jesus Christ all crossed the river in their lifetimes. At present, the lower Jordan River serves as an international border. The quality and quantities of water delivered by the Lower Jordan River have been extremely degraded during the last several decades. Since the implementation of water supply projects in Israel, Jordan, and Syria, no fresh surface water flows into the river except for negligible springs and rare flood events. As a result, the available water sources are limited to artificial deviation of saline springs from Lake Tiberias (“the saline carrier”), natural flows from adjacent saline springs, dumping and leakage of solid and liquid wastes, effluents from fish ponds, Figure 1. General location of the Jordan River and agricultural return flows from adjacent fields. The total discharge of the river into the Dead Sea in the past was about 1200 MCM/year (Klein 1998, 2 THE JORDAN RIVER BASIN Salik 1988, Dalin 1982, Sofer 1994). The amount has now declined to a mere 100-200 MCWyear The Jordan River basin is part of the Dead Sea rift (Tahal2000). valley, an active geological depression with The river is a resource shared by all peoples in elevations between 210 to 410 m bsl. Since the late the region. As such, it received a great deal of Miocene, several invasions of Mediterranean Sea attention in the peace treaty between Israel and water formed lakes that on evaporation resulted in Jordan (October 1994). Based on the agreements the salt deposits of Mt. Sedom (Zak 1967). The between the two countries and the desire to develop chemical, boron and strontium isotopic compositions the regional environment, changes in the operation of brines and saline springs that are found of the river and its surroundings are expected to take throughout the Rift Valley reflect chemical place in the near future. These changes include the evolution and modifications of the original desalination of saline water and treatment of sewage evaporated sea water (Starinsky 1974). These and other waste fluids. According to the water include sea water evaporation and salt crystallization balance calculation of Al-Weshah (2000), the total (e.g., low NdC1 and high Br/C1 ratios), gypsum discharge of the river into the Dead Sea is 175 precipitation and sulfate reduction (low sulfate), and MCM/year. If peace treaty allocations are included, extensive water-rock interactions such as the discharge is reduced to only 60 MCM/year. dolomitization (Ca enrichment and low 87Sr/86Sr The Jordan River rift valley serves as a base level ratios) and clay adsorption (low B/Li and high 6“B to which surface- and ground water drains from both values) (Starinsky 1974, Stein et al. 2000, Vengosh east and west. Shallow ground water represents an et al. 1991). The last phase of the fluviatile episode, interesting mixture of a variety of end members. 100 to 20 Ma BP, was the formation of the long and Although the expected future operational narrow Lake Lisan, a hypersaline water body with a modifications will affect the hydrology and sharp density stratification of freshwater overlying chemistry of the shallow ground water system, we hypersaline brines (Katz & Kolodny 1989, Stein et consider it as the more stable system among all other al. 1997). Since then, large variations occurred water systems around the river. Better understanding including extensive evaporation, level fluctuations of this complex hydrological and operational system (500-180 m below sea level) and salinity changes is necessary in order to predict the effect of the (Yechieli 1993). Currently the Jordan River flows expected future changes. through the Lisan Formation that is comprised of Figure 2 shows the Lower Jordan River region mark and gypsiferous sediments. and the Yarmouk River which marks the border The Jordan River is the largest river in the region between Israel, Jordan, and Syria. At present, the (Fig. 1). It originates from three sources, the Dan, outlet of Lake Tiberias and the Yarmouk are blocked Banyas, and Hasbani springs, from which it by dams. 72

Figure 2 . A map of the sampling sites along the Jordan River from Sea of Galilee and Alumot Dam in the north (left map) to the Dead Sea in the south (right map). Numbers refer to the sampling sites.

Water from saline springs at the shore of Lake Tiberias is carried by “the saline carrier” to the starting point of the Lower Jordan River at Alumot Dam. A total of 20 MCWyear of saline water and about 7 MCM/year of sewage is the initial discharge of the river. The total agricultural irrigated area that drains back to the river is estimated as 440 km2; 170 km2 on the west side and 270 km2 on the east. The total water consumption for irrigation was estimated to be 400 MCWyear on the east side (Salameh 1996) and 150 MCM/year on the west side (Tahal 2000). An analysis of the agricultural influence on the river mass balance is a major challenge of the current study. The above studies suggested that about 15% of irrigation water (i.e., 80 MCWyear) flows back to the river. This estimate will be tested with our results.

3 METHODOLOGY Since September 1999 we have made several trips along the river and vicinity. During these trips sample sites were identified and water samples were collected (Fig. 2). The fieldtrips included meeting with the local water authorities, professionals, consultants, and administrators. The trips to the Jordan River were conducted together with the Israeli Nature Reserve Authority (INRA). On the basis of this framework we have established a list of accessible sites, all coordinated with the military authority. Similarly, the Jordanian team identified and selected 15 sites for representing eastern inflows and groundwater discharge to the river system. It should be emphasized that the last hydrological year represents the second of two consecutive drought 73

MCM/year) and sewage effluents (6- 10 MCWyear) at the starting point of the river; and (2) an unknown constant inflow. The chemical variations suggest that the fraction of the second component gradually increases with distance (Fig. 3). Geochemical variations constrain the source of this second component. By simple mass balance calculations we compare the chemical composition of all of the identified inflows and the actual changes in the composition of the river water. Our calculations show that the best fit composition that is consistent with the chemical modification of the river is that of the Yarmouk River (Table 1). In contrast, the western surface inflows have relatively low S04/Cl, NdC1, Mg/Cl, and B/Cl as well as high CdC1 ratios (dashed line in Fig. 4) that are not consistent with our mass-balance calculations. Hence, the chemical shift observed in the upper section cannot be derived from the western surface inflow. Moreover, the fresh water inflows from the east (e.g., Wadi Arab) or even a mixture of western saline and eastern fresh water inflows are not able to explain the chemical variations as the dissolved salts are too low and are not consistent with the increase of Mg, SO4, and B contents. We expect that during winter floods this salinity would be even lower as indicated by the flood flow composition along wadis in the Jordan Valley (Salameh 1996). Consequently, we argue that the only measured inflow source that can affect the chemical composition of the upper 20 km of the Jordan River is that of the Yarmouk River. It should be emphasized that water quality along the Yarmouk River changes dramatically above and below Addasia dam. The quality of the upstream water is high (C1=135 mg/l) whereas the quality of downstream section is low (C1=840-1200 mg/l). This results from a combination of extensive upstream use of the river and diversion of the Yarmouk water to King Abdalla canal (-120 MCM/year) and Sea of Galilee (25 MCM/year) as part of the peace treaty between Israel and Jordan. Instead of the original fresh water, the Yarmouk River below Addassiyah Dam receives low quality effluents from local communities, agriculture return flow and fish pond effluents that control the unique chemical composition of the Yarmouk River. It seems that this chemical composition also affects the Jordan River along its upper 20 km. The gradual transition suggests lateral inflow of groundwater that has been contaminated by these types of effluents. The irnportance of Yarmouk River underflow is strengthened by the fact that the chemistry of the Jordan River is modified towards “Yarmouk River type water” before the actual entrance of the Yarmouk River to the Jordan River. The s7Sr/86Sr ratio (Fig. 5) decreases from 0.70775 to 0.70763 from 0 to 20 km downstream.

years and thus the samples represent base flow with minimal contribution of runoff. Moreover, the gates at the Dagania (Sea of Galilee) and Addassiyah (Yarmouk River) dams did not release a single drop of upstream fresh water to the Jordan River. Water samples were analyzed for major chemical constituents at the laboratory of the Israel Geological Survey and University of Amman. Strontium was separated by ion-exchange columns and the isotopic composition was measured using a MAT-261 mass spectrometer at the laboratory of US Geological Survey, Menlo Park, California. Boron isotopes were measured by negative thermal ionization mass spectrometry (Vengosh et al. 1989, 1999) using a MAT-261 at US Geological Survey, Menlo Park. Oxygen isotope ratio measurements were made on a VG SIRA-I1 mass spectrometer at the Geological Survey of Israel. Results for oxygen and boron isotopes are given in per mil values with respect to SMOW (Craig 1961) and NBS-SRM 951 standards, respectively. Analytical reproducibility of duplicates and replicate analyses are 0.1%0, 1%0,and 0.025%0 for 0, B, and Sr isotopes, respectively. 4 RESULTS AND DISCUSSION The chemical and isotope data of the Jordan River, between Alumot Dam and the Dead Sea, reveal three geographically distinct zones (Fig. 3): the upper (the first 20 km), central (20-65 km), and lower (below 70 km) sections. Table 1 summarizes the main geochemical characteristics of the Jordan River and selected identified inflows. 4.1 The upper zone The first 20 km of the river, below Alumot dam, is characterized by a gradual decrease of chloride and sodium and increase of Mg, so4, and B concentrations. This gradual modification is superimposed with the influences of local inflows. One conspicuous example of the influence of the surface water is the inflow of the fresh water of Wadi Arab (base flow during September with C1=174 mg/l) 12 km downstream, which reduces the salinity of the Jordan River. Nevertheless, these inflows have only a minor affect on the overall chemical composition of the Jordan River. The initial flow is characterized by high C1 (up to 2600 mg/l) and Ca and Na, and low Mg, SO4, and B concentrations with relatively low molar ratios of NdC1 (0.68), Mg/C1 (0.07), S0&1 (0.03), and B/C1 ( 8 ~ 1 0 -and ~ ) high CdCl(O.16) ratio (Table 1). The gradual changes and linear relationships in Figure 4 (zone A) reflect mixing of two components: (1) initial saline fluids that are derived from the artificial inlet of a blend of the “saline carrier” (20 74

c1 (mg/l)

so4 (mg/l)

2100-2500 1470-1600

170-185 440-530

0.66-0.70 0.75-0.88

0.03 0.10-0.13

1730-1960 1900 1370 1560

200 240 180 230

0.71-0.78 0.64 0.62 0.65

840-1200 174 108

580-850 100 80

1300-1670 1300-1570

B/CI

87Sr/86Sr 6"B

0.07 0.15

0.8 1.4-2.0

0.70775 0.70766

30 30

0.04 0.05 0.05 0.05

0.12 0.18 0.19 0.14

0.7 0.5 0.8 0.7

0.70782 0.70791 0.70785 0.70783

39.5 38 36 43

0.91-0.96 0.85 1.oo

0.23-0.26 0.2 1 0.26

0.27 0.29 0.46

5.3-6.0

0.70719

36.5

370-420 370-410

0.76-0.87 0.74-0.76

0.08-0.1 1 0.10-0.1 1

0.15-0.18 0.17-0.19

1.4-1,6 1.5-2.0

0.70771 0.70785

31.5-33 36.7

1330-1500 430-450

330-360 175

0.69-0.73 0.80-0.86

0.08-0.10 0.15

0.10-0.13 0.29

1.5-2.2 2.0-2.6

0.70776 0.70798

38 47.5

330 1500

160 470

0.95 0.80

0.18 0.11

0.25 0.14

River after 66 km

1430-1680

500-715

0.75-0.93

0.1 1-0.16

0.7-0.23

1.8-3.1

31

River after 76 km River after 9 1 km

1600-2300 1650-2400

670-970 680-1030

0.77-0.79 0.73-0.80

0.15 0.14-0.17

0.18-0.19 0.16-0.19

2.1-2.7 2.4-3.0

33 31

Rifter after 96 km

1740-2200

560-900

0.76-0.80

0.1 1-0.17

0.17-0.19

2.3-2.8

0.708150.70829 0.70814 0.708090.70820 0.70813

38,000 2250-2550 14,200

1800 890-1020 680

0.57 0.68-0.70 0.67

0.02 0.15 0.02

0.15 0.21 0.11

0.4 2.8-3.6 2.3

0.70796 0.70797

41.7 41.7

Na/Cl

Mg/CI

SOdCl

Upper Jordan Initial river River after 20 km

Western inflows Harod inflow Nimrod inflow Chanal 17 Hogla springs

Eastern inflows Yarmouk River Wadi Arab Wadi Ziqlag

Central Jordan River after 27 km River after 44 km

Western inflows Wadi Al-Malich Sukot spring

Eastern inflows Abu Thableh Zoor Tbdulla

Southern Jordan

31

Western inflows

I

Wadi Al-Ahemar Uga Tirtza well

' I I

I

I

I

1

I

Eastern inflows Wadi Makman Maliah Gdeida Zarqa River Rasif Aaraa

I

820 960 1360 4830 30,500

1340 1300 1120 3040 2330

1.o 0.8 1 0.80 0.87 0.56

0.60 0.50 0.30 0.23 0.03

0.45 0.38 0.23 0.23 0.22

Legend: 1. Molar ratios. 2. (x 1o-~). 3. Values reported in per mil (%o), 611B=[{(11B/'oB),,,,,,~,/(11B/10B)~~~951 } - 11 x1000. 4. Values represent sampling between September 1999 to September 2000. 5 . Sampling on March, May and August 2000. 6. Sampling on May 2000. 7. Sampling on September 2000. 8. Sampling on September 1999, March and May 2000. 9. March and May 2000.

75

I

-

I

-

I

Figure 3. Chloride variations along the Jordan River between September 1999 and September 2000. Note the three major salinity zones along the flow of the Jordan River. Distance in km is referenced to y coordinate rather than actual river length from its beginning at Alumot Dam.

The groundwater and western surface inflows along the upper 20 km have significantly higher s7Sr/87Sr ratios (0.7078 to 0.7091), which cannot account for the isotopic shift. In contrast, the Yarmouk River (0.707 16) and shallow groundwater below a fishpond (0.70741) have isotopic ratios that are consistent with the isotopic modification of the river. It seems that the anthro ogenic groundwater component has a low 87Sr/H:Sr signature that is different from that of local western saline springs with higher 87Sr/87Srratios. In the upper zone the gradual decrease in salinity is associated with a general (although with large fluctuations) increase in 6l80 values. In contrast, the 6l80 values of springs and observed runoff are low (-4%o). Hence, the oxygen isotopic modification is also inconsistent with western inflows. We observed extremely high 6l80 values in fishponds and in shallow groundwater below the fishpond. The large fluctuations of the 6"O values probably reflect both inflows of "0-enriched groundwater superimposed with surface evaporation. Assuming that the chemical compositions of conservative elements reflect mixing of the upstream (Alumot Dam) and Yannouk River type, we calculate the fraction (F) of the initial solute that is mixed along the first 20 km flow of the Jordan River by:

F = (Cmix - CYarmouk) / (Cinitial - CYarmouk) (1)Whereas Cmix is the concentration of conservative constituent in the Jordan River, Cymouk is the concentration in the Yannouk River type, and Cinitial is the original concentration at Alumot Dam. Our 76

Figure 4. Chloride versus sulfate and magnesium concentrations of the Jordan River (circles), western inflows (open triangles), eastern inflows (closed triangles, measured only in September 2000), and groundwater in the vicinity of the Jordan River (squares). Zone A represents the upper 20 km whereas zone C is southern section of the River. The dashed line represent mixing between the initial river at Alumot Dam and the Yarmouk River.

calculation for C1, SO4, Mg, and s7Sr/s7Srvariations show that the fraction of the original solute starting from Alumot Dam is gradually reduced to about 30 to 50% 20 km downstream. Hence the water of the Jordan River is significantly replaced by groundwater and/or agriculture return flows that is heavily controlled by human influence. In a parallel study (Shavit et al. 2001) we show a significant increase in flow rates at the southern section of the upper zone. Considering the water volume pumped out from the river and the negligible volume entering the river through its tributaries (e.g. Wadi Arab - 13 L/s), we suggest that a large volume of water enters the river through unmonitored inputs. Although a complete water mass balance was not possible at the present time, the significant increase in flow rate indicates that the subsurface contribution, either directly or through the local drainage system, is very significant. Based on these results Shavit et al. (2001) estimate the subsurface

section, 20 km below Alumot Dam, thus excluding these inflows as a major source. In contrast, the western inflows of Wadi A1 Malich further south and eastern inflows (Abu Thableh, Zoor Tbdulla; Table 1) have chemical compositions that are similar to that of the Jordan River (Table 1). Low salinity of the central Jordan River is apparently associated with inflow of low-saline water, mainly fiom the eastern side. The lack of major chemical changes along a large section of the river suggests that no other sources affect water quality. Hence, the quality of the central section of the Jordan River seems to be controlled by surface inflows. 4.3 The southern zone The beginning of this section varies with time; an increase of the salinity begins at a distance of 45 km below Alumot dam during the winter period, and moves southward (66 km) during the summer. The linear relationships between C1, SO4, and Mg concentrations of the southern Jordan River (Fig. 4; zone C) clearly reflect mixing processes. The salinity increase with distance (Fig. 3) suggests continue contribution of a high salinity source. Similar to the upper section, we explain this gradual change by continuous input of shallow groundwater from the Lisan formation rather than from individual surface inflows such as the Zarqa River, which is one of the largest inflows in this area. We have measured the chemical and isotopic (only the western side at this stage) of identified inflows and springs in the vicinity of the southern Jordan River. Our data suggest two types of saline water at both sides of the river (Table 1): (1) Hypersaline brines found in Wadi A1 Ah’mar (C1 = 38,000 mg/l), Tirza well (14,100 mg/l), and Aqraa (30,500 mg/l; eastern side) with typically low NdC1 and S04/Cl ratios. (2) Saline water with a C1 range of 1000 to 4800 mg/l and typically higher NdC1 and S04/Cl ratios. It seems that the major shift of the chemical composition of the lower Jordan River is derived from inflow of the second component. The rise in the salinity in the lower section is associated also with an 87Sr/87Sr increase (0.7081 to 0.7083), although we observed a slight decrease with further distance. This value is consistent with the g7Sr/87Sr ratios measured in saline inflows in this region (0.70796-0.70800). In the southern section the 6l80 values show a wide range of -5%0 to -1.5%0. A large range of 6l8O values was also observed in the western surface inflows while groundwater show typically lower 6l80 values. It seems that the wide range of 6l80 values in the Jordan River reflect alternate

Figure 5. x7Sr/x6Srvariations of the Jordan River as sampled on September 1999 (circles) and May 2000 (squares). Also include the ratios of western inflows (open triangles) and groundwater (lar e s uares) along the river. Note the slight decrease of the Sr/’ Sr ration in the upper 20 km and the high ratios towards the southern section of the river.

8 %

contribution in this region to be around 100 L/s per km (y coordinate rather than actual river length). The gradual chemical variations suggest a solute exchange of 2.5% to 3.5% per km (i.e., the rate of transition from the original solute to a 30% to 50% mixed-solution down stream). As mentioned above, our chemical data may be biased as our samples represent base flow in the second of two consecutive draught years. We expect that flood events would drastically change the hydrological balance of the river. Nevertheless, the chemical composition of the river, which is controlled mainly by the saline sources, would be less affected by low saline floodwaters. 4.2 The central zone The central zone of the Jordan River show low salinity level relative to the upper and lower sections (Fig. 3). Our results show that the geographical location of this low salinity changes with time. During winter (March, May) the low salinity occurs between 20 to 45 km, whereas in summer (August, September) the low salinity section stretches between 20 to 65 km. This phenomenon is consistent with the distribution of C1, Mg, Ca, and Na. However, The SO4 variations show only low levels along the 20 to 45 km interval and are not sensitive to this seasonal variations. The chemical composition (e.g., SOdC1, NdCl) of the central Jordan River is significantly different from that of the saline springs and western inflows (e.g., Hogla springs; Table 1) at the beginning of this 77

reduction of surface fresh water inflows would further decrease the dilution factor and the impact of the saline subsurface flow (Cl-1000 mg/l in the north, C1>3000 mg/l in the south) would further dominate the quality of the Jordan River.

contributions of groundwater (low 6l8O) and surface water, in addition to evaporation on the Jordan River itself in the arid environment of the southern Jordan River.

4.4 Identijkation of end-members by strontium and boron isotopes

REFERENCES

We use the isotopic compositions of strontium and boron to define the origin of different sources that affect the water quality of the river. The strontium isotope variations (Fig. 5) enabled us to reveal the influence of subsurface flows in the upper (an 87Sr/87Srdecrease from 0.70775 to 0.70763) and lower (an 87Sr/87Srincrease to 0.7081 to 0.7083) sections. While the low 87Sr/87Srratios in the north are attributed to anthropogenic sources, the higher values in the south reflect the isotopic composition of natural saline groundwater flows. The 6"B values of the Jordan River show relatively constant values of 29 to 33%0 (except one sample with 37%0).These values are in contrast to the relatively higher 6I1B values of the western inflows and groundwater. It also contradicts the large elemental boron variations that follow those of C1 changes along the different sections of the river. The relatively constant low 611B values of the Jordan River can be explained by (1) desorption processes that take place in the hyporheic zone and/or (2) overlaps in 6l'B signatures of the different salinity sources. The high correlation between elemental B and C1 favors the latter explanation. Nevertheless, the inconsistency between the measured 611B values of the apparent inflow saline sources, particularly in the southern section (6"B =40 to 43%0)and the low 611B of the Jordan River is not resolved.

Al-Weshah, R.A. 2000. The water balance of the Dead Sea: an integrated approach. Hydrological Process 14: 145-154. Craig, H. 1961. The isotopic geochemistry of water and carbon in geothermal areas. In: Nuclear Geology on Geothermal Areas, E. Tongiorgi, ed. Consiglio Nazionale Della Ricerche, pp. 17-53. Dalin, Y. 1982. Assessment of the expected floods to the Dead Sea. In: The Hydrology and the Energy Crisis. In: Proc. Ministry of Energy and Infrastructure, 5-24. Epstein, S. & T. Mayeda 1953. Variation of "0 content of waters from natural sources, Geochimica et Cosmochimica Acta4: 213-224. Katz, A. & N.Kolodny 1989. Hypersaline brine diagenesis and evolution in the Dead Sea- Lake Lisan system (Israel). Geochimica et Cosmochimica Acta 53: 59-67. Klein, M. 1998. Water balance of the Upper Jordan River Basin. Water International 23: 244-248. Marie, A. & A. Vengosh (2001). Sources of sabity in ground water from Jericho area, Jordan Valley. Ground Water 39: (in press). McCafferey, M.A., Lazar, B. & H.D. Holland 1987. The evaporation path of seawater and the coprecipitation of Brand K- with halite. Journal Sedimentary Petrology 57: 928937. Salameh, E. 1996. Water Quality degradation in Jordan. Royal Society for the Conservation of Nature, Amman, Jordan. Salik, D., 1988. The lower Jordan River. Horizons 25-26: 99110. Shavit, U., Holtzman, R., Segal, M., Vengosh, A., Farber, E., Gavrieli, I., ECO RT & T. Bullen 2001. Water sources and quality along the lower Jordan River, a regional study. To be presented in: 4th Symposium cum Industrial Forum, Preserving the Quality of our Water Resources, Vienna (Austria), 23-25 April 2001. Sofer, A. 1992. Rivers of Fire. Am Oved, Israel. Starinsky, A. 1974. Relation between Ca-chloride brines and sedimentary rocks in Israel. Ph.D. Thesis, Hebrew University, Jerusalem, Israel (in Hebrew). Starinsky, A., Katz, A. & D. Levitte 1979. Temperaturecomposition-depth relationship in rift valley hot springs, Hammat Gader, northern Israel. Chemical Geology 27: 233244. Stein, M., Starinsky, A., Katz, A., Goldstein, S.L., Machlus, M. & A. Schramm 1997. Strontium isotopic, chemical, and sedimentological evidence for the evolution of Lake Lisan and the Dead Sea. Geochimica et Cosmochimica Acta 6 1 : 3975-3992. Stein, M., Starinsky, A., Agnon, A., Katz, A., Raab, M., Spiro B. & I. Zak 2000. The impact of brine-rock interaction during marine evaporite formation on the isotopic Sr record in the oceans: Evidence from Mt. Sedom, Israel. Geochimica et Cosmochimica Acta 64: 2039-2053. TAHAL 2000. Flows in the lower Jordan River (in Hebrew). Vengosh, A. & R. Keren 1996. Chemical modifications of groundwater contaminated by recharge of sewage effluent. Journal Contaminant Hydrology 23: 347-360. Vengosh, A. & I. Pankaratov 1998. Chloridehromide and chloridelfluoride ratios of domestic sewage effluents and

5 CONCLUSION The Jordan River exhibits large variations in chemical and isotopic compositions along 100-km flow between the Sea of Galilee and the Dead Sea. These variations reflect continued rapid exchange with subsurface flows, in addition to surface inflows to the river. Discharge measurements also reveal a net addition of water along the upper section of the river (Shavit et al. 2001). The chemical data suggest that groundwater in the northern part is derived from human activities in the vicinity of the river, as reflected also in the composition of the Yarmouk River. In the southern part, saline ground waters that are derived from natural leaching of salts in the Lisan Formation control the salinity of the river. The impact of the groundwater component on the quality of the Jordan River adds additional constraint for future management of the river. Significant 70

associated contaminated groundwater. Ground Water 36: 8 15-824. Vengosh, A. & E. Rosenthal 1994. Saline groundwater in Israel: its bearing on the water crisis in the country. Journal ofHydrologv 156: 389-430. Vengosh, A., Starinsky, A., Kolodny, Y. & A.R. Chivas 1991. Boron-isotope geochemistry as a tracer for the evolution of brines and associated hot springs from the Dead Sea, Israel. Geochimica et Cosmochimica Acta 55: 1689-1695. Vengosh, A., Spivack, A.J., Artzi, Y . & A. Ayalon 1999. Boron, strontium and oxygen isotopic and geochemical constraints for the origin of the salinity in ground water from the Mediterranean Coast of Israel. Water Resource Research 35: 1877-1894. Yechieli, Y., Magaritz, M., Levy, Y., Weber, U., Kafri, U., Woellfel, W. & G. Bonani 1993. Late Quaternary geological history of the Dead Sea area, Israel. Quaternary Research 39: 59-67.

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Geochemical cycles, global change and natural hazards

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Wafer-Rock Interaction 2001, Cidu (ed.), 02001 SWefS & Zeitlinger, Lisse, ISBN 90 2651 824 2

The chemistry of rainwater in the Mt. Etna area (Italy): sources of major species A.Aiuppa Dipartimento CFTA, Universitd di Palermo, via Archiraj? 36, 90123 Palermo, Italy

P.Bonfanti Dipartimento di Scieizze della Terra, Universita di Catania, corso Italia 55, 95129 Catania, Italy

W.D’Alessandro Istituto Nuzionule di Geofisica e Vulcanologia,Sezione di Palermo, via La Malfa 153, 90146 Palerrno, Italy

ABSTRACT: Major ion content of 37 rainwater samples collected at Nicolosi was investigated. Measured pH values range from 3.80 to 7.22 and display a positive correlation with Ca2’ and an inverse correlation with NO;, suggesting that anthropogenic HNO, is the prevailing acidifying agent while Ca, likely as solid CaCO,, is the prevailing proton acceptor. Na/Cl ratios indicate a dominant marine origin for both species. m a y Mg/Na and Ca/Na ratios, generally exceeding seawater marker ratios, point to additional sources for K, Mg and Ca (soil dust, fertilisers etc.). Nitrate and sulphate concentrations display a nearly constant ratio indicating a common anthropogenic origin. Only a few samples are characterised by sulphate excess. The analysis of time series reveals a good correlation between excess sulphate in rainwater and SO, fluxes from the summit craters plume. Chloride contents also show a significant correlation with volcanic activity. (Allard et al. 1991, Pennisi & Le Cloarec 1998, Bruno et al. 1999). Yearly calculated CO, and SO, fluxes from the Etna plume (13 and 1.2 Mt/a respectively) correspond to about 10% of global volcanic emissions. It is consequently evident that the local atmosphere is strongly polluted by this big “natural emitter”. The aim of this work is to provide a baseline for the major ion composition of rainwater in the Etnean area. This topic is of interest for at lest two reasons: i) to assess the impact of plume emissions on the local atmospheric environment and on the chemistry of meteoric-derived groundwaters hosted in the permeable volcanites at Etna (Aiuppa et al. 2000); ii) to furnish data useful in the geochemical modelling of aqueous processes. In fact, a rigorous understanding of the chemical evolution of groundwaters flowing in a given area is possible only if the chemistry of rainwater, intended as the starting point of the hydrological cycle, is known.

1 INTRODUCTION The cycling of elements in the shallow geochemical spheres (atmosphere, ocean and crust) is strictly linked to the water cycle. Through rainwater, chemical substances in the atmosphere are transported to the earth surface. As a result of the dissolution of these reactive chemical species, rainwater becomes aggressive with respect to rocks, giving rise to intense rock leaching and finally metal transport to the oceans through rivers. The origin of chemical constituents present in the atmosphere and then dissolved into forming rain droplets is quite complex, as several natural and anthropogenic sources release elements to the atmosphere. The increasing importance of human activity in releasing chemical substance to air has required much effort in the comprehension of the mechanisms governing rainwater chemistry. In particular, a better understanding of natural chemical fluxes to the atmosphere is needed to better assess the environmental impact of the anthropogenic source. Seawater spray, soil dust resuspension and organic matter decay are among the most invoked natural sources. Volcanic activity may also modify the natural chemical fluxes in the atmosphere. Mt. Etna, the biggest volcano in Europe, is in a persistent activity state since the last 200,000 years, producing a continuous degassing from central craters. Recent studies showed that volatile species as s, C1, F and trace metals, together with major compounds (H,O and CO,) are continuously released by the Etnean plume

2 METHODS 2.1 Sampling and analytical techniques Rainwater was collected by means of an automatic, wet only, sampler. 37 samples were obtained during a one year survey (from December 1990 to December 1991), the sampling frequency being 1 week. During the investigated period, 17 non-rainy weeks occurred, mainly in the summer season (June-

83

corded in the months from June to August. Total precipitation was 902 mm over the period of 54 weeks.

August). The amount of water was determined gmvimetrically, whereas pH was measured in the field with a portable instrument. Analyses of the major chemical composition were performed by atomic spectrophotometry (NayK, Mg and Ca) and HPLC (Cl, NO3 and SO4).

2.4 Volcanic activig During the sampling period, Mt. Etna displayed a variable activity status. In December 1990 continuous strombolian activity characterised the summit craters, while the period January-August 1991 was of relative quiescence. In September, strombolian activity at the summit craters resumed and increased in intensity until December, 14 when an eruption, that lasted 473 days, took place (Calvariet al. 1994). The first hours of this eruptive phase were characterised by intense explosive activity which produced a dense ash cloud and whose fine products were dispersed on the slope of the volcano in a southwestern direction (Calvariet al. 1994). During the same period, COSPEC measurements showed very high SO2 plume fluxes in December 1990, very low values in the period JanuaryNovember 1991 (except for a sharp positive peak on March), and again high values in December 1991 (Bruno et al. 1999).

2.2 Location and geological background The sampling station used throughout the whole survey was locatedat the outskirts of the Nicolosi village, at an altitude of about 670 m a.s.l., on the southern flank of Mt. Etna volcano. About 10 km SSE is the major urban agglomerate of Catania (-700,000 inhabitants) with many industrial plants while the Ionian sea is 12 km to the east and Mt. Etna's summit craters are 15 km to the north. Around Nicolosi village several basalt quarries are present and land is used mainly for vineyards. Fig. 1 displays a schematic lithologicmap of the area.

3 RESULT AND DISCUSSION Table 1. Chemical composition of the sampled rainwaters (mgil) and enrichment factors CEF) relative to seawater. pH Na K Ca Mg C1 NO3 SO4 min 3.80 0.10 600 mW/m2), low density bodies often considered volcanic centres (e.g., Torre Alfina, correlated with positive heat flow anomalies, Vesuvius, etc.), these features are unlikely. In presence of a 30-40 km wide low-velocity and addition, the required proportions of crust (3He/4He conductive body (Gianelli et al. 1997). Most of these - 0.03 Ra) to be added to MORB mantle (-8 Ra) for geophysical data support the presence of a partially lowering the 3He/4He of melt to 3 + 4 Ra, would molten granite below Larderello (Ganelli et al. largely modify the composition of melt and possibly 1997). From a geodynamic viewpoint, the observed freeze it. Therefore, a mantle enrichment of U and Th anomalies can be reconciled with a model of mantle by metasomatic fluid addition is certainly the most uplifting at the front of the subducting Adriatic plate, plausible mechanism to explain the low 3He/4He that causes melting of the lithosphere with addition observed values. This conclusion is hlly in of new asthenospheric material, isotherms uprising agreement with the above summarised characteristics and granitization of the southern Tuscany crust

104

(Gianelli et al. 1997 and references therein). In this setting, fluids coming from mantle can easily reach the surface, facilitated by the extensional tectonics of the area. Helium isotope ratio in gas emissions of the Tuscany region rapidly decreases towards the East (Fig. l), reaching 0.09 Ra (Rapolano Terme) only 60 km afar from the Larderello area and the pure crustal end-member 0.02 Ra in the Umbria region. This sudden decrease of mantle gas contribution puts a strong constrain on the localisation of the mantle uplift and on its extension beneath the Apennine crust. No mantle uplift results beneath the zone where helium isotope ratio reaches typical crustal values. The gas output in the main foredeeps of the Apenninic orogene (POBasin and Bradanic Through) is scarce and (2%-dominated. In the PO Basin, values range between 0.005 and 0.04 Ra (0.005-0.03 Ra in gas dissolved in spring waters, Marty at al. 1992; 0.02 - 0.04 Ra in free gas from drilled wells, Elliot et al. 1993), clearly showing that He in the POBasin has a typical crustal origin without any appreciable mantle contribution. The lack of mantle helium in the PO Basin is probably due to the genesis of this basin, which was formed by crustal loading. On the contrary, basins formed by crustal extension (e.g. Pannonian Basin, Rhinegraben) display a mantle helium signature because active tectonics facilitates the route to the surface for deep fluids (O’Nions & Oxburgh 1988). 2.3 General distributioti

The main feature observable from the 3He/4He distribution in peninsular Italy is the presence of a higher mantle contribution in the perithyrrenian margin with the respect to the inner or periadriatic regions. In addition, as already noted, in the perityrrhenian margin of the Italian peninsula a series of recent or active volcanoes are present (Fig. l), clearly testifjling the relatively easy ascent to the surface of mantle melts in this area. Gas output, mainly C02-dominated, is abundant in the perityrrhenian and Apenninic sectors, while it considerably decreases towards the East where consequently few data are available (Chiodini et al., 2000). The helium isotope distribution in the Italian peninsula is coherent with the pattern of some geophysical parameters like the heat flow and the seismic wave attenuation. Heat flow values progressively decrease from the perityrrhenian margin (1 00- 150 mW/m2) towards the Adriatic sea (< 40 mW/m2, Cataldi et al. 1995). In addition, recent studies carried out by Mele et al. (1997) on the propagation of Pn and Sn seismic waves, have shown a high-attenuation zone in the upper mantle beneath the southern Tyrrhenian sea, the Apennine chain and the western Italy.

Figure 1. 3He/4Hevalues (expressed as inultiples of Ra) for the peninsular Italy. Data after: Hooker et al. 1985; Sano et al. 1989; Marty et al. 1992; Graham et al. 1993; Eliot et al. 1993; Vaselli et al. 1997; Minissale et al. 1997; Italian0 et al. 2000. 0.15 and 0.26 Ra values in Umbria region are unpublished data (Montecastello di Vibio and S. Faustino gas emissions, respectively). A more detailed distribution of 3He/4Hedata in Latiuin and Tuscany regions is given by Minissale et al. 1997, where is even clearer the decrease of 3He/4Hevalues from the perithyrrenian side of Italy towards the East. Geologc elements after Cataldi et al. 1995, modified. In white the Apenninic units. 1: Apenninic foredeep; 2: Adriatic-Apulian foreland; 3: main Apenninic thrust front; 4; recent potassic volcanics. L = Larderello, V = Vesuvius, I = Ischia, PF = Phlegrean Fields, VL = Vulture

No attenuation is recorded on the periadriatic margin of Italy and beneath the PO Plain. In the Authors’ opinion, the observed shear wave attenuation suggests that the lithospheric mantle beneath the internal units of the Apennines and western Italy is contaminated by the advection of relatively hot material at shallow depth, which drastically changes its thermal structure. These geochemical and geophysical distributions are interpreted in the key of the geodynamic model that describes a mantle uprising, with consequent crustal thinning (Fig. 2), in the Tyrrhenian and perityrrhenian area, as a consequence of the westward subduction of the Adriatic plate beneath the Apennines (e.g. Doglioni et al. 1996). Mantle and/or mantle fluids uprise in the ferityrrhenian area may reconcile the distributions of He/ He values, of heat flow and of seismic wave attenuation in the Italian peninsula.

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3 CONCLUSIONS

He isotope ratio in the volcanic areas of peninsular Italy displays values I 4.6 Ra, both in free gas and fluids included in recent lavas. A mantle metasomatised by crustal-derived fluids is the most plausible hypothesis to explain these low values with respect to a MORB-type mantle. This interpretation fits with petrological and geochemical data of volcanic rocks of the region, even though fbrther studies are needed to better constrain the timing of metasomatism and the eventual different episodes. Tectonically active areas of the Italian peninsula (Tyrrhenian coast and Apenninic chain) show total gas output and mantle gas contribution much higher than tectonically stable areas (foredeep and foreland). The general 3He/4He distribution in the peninsular Italy clearly shows that mantle fluid contribution is greatest in the perityrrhenian regions and decreases towards the periadriatic regions. This trend is coherent with other distributions, such as the recent and active volcanism, the crustal thickness, the heat flow density and the seismic wave attenuation. All these distributions can be framed in the context of the Adriatic subduction beneath the Apennines.

REFERENCES Allard, P., Baptiste, J.B., D’Alessandro, W., Parello, F., Parisi, C.B. & C. Flehoc 1997. Mantle-derived helium and carbon in groundwaters of Mount Etna, Italy. Earth Planet. Sci. Lett. 148: 501-516. Cataldi, R., Monelli, F., Squarci, P., Tafl?, L., Zito, G. & C. Calore 1995. Geothermal ranking of Italian territory. Geotherniics 24: 115-129. Chiodini, G., Frondini, F., Cardellini, C., Parello, F. & L. Peruzzi 2000. Rate of diffuse carbon dioxide Earth degassing estimated from carbon balance of regional aquifers: the case of central Apennine, Italy. J. Geophys. Res. 105: 8423-8434. De Fino, M., La Volpe, L., Peccerillo, A., Piccarrela, G. & G. Poli 1986. Petrogenesis of Monte Vulture volcano (Italy): inferences from mineral chemistry, major and trace element data. Contrib. Mineral. Petrol. 92: 135-145. Doglioni, C., Harabaglia, P., Martinelli, G., Monelli, F. &Zito, G. Zito 1996. A geodynamic model of the Southern Appennines accretionary prism. Terra Nova 8: 540-547. Elliot, T., Ballentine, C.J., O’Nions, R.K. & T. Ricchiuto 1993. Carbon, helium, neon and argon isotopes in a PO Basin (northern Italy) natural gas field. Chem. Geol. 106: 429-440. Gianelli, G., Manzella, A., & M. Puxeddu 1997. Crustal models of the geotherinal areas of southern Tuscany (Italy). Tectonophysics 281: 221-239. Graham, D.W., Allard, P., Kilburn, C.R.J., Spera, E.J. & J.E. Lupton 1993. Helium isotopes in some historical lavas

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from Mount Vesuvius. J. Volcanol. Geotherm. Res. 58: 359-366. Hooker, P.J., Bertrami, R., Lombardi, S., O’Nions, R.K. & E.R. Oxburg 1985. Helium-3 anomalies and crust mantle interactions in Italy. Geochim. Cosmochim. Acta 49: 2505-25 13. Inguaggiato, S., & G. Pecoraino, 1998. Chemical and isotopic features of gas manifestations at PNegrean Fields and Ischia Island, Italy. In G.B. Arehart & J.R. Hulston (eds), Water-rock interaction, 633-636, Rotterdam: Balkema. Italiano, F., Martelli, M., Martinelli, G. & P.M. Nuccio 2000. Geocheinical evidence of melt intrusions along lithospheric faults of the Southern Apennines (Italy): geodynamic and seismogenic implications. J. Geophys. Res. 106: 13569-13578. Marty, B., O’Nions, R.K., Oxburgh, E.R., Martel, D. & S. Lombardi 1992. Helium isotopes in Alpine regons. Tectonophysics 206: 71-78. Marty, B., Trull, T., Luzziez, P., Basile, I. & J.C. Tanguy 1994. He, Ar, 0, Sr and Nd isotope constraints on the origin and evolution of Mount Etna magmatism. Earth Planet. Sci. Lett. 126: 23-39. Mele, G., Rovelli, A., Seber, D. & M. Barazangi 1997. Shear wave attenuation in the lithosphere beneath Italy and surrounding regions: Tectonic implications. J. Geophys. Res. 102: 11863-11875. Minissale, A., Evans, W.C., Magro, G. & 0. Vaselli 1997. Multiple source components in gas manifestation from north-central Italy. Chem. Geol. 142: 175-192. O’Nions, R.K. & E.R. Oxburgh 1988. Helium, volatile fluxes and the development of continental crust. Earth Planet. Sci. Lett. 90: 331-347. Peccerillo, A., 1999. Multiple mantle metasomatism in central-southem Italy: geochemical effects, timing and geodynarnic implications. Geology 27: 3 15-318. Sano, Y., Wakita, H., Italiano, F. & P.M. Nuccio 1989. Helium isotopes and tectonics in Southern Italy. Geophys. Res. Lett. 16: 51 1-514. Tedesco, D., Allard, P., Sano, Y., Waluta, H. & R. Pece 1990. Helium-3 in subaerial and submarine fumaroles of Campi Flegrei caldera, Italy. Geochim. Cosmochini. Acta 54: 1105-1116. Vaselli, O., Tassi, F., Minissale, A., Capaccioni, B., Magro, G. & W.C. Evans 1997. Geochemistry of natural gas manifestations from the upper Tiber Valley (Central Italy). Miner. Petrogr. Acta XL: 201-212. Villeinant, B. & C. Flehoc 1989. U-Th fractionation by fluids in K-rich magma genesis: the Vico volcano, Central Italy. Earth Planet. Sci. Lett. 91: 312-326.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Water-rock interactions during seismic an volcanic activity recorded at Mount Etna by continuous roundwater monitoring F.Quattrocchi , G .Di Stefano , G.Galli, L.Pizzino & P.Scarlato Istituto Nazionale di Geofsica, Via di Vigna Murata 605, 00143, Roma, Italy P.Allard & D .Andronico Sisrema Poseidon, Via Monti Rossi 12, 95030, Nicolosi (CT)

D .Condarelli & T.Sgroi Istituto lnternazionale di Vulcanologia,Piazzale Roma 2, 95123 Catania

ABSTRACT: In this paper we present and discuss geochemical monitoring data for the A q u a Dfesa water well on the Southern slopes of Etna volcano, which were gathered by the Geochemical Monitoring System N (GMS 11) prototype. This automatic monitoring station provides, every ten minutes, data recording of groundwater temperature, electrical conductivity, pH, redox potential, dissolved C02, atmospheric pressure and air temperature. An automatic sampler collects, every 3 days, 2 bottles of water, acidified and not, for subsequent laboratory analyses (including B, NH4, Fetotand dissolved C02). During the period investigated (March to September 2000) recurrent paroxysmal eruptions occurred at the South-East crater (SEC) of Etna. The A q u a Difesa GMS I1 station recorded a few geochemical anomalies whose relationships with either volcanic or/and seismic activity are evaluated. We find a possible correlation between the greatest geochemical anomalies, in April 2000, and the most powerful lava fountains episodes at the SEC. Although preliminary, the results open new perspectives to understand better the response of etnean aquifers to volcanic crises. Further continuous monitoring at this and other remote sites on Etna’s flank should help us to characterise the dynamics of water-gas-rock interactions associated with volcano-tectonic processes. 1 INTRODUCTION The continuous monitoring of active volcanoes and the forecasting of their eruptions have been among the priorities of international programmes of Earth Sciences and hazard reduction over the last decades (Automatic Geochemical Monitoring of Volcanoes EC Program 1996-98, Contract N. ENV4-CT960289). Geochemical monitoring of volcanoes has produced a huge background of data, that has allowed us to elaborate and test various theoretical models of magma degassing processes. However, as for geophysical parameters, it is still necessary to conduct continuous monitoring of geochemical parameters at volcanoes, in order to retrieve better the precursory signals of eruptions. Several efforts have been made in this direction (e.g., Toutain et al. 1992, Quattrocchi et al. 2000 and references herein), thanks to the availability of lower cost components for instrumental devices. In Italy, the needs of civil defense and land management have promoted such efforts with the support of the Italian Civil Protection. The requirement for multiparametric continuous geochemical monitoring of volcanoes is obvious, but usually excluds hot gas emissions that either are poorly accessible or too corrosive for the instrumentation. Most recent efforts were thus

directed towards continuous monitoring of either volcanic groundwater, which can integrate and reveal changes in magmatic gas inputs (e.g., Quattrocchi et al. 2000), or/and soil gas emanations Here we report and discuss data that were obtained through continuous monitoring of a water well on the southern flank of Etna volcano, using a multiparametric prototype: the Geochemical Monitoring System I1 (GMS 11). 2 METHODS 2.1 Site selection criteria and geochemical frame. The A q u a Difesa water well (200 m deep, flow: 50 l/sec) is located near Belpasso city, at 600 m a.s.1. on the Southern flank of Etna (UTM location: 33 SVVB957625). Two pumps (the deepest routinely switched off) tap two different aquifer strata, the deepest groundwater being warmer, richer in gas and more mineralised than the shallower one. This configuration allows us to test the specific response of the sensors and, once established, its influence can be discriminated from true volcanic effects (Quattrocchi et al. 2000). The choice of the well was done on the basis of previous chemical and isotopic investigations addressed to study relationships between magmatic input and the groundwater 107

system (e.g., Allard et al. 1997 and references herein, Quattrocchi et al. 2000). The Southern and Eastern sectors of the volcano, and particularly the Paterno-Belpasso area, where the Acqua Difesa well is located, were discovered to be preferential zones where groundwaters are the most enriched in magma-derived volatiles and where diffuse soil degassing of magma-derived carbon-dioxide is the most extensive (pc02 up to 1.2 atm, 3He/4Hemantle signature up to 6.6 Ra). A few kilometres downslope south of the well is the Salinelle di Paternd, where mud volcanoes and waters enriched in B, Fe, Li, Si02, NH4 and heavy metals are fed by a hydrothermal system of mixed magmaticsedimentary origin (Chiodini et al. 1996). These areas are of special interest to detect any episodic increase of deep end-members into the overlying main aquifer of meteoric origin that is typically tapped at the Acqua Difesa well. This may occur, for example, as a consequence of episodic uprising of a B-NH4 rich gas phase, originating from a boiling hydrothermal aquifer (normally with around 0.3 ppm of B), as suggested by the Cl/B ratio in the Paterni, area, which is lower than the rock ratio (Chiodini et al. 1996). Variations in the stress-strain and fracture events may also modify the mixing ratios of components between normally isolated or sealed aquifer strata. Finally, an increased input of magmatic gas due to the rise and decompression of a large magma batch might also be reflected in significant variations in water chemistry, particularly in the partial pressure of CO2 and helium. Therefore, our geochemical monitoring at the Acqua Difesa well is aimed at evaluating such possible geochemical variations in groundwater in relationship to the active volcano-tectonic processes.

microprocessor and 24 bottles is currently on-line, collecting every 72 hours 2 bottles of 500 ml which are either acidified (to pH 2.0) or not. NI& was measured by a colorimetric method: at pH 12.6 the ammonium ions react with ipo-chloride and salicilatic ions in the presence of sodic nitroprussiate, as catalyst, giving indofenolic bleu (field of sensitivity: 0.015-2.0 mg/L). The Boric Acid (B in mg/L) determination was carried out by a colorimetric method based upon the reaction with Azometina-H (field of sensitivity 0.05-2.5 mg/L). The total dissolved Fe was determined by using the 1.10 phenantroline colorimetric method. Dissolved CO2 was measured by ORION ion selective electrode, with a buffered pH control.

2.2 Continuous and discrete monitoring methods Details of the GMS I1 prototype are described in Quattrocchi et al. (2000). With respect to the former GMS I (details in the previous paper) this second prototype mainly differs in that the sensors are managed separately, thus eliminating the use of closed pre-assembled probes which prevent an easy access to the signal path as a whole. The entire path of the signal is accessible, conceiving the acquisition channels the base units, on which the entire system has been built up. In this way, the GMS I1 remote station can operate with variable sensors and/or instruments on line and in complex configurations (for groundwaters, soils or fumaroles monitoring). The basic GMS I1 configuration installed at the Acqua Difesa well may detect at present: water temperature, pH, Eh, electrical conductivity, dissolved Col, Rn, He, H2S, air pressure and temperature. Details are given in Quattrocchi et al. (2000). An auto-sampler (ISCOTM)equipped with a

Figure 1. Continuous monitoring trends recorded at the A q u a Dfesa well during the period April-May 2000; text for details.

3 RESULTS 3.1 Hydrogeochemical patterns and anomalies Taking into consideration the continuous monitoring data set, from 1/3/2000 to 15/9/2000 the following main geochemical anomalies were observed (gray zones and horizontal lines in Fig. 1 referred to the selected period April-May 2000): i) while the other signals remained almost constant, a slight longperiod drop of the Eh signal (around -100 mV) was observed, spanning the second half of March 2000. 108

After the Eh values remained around 280 mV; ii) the most apparent geochemical anomaly throughout the discussed period appeared as short-term anomaly on the early morning of 18/04/2000. At 6.00 a.m. the CO2 signal started to rise first together with the electrical conductivity, these reaching 1400 ppm and around 1400 pS/cm respectively (well outside the surrounding fi f 2 (T range; for the CO2 anomaly see the line with external arrows in Fig. 1). Almost contemporaneously the Eh dropped of around 150 mV in 1 hour (minimum at 10:42 in the morning, also exceeding the fi -+ 2 (T range), while the pH dropped by around 0.2 pH units. However, the temperature remained constant. The overall explanation of this anomaly may be a sharp input of reduced and acid cold gas, that may explain the observed increased salinity (i.e., enhanced waterrock interaction in this sector of the aquifer), in a very abrupt manner; iii) the most apparent long-term geochemical change affected the Eh signal which exhibited a strong negative variation from May 1 to May 13 and remained constant during the next two months; iv) a positive spike in the CO2 signal occurred on May 17 (line with internal arrows in Fig. 1); v) two other similar CO2 positive spikes took place: one between June 22 and June 26 and the other between June 28 and July 1”; vi) a long-trend Eh positive variation started on August 28, allowing the Eh signal to recover back to the levels recorded before the early May negative change (see above).

Figure 2. A q u a Dfesa discrete chemical analyses (3-6/2000).

We have also taken in consideration the results of discrete monitoring (29/2-2 1/06/2000) made up of chemical analyses (Fetot, B, NH4, dissolved CO$ carried out on the automatically stored water samples. We observed the following trends: i) the dissolved CO2 values never moved from the mean fi f 2 (T range; ii) the Fetot,B and NH4 values (Fig. 2) remained inside the A rfr 2 (T range during the observation period as a whole with the exception of the 18/03/2000 data, when Fetotreached 5.14 mg/L (mean: 2.32 mg/L), the B reached 0.69 mg/L (mean: 0.42 mg/L) and NH4 reached 0.08 mg/L (mean: 0.04 mg/L). This anomaly may be explained by a short-

term episode of: i) enhanced water-rock interaction processes ii) a spike of reduced and B-rich gases and iii) a slight mixing between the typical Acqua Difesa groundwater and a more reduced and B-richer liquid component, i.e., near to the deep reservoir endmember (see Chiodini et al. 1996). 3.2 Volcanic and seismic activity On January 26,2000 the Southeast Crater (SEC), the youngest sub-terminal cone of Mt. Etna volcano, reactivated after about 3 months of quiescence. Until June 24, 64 lava fountain episodes occurred, often preceded (but usually followed) by lava flows (reaching the Torre del Filosofo refuge on 1214/3/2000), and interspersed with periods of total or partial rest. We defined this activity “the 2000 First Semester eruption at SEC”, consisting of the periodic ascent of fresh magma batches that filled a shallow plumbing system, and was divided it into at least two main stages. During the first one, until February 23, the fire-fountain episodes occurred more frequently than the second one, depending on repose time, duration and eruptive style of each episode. During the whole period, volcanic tremor was characterised by a progressive increase of amplitude reflecting the behaviour of explosive phases at SEC. Two months later (August 28-29) another fire fountain event occurred, but the day after violent strombolian activity at SEC didn’t reach the same intensity of fire fountaining. Figure 3 compares the timing of the 65 explosive events (plus the paroxysmal activity of August 29) with temporal progress of Reduced Displacement (R.D.) linked to volcanic tremor, calculated on surface waves through an equation provided by Fehler (1983). The tremor started to uprise at the end of March 2000. Each lava fountain episode can be divided into three phases: a) resumption phase (gradual increase in volcanic tremor amplitude, increase in explosive activity, effusion from one or two vents; b) paroxysmal phase (transition from strombolian activity to sustained lava fountains reaching an height of 800 m, ash, lapilli and steam rising up to 26 km); c) conclusive phase (drop in volcanic tremor, re-establishment of the normal seismic pattern). Among the 65 episodes, the April 16, 2000 fire fountain (the 50th from January 26) represents the most spectacular explosive episode. The paroxysmal phase lasted three and a half hours, and was characterised by a variable eruptive pattern with an oscillating eruption column ranging between 2 and 6 km (the highest observed during the First Semester 2000 eruption) above the cone. We believe that the previous considerations make this fountain the most energetic episode of the studied time period (March 1-September 15, 2000). We also deduce the importance of this episode from the pattern of volcanic tremor: in fact amplitudes are the highest

109

ones and duration of the paroxysmal phase the longest one of the studied period. It demonstrates that tremor amplitude reflects the pattern of volcanic explosive phases. It’s possible that the relatively long rest before this event (about 10-20 days), compared with the other episodes, permitted a greater accumulation of gas and new hot magma replenished the shallow eruptive system.

Figure 3. Reduced Displacement (cm2) linked to the volcanic tremor calculated at EMF station (1/1- 15/9/2000).

As regards the seismic activity as a whole, the most relevant seismic events which took place in the SSW slope (where is concentrated the strain release of the volcanic structure) during March-September 2000: i) a seismic swarm of maximum intensity (111) on 26/03/2000 (El in Fig. 1); ii) a highest magnitude earthquake (Md 3.O), recorded on 15/04/2000, within 10 km of the well (lat. 37.75N, long. 14.92E, ING data bank, E2 in Fig. 1); iii) a seismic swarm (E3 in Fig. 1) occurred on 06/05/2000 at the same time of a paroxysmal phase of volcanic activity; iv) another relative strong event located near Nicolosi on 27/05/2000 at 13:19 local time (E4 in Fig. 1).

first half of March 2000. This was just a few days before the on-set of a period of enhanced calculated Reduced Displacements and before the most energetic seismic swarm occurred along the Southern slope of the volcano. We hypothesise that volcano-faulting, volcanic stress-field changes and hydrofracturing that trigger swarm seismicity create possibly gaseousheat input along pathways to the surface through aquifers; b) the most apparent geochemical anomalies recorded by continuous monitoring (namely those of the 18/04/2000 and the 1-13/05/2000) occurring during the period of the highest calculated Reduced Displacement of each tremor-lava fountain episode. In particular, it was observed in coincidence with the most energetic event on 16/04/2000. The two main geochemical anomalies in the six months (1/3 to 15/9/2000) may be both explained by short-term episodes of reduced and acid gases input which modify the overall physico-chemical patterns of the main aquifer and the water-rock interaction processes. They occurred just when a huge quantity of gas was emitted at SEC, together with lava fountains and high tremor. One may thus hypothesise that in April and May 2000, pressurised volatiles interacted with shallow fluids reservoirs, creating geochemical anomalies at the surface. Further monitoring using also the wider Poseidon System, geochemical network should allow us to improve our understanding of these observations. REFERENCES Allard, P., Jean-Baptiste, P., D’Alessandro, W., Parello, F., Parisi, B. & C. Flehoc 1997. Mantle-derived helium and carbon in groundwaters and gases of Mount Etna, Italy. Earth Plan. Sci.Lett. 148: 501-516. Chiodini, G., D’Alessandro, W. & F. Parello 1996. Geochemistry of the gases and the waters discharged by the mud volcanoes of Paterno, Mt. Etna (Italy). Bull. Volcanol. 58: 51-58. Dall’Aglio, M., Pagani, F. & F. Quattrocchi 1994. Geochemistry of the groundwater in the Etna region before and after the paroxysmal phase of the eruption of December 1991: implications for geochemical surveillance of Mt. Etna. Acta Vulcanol. 4: 149-156. Fehler, M. 1983. Observations of volcanic tremor at Mount St. Helens volcano. J. Geophys. Res. 88: 3476-3484. Quattrocchi, F., Di Stefano, G., Pizzino, L., Pongetti, F., Romeo, G., Scarlato, P., Sciacca, U. & G. Urbini 2000. Geochemical Monitoring System 11 prototype (GMS 11) installation at the “Acqua Difesa” well, within the Etna region: first data during the 1999 volcanic crisis. J. Volc. Geoth. Res. 101: 273-306. Toutain, J.P., Baubron, J.C., Le Bronec, J., Allard, P., Briole, P., Marty, B., Miele, G., Tedesco, D. & G. Luongo 1992. Continuous monitoring of vocanic gases in a water-well at Vulcano, Southern Italy. Bull. Volcanol. 54: 147-155.

4 DISCUSSION AND CONCLUSIONS It is very difficult to firmly establish a causal correlation between geochemical and geophysical data sets recorded throughout an active volcanic area, mostly if data are gathered only by using a single geochemical station. At this preliminary stage we may simply compare the two different data sets in order to assess the possible “response” of the Acqua Difesa test-site to the ongoing volcanotectonic processes, as done also in the past Etna volcanic activity. The data reported here, during “the 2000 First Semester Eruption at SEC” suggest a possible correlation between some of the anomalies at the A q u a Difesa well and dynamic changes in the aquifer linked with the volcanic andor seismic crises, during the same period. Schematically we may have observed: a) an episode of a relatively apparent input of deep component pathfinder elements (B, NH4, Fe) related to either an uprising gaseous phase or to a deep groundwater during the 110

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2

The Ardea Basin fluid geochemistry,hydrogeology and structural patterns: new insights about the geothermal unrest activity of the Alban Hills quiescent volcano (Rome, Italy) and its geochemical hazard surveillance F.Quattrocchi, G.Galli & L.Pizzino Istituto Nazionale di Geofisica e Vulcanologia, Via di Vigna Murata 605, 00143, Roma, Italy

G Capelli, D .De Rita, C .Faccenna, R.Funiciello, G .Giordano, D. Goletto & R.Mazza University “Roma Tre”, Dip. Sci. Terra, Largo S L . Murialdo, 00146, Roma, Italy

C .Mancini University “La Sapienza”, Dip. Ing. Nucl., Piazza S. Pietro in Vincoli 10, 00183, Roma, Italy

ABSTRACT: A multidisciplinary task study was undertaken within the Ardea Basin, located at the SouthEastern boundary of the Alban Hills quiescent volcano (central Italy), encompassing fluid geochemistry, hydrogeology, structural geology and seismotectonics. The work is aimed at evaluating: i) the active geodynamics and geothermal unrest activity of the volcano; ii) the Natural Gases Hazard (C02, H2S, CH4 and other dangerous peri-volcanic gases); iii) the Rn-indoor hazard prone sectors; iv) the effects of anthropogenic depletion of shallow aquifers on the Natural Gases Hazard. The Ardea Basin, bordered by a complex fault system, and interpreted as a transfer related basin, is marked by a high CO2 flux and by hydrothermal and gas-emitting sites, which were revisited during this work allowing new geothermometric and chemical equilibria evaluations and giving renewed insights into the geothermal reservoir. The pC02 and 222Rnin groundwater were mapped to finally assess a moderate to low NGH, excluding the Fossignano and Zolforata pools. The main positive anomalies are located along the border of the basin, corresponding to NE transfer fault systems.

1 INTRODUCTION

addition, our unpublished geochemical data (major, minor, trace elements and He-C isotopic ratios) were reworked. Suitable and tested methods have been adopted during surveying (Capelli et al. 1999, Pizzino et al. 2000, Mancini et al. 2000). The need to accomplish a geochemical-hydrogeological micro-zonation, aimed at natural hazards assessment, arose from the detailed study of other sectors of the Alban Hills volcano, where episodes of strong peri-volcanic degassing occurred (i.e., November, 1995 and September 1999, over the Ciampino-Marino area, see Fig. 1, details in Pizzino et al. 2000). This study was indeed recommended by the Civil Protection Department, and also for future continuous monitoring. A recently funded GNV project (Gruppo Nazionale di Vulcanologia) is dealing with the diffuse gaseous emissions throughout Italian volcanoes and fault systems; the main task being to locate in detail the most dangerous NGH prone-areas and to calculate the diffuse CO2 fluxes as a whole in each area (Chiodini & Frondini 2000). In order to evaluate fully the NGH assessment, we have to know in detail the ongoing geodynamic processes, the hydrogeological patterns, the fluidrock interaction processes and the possible convection regime condition which is set up.

The Alban Hills quiescent volcano is affected by unrest activity, that consists mainly of seismic swarms, deformation of the volcanic structure and spread degassing (Delaney et al. 1996, Kerner et al. 2000, Chiodini & Frondini 2000). Its Natural Gases Hazard (NGH) is defined as “hazard by diffuse gaseous exhalations of natural origin and Rn-indoor hazard” (see also Pizzino et al. 2000). The study focused mainly on the nature and evolution of endogenous fluids (Giggenbach et al. 1988, Quattrocchi & Venanzi 1989, Duchi et al. 1991, Pizzino et al. 2000) circulating inside the shallower reservoirs of the volcano, in which mixing occurs between meteoric circulation and deep fluids, mainly gases of volcanic and thermo-metamorphic origin (peri-volcanic type). The depletion of the main aquifer by anthropogenic activity may enhance the NGH, as a consequence of gadwater ratio rising and gas solubility decreasing as a whole with time (Pizzino et al. 2000, Capelli et al. 1999). This work discusses a new detailed field survey, solely inside the Ardea Basin (I.G.M. sheet N. 158, IV square), which characterises groundwater (around 60 sites), by measuring hydrogeological parameters, physic-chemistry, CO2 and 222Rn. In 111

2 SEISMOTECTONICS OF THE ARDEA BASIN Along the NW-trending Roman Comagmatic Province (RCP, see De Rita et al. 1988, Faccenna et al. 1994, Kerner et al. 2000 and references herein), in which the Alban Hills volcano is located, the carbonate basement underlying the volcanic structures has different seismogenic behavior, probably due to a different pore pressure regime. Within this framework, the Alban Hills volcano is considered the most hazardous seismic structure inside the RCP (Amato et al. 1994), probably as a consequence of a huge fluid circulation, that in the past involved phreatomagmatic eruptions, and today implies a fading offgeothermal system (Giggenbach 1988). The role of fluids in earthquake triggering was recently recognised as very important especially in quiescent'active volcanic structures, geothermal areas and extension structures, as the Alban Hills volcano (Fournier 1987, Delaney et al. 1996). Actually an open debate is ongoing about the recent activity of the volcano and its quiescence (De Rita et al. 1988). Some authors stress the unrest activity of the volcano (Delaney et al. 1996) and recently some others have reviewed critically the stratigraphic, petrologic and geo-chronologic data to finally hypothesize an ongoing renewed volcanic cycle (Kerner et al. 2000). The Tyrrhenian margin is characterized by a NESW oriented stretching regime with main NW-SE faults segmented by NE-SW transfer fault zones that border narrow and asymmetric basins. Accordingly, the Ardea Basin (Fig. 1) was defined as a transfer related basin, (Faccenna et al. 1994) developed during the end of the Pliocene to early Pleistocene. It is a morphological depression, as confirmed by the stratigraphical and hydrogeological reworking of the present study. It is localized between Pomezia and Aprilia towns: 30 Km long and 10 Km wide elongated in a N40°E direction. The basin is bounded by a NW-dipping master fault in an overall half graben asymmetric structure. Antithetic NE-S W orientated faults are developed within a small rollover anticline structure. The North-western border is marked by a N40E normal fault with a roll-over structure and a SE vertical throw. The fault is covered by Galerian sub-horizontal un-deformed deposits. The Ardea Basin is one of the deepest basins of the Tyrrhenian margin: a lot of Pleistocene terrigenous and volcanic deposits (at least three pyroclastic flows, reaching around 100 m thickness) accumulated and reach anomalous thickness (up to 1600 m, see Faccenna et al. 1994, and references herein) with respect to the adjacent areas. As it occurs along other transfer faults, either along the North-Western normal fault or along the ArdeaFossignano master fault, the intersection of differently-oriented structures strongly affects the geochemical patterns of fluids: the former is affected

by the Zolforata gaseous pools (ZFR in Fig. l), while the latter is affected by the Fossignano gaseous pools, namely the FSG and FSGD sites in Figure 1. At the last site, an eruption in 1995 produced a fluid with 27 O C , pH = 3.0 and with a CHq content the highest found throughout the Alban Hills volcano, inferring a mixing between biogenic and thermogenic components. In the vicinity are located the thermal Ardea groundwaters (ARD and ARDH in Fig. 1).

Figure 1. Geological patterns of the investigated area.

3 RESULTS 3.1 Hydrogeological patterns and aquifer depletion The hydrogeological patterns of the Alban Hills structure substantially reflect the complexity of the overlapping and alternation between lava bodies and pyroclastic products (De Rita et al. 1988, Capelli et al. 1999). Radial patterns towards the coast throughout the Ardea Basin sector have been confirmed and detailed in this study. For this sector, a new geo-referenced data bank, exploiting GIS tools was carried out, using data from around 700 bore-holes (original references, stratigraphy, hydrogeology, chemistry, facilities, etc. . .). Two profiles were drawn and interpreted (Fig. 2) allowing us to define in details the Ardea Basin structure (e.g., boreholes 736 and 56) and the bordering faults (e.g., boreholes 298 and 505). The sandy-clay strata (Pliocene-Calabrian age) 112

underlying the volcanic deposits were found gently dipping north-eastward, and also uplifted southwestward, to finally outcrop near Pratica di MareTor Caldara-Lavinio along the coast. The contact between the volcanic deposits and the NeogeneQuaternary clayey deposits was developed following an irregular surface located at a depth around - 100 m (S. Procula, where the maximum thickness of the main aquifer was found at 120 m) and + 40 m a.s.1. (Padiglione sector). The top of the clayey stratum is variably undulating, with a wide depressed area (Ardeatina Depression) characterised by a trapezoidal shape NE-S W oriented, with the biggest base towards the Alban Hills summit. Inside this depression the hydraulic head reach maximum potential. The top of the saturated zone (free phreatic surface) was reconstructed by using 4 data sets: 1956, 1970, 1983, 1999 (our survey involved 56 sites). It allowed us to discriminate the prone-sectors affected by anthropogenic depletion during the last 40 years (around 7000 wells were ,drilled in the area), namely the Bosco di Padiglione sector (maximum depletion at 50 m), the Pratica di MareTorre S. Anastasio area, the Campoleone sector and the Ardea town. 3.2 Hydrogeochemical patterns and anomalies The main aquifer of the Alban Hills volcano, confined mainly within the volcanic rocks, is essentially cold; nevertheless it locally receives a peri-volcanic gaseous input from the underlying basement making it slightly hot (i.e., sectors such as Ciampino-Marino, Capannelle, Valleranello Morena, Frattocchie, etc.. .). The geochemical anomalies observed in groundwater during the last 20 years are hypothesized to be the result of episodes of enhanced phase separation processes at depth between a hyper-saline brine and a vapor (Fournier 1987). This processes seems to be triggered and correlated with seismicity linked to extensive and transtensive tectonics. Only within the Ardea Basin did we encounter drillings (i.e., ARDH) reaching around 55 'C. At this site complete chemical and isotopic analyses were carried out together with reworking chemical equilibria and geothermometric algorithms (Solmneq code). We found different results with respect to the previous papers (Giggenbach et al. 1988, Duchi et al. 1991): a medium enthalpy reservoir with 220 'C. Very high saturation indexes were calculated, mainly for the sulphide species, confirmed by an apparent pyrite precipitation at the bottom of the well (crystals growth on the sandstone at 80 m). The isotopic data, gathered by the Geochemical Seismic Zonation EC Program (Contract. ENV4-CT96-029 1) are

following. ARDH: i?i1*0= -5.84%0 613C=+7.82%o, R/Ra(3He/4He)=0.89,4He/2@Ne= 2.95, He= 4. 14.10-l2 mol/L; ARD: 613C=+6.88 %o; FSGD: 613C=+12.1 %o. They suggest a deep in ut and a slight mantle signature. Moreover, the 422Rn content i.e., the highest discovered within the Alban Hill (up to 1185 Bq/L), infers the existence of a fast hydrothermal circulation and a widespread fiacture system. Table 1. Gas-phase analyses of the Ardea Basin pools Site CO2 H2S CH4 02+Ar N2 Ne 222Rn _ _ _ _ _ ~ ~ ~ % ppm ppm YO % ppm BqIL

FSG 29.3 120 362 FSGD 74.0 16000 3200 ZFR 71.8 11000 190

13.2 5.2 10.2

-

57.0 13.0 98 20.3 3.0 222 28.9 10.0 110

Apart from the FSGD and ZFR highly hazardous pools, the NGH assessment was evaluated moderately low by mapping the pC02 and 222Rnin the groundwater. The NGH prone sectors are: the Ardea and Bracona villages (pC02>0.5 bar) and the Acqua Buona sector (pC02>1 bar). 222Rn values greater than 80 Bq/L were found in the Bracona and Ardea villages and South from S. Procula. These pC02 and 222Rnanomalies are elongated in a NESW direction, following the Ardea Basin Eastern master fault. As foreseen, the Radon anomaly distribution is wider, due to its complex geochemical mobility (see Pizzino et al. 2000). The Factor Analysis, carried out for all the samples, strictly associates the pC02 and Rn variables, confirming also in this sector the role of CO2 as carrier to transport rare gases of shallower origin. They are pathfinder of convection conditions settings: the discovered thickness of the volcanities -seat of Urich minerals- alone is not enough to justify the high 222 Rn content. The groundwater salinity is enhanced by two main processes: i) a slight mixing with seawater, limited within a thin sector along the coast; ii) an enhanced leaching by the input of acidic and reduced gases. This last process is seem in at the S. Stefano thermal waters and at the R2713, Q33 Piana Giardino and Q 16 Acqua Buona upstream. 4 CONCLUSIONS The Ardea Basin is confirmed as a very important peri-Thyrrenian transfer structure at the intersection of the main local NE-SW fault system and the regional one (NW-SE trending), from reviewing new hydrogeologic, stratigraphic and geochemical data. The structural patterns exert powerful consequences on the fluid geochemistry patterns: deep fluids are emitted just along the intersection sites of the Ardea Basin master faults. The recently drilled ARDH well 113

Figure 2. Geologic profiles of the Ardea Basin (see text)

taps up to date the groundwater nearest to the deep reservoir end member located inside the Alban Hills carbonate basement. Here new geothennometric constrains allow us to state the presence at depth of a medium enthalpy reservoir (220 “C). The new geochemical and hydrogeological data gathered during the 1999 survey allowed us to reconstruct the steps of aquifer depletion in the last 40 years. As a whole the situation is less severe than foreseen, compared to other sectors of the volcano. Aquifer depletion is not widespread, being 40 m at the maximum within the sectors of S. Procula-Via Laurentina, where fortunately the pCO2 is not high enough to reach hazardous NGH thresholds. We did not find, as the Ciampino-Marino area, sectors characterized by a high NGH, excluding the Ardea and Brucona villages and the vicinity FSGD and ZFR pools. Here recommend a Rn indoor survey in and a CO2 flux survey.

Chiodini, G. & F. Frondini 2000. Carbon dioxide degassing from Alban Hills volcanic region, Central Italy. Chem. Geol., in press. Delaney, P., Amato, A., Borgia, A., Chiarabba, C. & F. Quattrocchi 1996. The Restless Volcano of the Alban Hills, Central Italy. Proc. AGU Meeting 1996, San Franc., U.S.A. De Rita, D., Funiciello, R. & M. Parotto M. 1988. Carta Geologica del Complesso Vulcanico dei Colli Albani. Prog. Fin. “Geodinamica”, CNR, Rome, Italy Duchi, V., Paolieri, M. & A. Pizzetti 1991. Geochemical study on natural gas and water discharges in the Southern Latium (Italy): circulation, evolution of fluids and geothermal potential in the region. J. Volc. Geoth. Res. 47: 22 1-235. Faccenna, C., Funiciello, R., Bruni, A., Mattei, M. & L. Sagnotti 1994. Evolution of a transfer-related basin: the Ardea Basin (Latium, central Italy). Basin Res. 6: 35-46. Foumier, R.O. 1987. Conceptual models of brine evolution in magmatic-hydrothermal systems. In ‘‘ Volcanism in Hawaii, U.S.G.S. Prof Paper N. 1350, Chapter 55: 1487-1506. Giggenbach, W.F., Minissale, A.A. & G. Scandiffio 1988. Isotopic and chemical assessment of geothermal potential of the Colli Albani area, Latium region, Italy. Applied Geochemistty 3: 475-486. Kemer, D.B., Marra, F. & P.R. Renne 2000. The history of Monti Sabatini and Alban Hills volcanoes: groundwork for assessing volcanic-tectonic hazards for Rome. J. Volc. Geoth. Res., in press. Mancini, C., Quattrocc:: F., Guadoni, C., Pizzino, L. & B. Porfidia 2000. study throughout different seismotectonical areas: comparison between different techniques for discrete monitoring. Ann. Geof 43 (1): 1- 28. Pizzino, L., Galli, G., Mancini, C., Quattrocchi, F. & P. Scarlato 2000. Natural Gases Hazard (COz, z22Rn)within a quiescent volcanic region and its relations with seismotectonics: the case of the Ciampino-Marino area (Alban Hills volcano, Rome. Natural Hazard, in press. Quattrocchi, F. & G. Venanzi 1989. Sulla scelta di un sito per il monitoraggio di parametri idrogeochimici per 10 studio di premonitori sismici nell’ area dei Colli Albani. Proc. 8”’ GNGTS Meeting, 1989, CNR - Roma: 259-266.

REFERENCES Amato, A., Chiarabba, C., Cocco, M., Di Bona, M. & G. Selvaggi 1994. The 1989-1990 seismic swarm in the Alban Hills volcanic area, central Italy. J. Volc. Geoth. Res. 61: 225-237. Capelli, G., De Rita, D., Cecili, A., Giordano, G., Mazza, R., Rodani, S. & Bigi G. 1999. Cities on volcanoes: groundwater resources and management in a highly populated volcanic region. A GIS in the Colli Albani region, Rome, Italy. Proc. of the Intern. Symp. Eng. Geology, Hydro. and Natural Disaster with emphasis on Asia. Kathmandu, Nepal, September, 28-30, 1999, NGSIAEG-IUGS- IDNDR.

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Response of an artesian well in southern Armenia to the 1400 km distant Izrnit earthquake of August 17,1999 H.Woith, R.Wang, C.Milkereit & J.Zschau GeoForschungsZentrum Potsdam, Germany

U .Maiwald & A .Pekdeger FU Berlin, FB Geowissenschaften, Germany

ABSTRACT: Co- and postseismic variations of the specific electrical conductivity were observed at an artesian well in southern Armenia after the M=7.4 Izmit earthquake of August 17, 1999. The epicentral distance was 1400 km. The static strain at the observation site was calculated to 1.3 x lO-", which is quite low when compared to periodic strain variations in the order of 10" induced by the earth tides. Nevertheless, the conductivity dropped by 3 % within one hour after the earthquake, further decreasing by another 3 % during the consecutive 3 weeks before it slowly recovered within a few months to reach the pre-event value. Additionally, the water discharge increased by 25 ?o' immediately after the earthquake. We developed a model that explains the observations in terms of a changed mixing ratio between two fluid systems as a result of a pressure disturbance triggered by the seismic waves passing the observation site. 1 INTRODUCTION The reaction of hydrogeological systems to earthquakes is a part of the earthquake research project READINESS (Realtime Data Information Network in Earth Sciences Woith et al. 1998) which aims to detect seismo-tectonically driven changes of physico-chemical properties of thermal and mineral waters. Most thermal/mineral waters are mixtures of fluids, usually with deep and shallow water components, so continuous monitoring of physicochemical ground water parameters should enable the detection of compositional anomalies related to hydrologic pressure changes in the different fluid the reservoirs. 2 TECHNICAL SETUP

READINESS monitoring network The present configuration of the monitoring network is shown in Figure 1. Ten thermal and mineral water sources in Turkey, Armenia, and Israel are being monitored continuously for water temperature, specific electrical conductivity, pH, redox potential, water/gas quantity/water level, and radon activity. The sampling interval is between 5 and 15 minutes. Every hour the data are transmitted via satellite to the partner institutions. The Armenian site KAT was known to be a highly sensitive spot responding to all major earthquakes that took place in the ArabianEurasian collision zone during the past 35 years

(Balassanian 1999) and was thus proposed by the scientific group of the National Survey for Seismic Protection, Yerevan as a promising target for continuous ground water monitoring.

Monitoring site KA T The monitoring site KAT (46.21°E, 39.16"N, 1495 m above sea level) is located near to the mining village Kajaran in the southern part of Armenia close to the Iranian border. Although convergent tectonics predominate in the Caucasus collision zone (Reilinger et al. 1997), the area of Kajaran is located in one of the most structurally complex regions of Armenia which includes graben structures. Typical geological features are widespread magmatic intrusions (Balassanian 2000). About 35 years ago a borehole was drilled through 18 m of alluvial deposits into volcanic rocks to a depth of 147 m. Mineralised fluid enters the borehole from the fractured aquifer at a depth of about 80 m. The uppermost 18 m of the borehole are cased with a steel tube of approximately 12 cm diameter. The rest of the borehole is unlined with a diameter of 9 cm. Although the name of the borehole is "Lernadzor-2" it will be referred to as KAT throughout the text. In 1983 a second borehole - ("Lernadzor-13/83',, labelled as LER in the following text) - was drilled at the same location to the same depth. Although the lateral distance between both wells is only 4 m, the produced fluids differ markedly in their composition. Both boreholes 115

Figure 1. Earthquakes (circles) and READINESS monitoring sites (diamonds) in the eastern Mediterranean. The “beachball” depicts the location and focal mechanism of the August 17. 1999 M=7.4 Izmit earthquake. The distance between the epicentre and the Armenian monitoring site KAT is about 1400 km.

are artesian, the water is free flowing without any additional pumping at rates of 7.2 litre/minute at KAT and 0.3 litre/minute at LER. Whereas KAT produces a water with a conductivity of around 1.36 mS/cm at a temperature of 15.4 “C, the corresponding values for LER are 4.48 mS/cm and 11.2 “C, respectively. Furthermore, LER has a gas content of about 50% free CO2, The water from KAT is Ca-Na-HCOj-SO4 dominated (Fig. 2) with relatively high radon and helium concentration of about 1500 Bq/l and 7900 1o - ml/I, ~ respectively.

Figure 2. Chemical composition of ground waters issuing from the artesian wells KAT and LER, as well as from a nearby river. The diameter of the inner circles is proportional to the specific electrical conductivity.

3 CO- AND POST SEISMIC ANOMALIES CO- and postseismic variations of the specific electrical conductivity as well as the flow rate were observed at KAT related to the M=7.4 Izmit earthquake of August 17, 1999. The epicentral distance was 1400 km.The conductivity dropped by 40 pS/cm (= -3 %, the accuracy of the measurements being *0.5 %) within one hour after the earthquake, further decreasing by another 40 pS/cm during the consecutive 3 weeks (Fig. 3) before it slowly recovered within a few months to reach the preevent value. The flow rate increased from 1.6 to 2.0 litre/minute (= +25 %) within 10 minutes after the earthquake (Fig. 4). After another 10 minutes the flow rate was reduced to 1.75 litre/minute. About 4 hours after the event the flow rate stabilized at 1.9 litre/minute. Whereas the water discharge increased immediately after the earthquake, i.e. within the sampling interval of 10 minutes, the decrease of the conductivity starts 70 minutes after the earthquake (Fig. 4). Nevertheless, we call this drop ”co-seismic” to distinguish it from a nearly exponential “postseismic” decay during the succeeding 3 to 4 weeks. The time delay is due to the fact that the borehole volume has to be replaced before any changes in the water quality are detectable at the surface. The described earthquake related anomaly is not a unique phenomenon for the site KAT. Between April 1995 and October 2000, in total 8 similar 116

members in this mixing scenario are the river water and the fluid issued at LER (see Table 1). Table 1. Water temperature, specific electrical conductivity, and pH of a nearby river, a shallow spring, KAT and LER. river spring KAT LER Temperature "C 5.4 8.8 15.4 11.2 Conductivity pS/cm 480 901 1360 4480 PH 7.47 7.25 6.23 6.08 Figure 3. Response of the specific electrical conductivity and the discharge of the artesian well KAT to the Izmit earthquake of August 17, 1999, which occurred at a distance of 1400 km. The resolution of the conductivity meter is I pS/cm; its accuracy is *0.5%. The flow meter measures about 15% of the total water discharge in a by-pass to the main flow.

anomalies were recorded (Woith et al. 2001). What could be the physical mechanism which changes the chemical composition of a mineral water in response to a distant earthquake? Is it the strain or the passing seismic wave shaking the local hydrological system? The static strain at the observation site KAT was calculated to 1.3 x 10'" for the Izmit event, which is quite low. Furthermore, the earth tides cause periodic variations of the specific electrical conductivity in the order of 0.1 % at KAT (Woith et al. 1998), whereas the earthquake induced strain in the order of 10-" reduces the conductivity by 6 %. From water level andor water temperature monitoring similar discrepancies have been observed before (Bower 1989, Igarashi & Wakita 1991). We are developing a physical model which is able to explain the observations based on a combined geological-hydrochemical model (Wang et al. 2001, Woith et al. 2001).

Mixing 78 % of the river water with 22% of the mineral water produced by LER would yield a conductivity of 1360 pS/cm as observed from KAT. A similar mixing ratio is obtained for the gas phase. We used the thermo-dynamic program PHREEQC (Parkhurst et al. 1980) to model a mixture of the river water with the highly mineralised LER water. The measured pC02 of KAT plots on the mixing line where 30 % deep fluid mix with 70 % of the shallow water. In order to decrease the conductivity of KAT by 3 %, it would be necessary to change the mixing ratio by 1 % in favour of the shallow water component (Fig. 5).

Figure 5. Relative conductivity variations at KAT as a fbnction of the mixing ratio between the river water and the mineral water of LER.

Figure 4. Co-seismic behaviour of the water discharge and the specific electrical conductivity related to the Izmit earthquake. The sampling interval at KAT is 10 minutes.

4 MIXINGMODEL We assume that the mineral water issued at KAT is a mixture of a shallow ground water component with a deep fluid, which is gas rich (C02). Possible end

The question then becomes one of: how can a distant earthquake change the mixing ratio by lm2%? The pressure in the shallow ground water system has to be increased relative to the deep water reservoir. This is unlikely, because at a distance of 1400 km both reservoirs should be affected in a similar way. But if we assume that the cold water circulates in a macro fracture whereas the hot fluid migrates along micro fractures by diffusion entering the main, water bearing fracture, an explanation and model of the 117

Potsdam. We thank an anonymous reviewer for valuable comments and suggestions.

observations becomes evident. According to (Wang et al. 2001) a pressure increase in the macro fracture of several hundreds Pa is enough to change the mixing ratio by 1 % in favour of the shallow water component thus reducing the conductivity of the mixed fluid by 3 %. Can a passing seismic wave increase the pore pressure? We propose a mechanism called “advective overpressure“ (Sahagian & Proussevitch 1992). This mechanism requires the existence of a free gas phase, i.e. bubbles. Let us assume that some of the gas bubbles are attached to grain boundaries. The seismic wave will disturb the rock matrix causing gas bubbles to leave their places on the grain surfaces and start to rise. Because the system is largely confined, the rising bubbles will be trapped at the top of the fractured zone (10 m of hydrothermally altered, fractured rocks according to the geological profile) and thereby lead to a significant pressure increase. We do not have direct evidence for the existence of a free gas phase at a depth of 80 m. The carbonate system is disturbed. The pH, pC02, H2C03, and HC03- are not in equilibrium. An enthalpy-chloride-diagram after Nicholson (1993) indicates steam loss for the reservoir fluid of LER. The steam including CO2 might enter the macro fracture system short before the mixed fluid enters the borehole leaving no time for an equilibration.

REFERENCES Balassanian, S. (2000) Earthquake prediction research for current seismic hazard assessment. In: Earthquake Hazard and Seismic Risk Reduction, Vol. 12 (Ed. by S. Balassanian, A. Cisternas and M. Melkumyan), pp. 169209. Kluwer Academic Publishers, Dordrecht/Boston/London. Advances in Natural and Technological Hazards Research. Balassanian, S. Y. (1999) The anomalous daily dynamics of local geophysical and geochemical fields (ADF) effect study in the connection with earthquake preparation and occurrence. In: Spontaneous global& synchronized variations of physical parameters., Vol. 24 (Ed. by Rokityansky), pp. 74 1-752. Elsevier Science, Oxford, United Kingdom. Physics and Chemistry of the Earth. Part A: Solid Earth and Geodesy. Bower, D. R. (1989) Tidal and coseismic well-level observations at the Charlevoix Geophysical Observatory, Quebec. Tectonophysics, 167(2-4), 349-36 1. Igarashi, G. & Wakita, H. (1991) Tidal responses and earthquake-related changes in the water level of deep wells. Journal of Geophysical Research, 96(B3), 4269-4278. Nicholson, K. (1993) Geothermal Fluids, 263 pp., Berlin: Springer Verlag. Parkhurst, D. L., Thorstenson, D. C. & Plummer, L. N. (1980) PHREEQE - a computer program for geochemical calculations, 210 pp., US Geological Survey. Reilinger, R. E., McClusky, S. C., Souter, B. J., Hamburger, M. W., Prilepin, M. T., Mishin, A. A., Guseva, T. & Balassanian, S. (1 997) Preliminary estimates of plate convergence in the Caucasus collision zone f?om Global Positioning System measurements. Geophysical Research Letters, 24( 14), 1815-18 18. Sahagian, D. L. & Proussevitch, A. A. (1992) Bubbles in volcanic systems. Nature, 359,485. Wang, R., Woith, H., Milkereit, C. & Zschau, J. (2001) Modeling ground water anomalies induced by distant earthquakes. Journal of Geophysical Research (in preparation). Woith, H., Milkereit, C., Wang, R., Zschau, J., Igumnov, V. A., Balassanian, S., Maiwald, U. & Pekdeger, A. (2001) Coand postseismic conductivity anomalies monitored at an artesian well in southern Armenia. Journal of Geophysical Research (in preparation). Woith, H., Milkereit, C., Zschau, J., Maiwald, U. & Pekdeger, A. (1 998) Monitoring of thermal and mineral waters in the frame of READINESS. In: Water Rock Interaction, Vol. WRI-9 (Ed. by G..B. Arehard and J. R. Hulston), pp. 8098 12. A.A. Balkema, Rotterdam.

5 CONCLUSIONS Some hydrogeological systems are very sensitive to strain since it is observed that earth tides induce periodical variations of the electrical conductivity of the waters. Most natural mineral and thermal waters are mixed fluids and the mixing ratio is likely to be sensitive to pressure disturbances in either of the reservoirs. The higher the compositional contrast between the two mixed components, the more evident the response. If a free gas phase is present, the pore pressure within a hydrogeological system may be increased by small disturbances such as a passing seismic wave generated at large distances away from the monitoring site. The site KAT in southern Armenia fulfills all the criteria listed and is thus a good monitoring site for possible earthquake precursors. ACKNOWLEDGEMENTS Many thanks are due to the scientific and technical staff of the Armenian National Survey for Seismic Protection, Yerevan. Financial support was given by the GeoForschungsZentrum

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Discrete and continuous monitoring of groundwaters in the seismic area of the Umbria region (Italy) A .R .Zanzari Dipartimento Scienze della Terra, Perugia, Italy

A .Martinelli Agenzia Regionale per la Protezione Ambientale (ARPA-Umbria),Perugia, Italy

R Cioni , M .Guidi, B .Raco & A Scozzari IGGI-CNR, Pisa, Italy

F. Quattrocchi & G.Galli Istituto Nazionale di Geojisica, Rome, Italy

C .Mancini DINCE Universita “La Sapienza ”, Rome, Italy

ABSTRACT: this paper presents the Geochemical and Hydrologic Monitoring of Regional Groundwater for the Definition of the Response of Hydrologic Systems to Earthquakes (MICRAT) project and discusses its preliminary results. The objective of the project is to monitor water points in the seismic areas of the Umbria region (Central Italy) with a view to defining the aquifer’s response to seismic activity. In order to obtain such kind of information discrete sampling of water twice a month and gas every month, and continuous data acquisition from a network of 8 automatic monitoring stations at selected sites are being carried out since April 2000. Preliminary results of gas analyses are: 1) springs with rapid circulation in carbonate aquifers show atmospheric dissolved gases in equilibrium with both temperature and altitude of the recharge area and presence of some excess air; 2) waters from very deep or from reducing carbonate levels bearing aquifers have C€& and H2S dissolved and an atmospheric gas content lower than expected, but a higher ArN2 ratio which suggests a stripping process due to bubbling gas. 1 INTRODUCTION AND AIM OF THE PROJECT The area between the Umbria and Marche regions of Central Italy belongs to the Appennine mountain chain, a geological environment characterised by fiequent seismic activity which has caused in the past and even very recently considerable damage and a great number of casualties. As a result of major earthquakes the local aquifers suffer some changes in the potable quality of their exploited waters as well as in their flow and geometry. Owing to the importance of those aquifers to provide a drinking water supply, the Government of the Umbria region has financed the MICRAT project for the hydrologic and geochemical monitoring of some main water points (springs and wells). The objective of this project is to identify the kind of hydrologic and geochemical response of the aquifers to seismic stress and to define a protocol of permanent monitoring of the water points to be applied in the future. This objective will be achieved through: 1) the installation of a network of stations for the continuous monitoring of hydrologic, physico-chemical and chemical parameters and 2) periodic analyses of selected water points (twice a month) and of dissolved gas (every month).

2 CONTINUOUS MONITORING NETWORK The development of continuous monitoring instrumentation is a scientific open field since the 196Os, after the first discovery of the possible use of Radon as an earthquake precursor and stress-strain variation pathfinder (Thomas 1988, Galli et al. 2000). 2.1 Monitoring ofphysico-chemical and chemical parameters The water stations have been designed and assembled in such a way as to obtain the optimal signal stability. Particular care has been devoted to the design and building of an amplifier for selective electrodes whose drift will be 0.02 unitdmonth or less in order to give significance to small pH variations induced by changes of CO2 partial pressure in the water. They can measure up to a maximum of seven parameters. One of the stations is devoted to the detection of free gases. The measured parameters are summarised in Table 1. 2.2 Radon monitoring The monitoring stations measure Radon concentration in groundwater by revealing a particles emitted 119

by 222Rnand its short lived daughters 218Poand 214P0 by means of a scintillation chamber developed by DINCE and ING coupled with a photomultiplier (Qattrocchi et al. 2000). 222 Rn is quantitatively stripped by bubbling air from a fixed volume of water and sucked into the scintillation chamber. The apparatus performs a complete automated cycle and data acquisition every 6 hours. The Radon and water stations have common power supply and data logger. For calibration purposes water enriched with Radon (- 800 Bq/l) is been split to the continuous Radon monitor and to calibrated devices such as Marinelli Beakers or Degassing Portable Systems (Galli et al. 1999). The first Rn concentration data in Capovena groundwater have been plotted versus time in figure 1. Table 1. Continuotts stations and i i i e a s i i r e d ~ - - * ~ ~.--~--*~..-..-..,...,.... .!?!ace!!.?!??e ........A!??!Ysed.!u!d ................Yeas!!red.Pra!!?eters ........... P

T, pH, Cond., Eh, dissolved CO2, CH4,222Rn Fonteccliio Mineral water well Cond., T, pH, Eh, p C 0 ~ T, pH, Cond., Eh, pC02 Mocaiaila well Capodacqua spring pH, Eh Capovena spring pH, Eh 2 2 2 ~ i Bagiiara spring T, pH, Eh, Cond. Clihuuio spring T, pH, Eh, Cond 4 Air, CO2, CH4 Triponzo

Thermal water well

3 GEOCHEMICAL STUDY Figure 1 Variations with time of T, pH, Eh and 222Rnconcentration in Capovena groundwater; the voltage supplied by the battery pack is also shown.

3.1 Water chemistry

The monitored area (Fig. 2) extends along the central Appennine ridge where important aquifers feed the main regional water systems. It is possible to distinguish three main types of aquifers separated by impervious clayey levels. The first is characterised by fractured limestone and gives waters with low salinity and a marked calcium carbonate character. The second type is represented by early Jurassic-late Triassic formations (Barchi et al. 1996). Here, the waters have a calcium sulphate-carbonate composition and are characterised by longer circulation, higher salinity and higher contents in magnesium and strontium as compared to the first type. Water mixing of both aquifers also occurs. The third type is a very deep aquifer in Triassic evaporites giving high salinity waters with a sulphate-chloride character (Giaquinto et al. 1991, Quattrocchi et al. 2000). Two water points of the third type have been studied: the first (Triponzo) is a sulphate- prevalent thermal water with a reducing character, while the second (Stifone) is a very large flowing spring with a high salinity (NaCl)-prevalent character and high pC02 (Frondini 1995).

In the northern part of study area the outcropping formations are fluvial and lacustrine sediments, a sandy-mar1 formation and limestone. The geochemical character of the waters reflects the lithological nature of these aquifers with medium salinity, high ratio Mg/Ca and relatively high pCOz. The Fontecchio spring has a special chemistry with high Na-HC03 content. Twenty four water points representative of the main aquifers with extreme chemical characters have been considered. They have been selected on the basis of their regional importance and their proximity to active faults (Boncio & La Vecchia, 2000, Boncio et al. 2000, Lavecchia et al. 2000). The main characters of analysed waters are represented in the Langelier-Ludwig diagram of figure 3 . 3.2 Dissolved gas study

As the mobilization of a gaseous phase can be expected during the accumulation and release of 120

Figure 3. Langelier-Ludwig diagram for the monitored springs. Figure 2 . Map of inoiutored points. Full symbols represent continous monitoring stations.

tectonic stress, a systematic study of the distribution of dissolved gases in selected water points is being carried out. So far more than 60 gas samples have been collected and analysed for Ar, N2, CH4 and CO2 by gaschromatography. All samples have N2 as the main component except those of Stifone (CO2 dominated) and Fontecchio (CH4 dominated). Referring to the Ar versus N2 graph (Fig. 4) some relevant observations can be made: -The majority of samples plot along a narrow belt which, starting from the theoretic contents of Ar and N2 dissolved in infiltrating water at the average temperature and pressure conditions of the recharge area mean altitude, follow a trend towards the characteristic point for air. -The amount of excess air varies from one water to another reaching rather high values. This can be related to the circulation rate of the infiltrating meteoric water which is expected to be higher in the absence of water saturated zones within the aquifer (Cioni et al 1998). -A few points with a water chemistry far from the typical character of carbonate aquifers show different behaviour. They are: a) Triponzo spring (discharge temperature 29"C, high salinity and slightly reducing character) has a very small amount of excess air, a presence of Methane and a calculated recharge temperature between 13 and 17 "C at mean recharge altitude.

121

b) The Fontecchio water where the presence of a considerable content of methane, corresponding to a partial pressure a little greater than 1 bar, suggests the presence of degassing. The formation and separation of Cl& bubbles induces the stripping of other dissolved gases thus decreasing their content in the water. A simple stripping model explains the position of the points in figure 4. c) The Tili well which intercepts an artesian aquifer with a principal recharge area from the neighbouring limestone hills and is confined '-y impermeable formations containing levels rich of organic matter (Zanzari 1998). Dissolved gases show a total absence of 0 2 and a presence of CH4 and their position in the Arm2 plot suggests an enrichment of N2 together with CH4 from the reducing organic levels.

Figure 4. Ar versus N2 dissolved in the samples. Results of 7 different sinplings are reported.

4 CONCLUSIONS

Data collected by the monitoring stations show a substantial constancy of the physico-chemical parameters and of 222Rnduring the sampling period. Discontinuous chemical data of waters collected in the first 5 months show very small variations (e.g. pH and HC03) in only a few water points. Larger changes have been observed in two waters resulting from the mixing of a shallow component and a deeper sulphate component. Dissolved gases have provided relevant information on recharge conditions and on water circulation patterns. The effects of stripping of dissolved atmospheric gases, due to the release of a methane rich gaseous phase, have been recognised. The mobilisation of a reduced gas phase rich in CH4 and N2,probably due to decomposition of organic matter, seems to be responsible for the deviation of some samples from the regional behaviour. REFERENCES Barclii, M., Conversini, P. & G.S. Tazioli 1996. Schema idrogeologico delle eniergenze del Clitunrio e del Teinpio del clihiiuio, Uiiibria Orientale. Quad. geol. Appl. 23: 37-48 Boncio, P, Brozzetti, F. & G. Laveccliia 2000. Architecture and seismotectonics of a regional Low-Angle Nornial Fault zone in Central Italy. Tectonics. 19:1038-1055. Boncio, P. & G. Laveccliia 2000. A structural iiiodel for active extension in Central Italy. J. Geodynanrics 29:233-244. Cioni, R., Guidi, M., Bigazzi, G., Corbin, J.C., Cluodini, G., a Raco, B. & G. Magro 1998. Dissolved gas study. In “a multi-disciplinary global approach of groundwater flows in karstica areas and its consequences for water resources and environment studies”. The Katrin Project, EU contract ERR-C’HRY--C‘T91-0567 &f. M. A4onnin Ed.), Final Report: 8-1 1, 90-97, 124-136, 180-185. Frondini, F. 1995. Geocheinistry of ground water in SoutliCentral Umbria. Plinius 13:79-83. Galli, G., Guadoni, C. & Manciiu, C. 1999. Radon grab sampling in water by means of radon transfer in activated clwcoal collectors. Proceedings of the Fourth International Conference on Rare Gases Geochetriistry. 22 : 583-587. Galli, G., Maiicini, C. & Quattrocclii, F. 2000. Groundwater radon continuous monitoring system (a scintillation counting) for natural hazard surveillance. Pure and Applied Geophysics. 157:407-433 Giaquinto, S., Marclietti, G., Martinelli, A. & E. Martini 1991 (eds). Le acqiie sottemiee in Umbria. Petugia: Protagon. Laveccliia, G., Boncio, P. & F. Brozzetti 2000. Analisi delle relazioiii tra sisnuciti e Struthire tettoniclie in UmbriaMarclie-Abnlzzo fiiializzata alla realizzazione della niappa delle zone sisinogeneticlie. In G. Galadini, C. Meletti & A. Rebez (ed), Le ricerche del GNDT nel caiiipo clella pericolosita sistnica. Quattrocclii, F., Di Stefano, G., Pizzino, L., Pongetti, F., Romeo, G., Scarlato, P., Sciacca, U.& G. Urbini 2000. Geochemical Monitoring System I1 prototipe (GMS 11) installation at tlie “Acqua Difesa” well, witlun the ETNA region: first data during the 1999 volcanic crisis. J. Volcanol. Geothernr. Res. 10 I: 273-306. Quattrocclii, F., Pik, R., Pizzino, L., Guerra, M., Scarlato, P., Angelone, M., Barbieri, M., Conti, A., Marty, B., Sacclu,

122

E., Zuppi, G. M., & Lombardi, S. 2000. Geocliernical clianges at tlie Bagni di Triponzo thermal spring during the Umbria-Marclie 1997-1998 sequence. Journal of SeisniolOS, 4: 567-587. Thomas, D. 1988. Geochemical precursors to seismic activity. Pure andApplied Geophysics, 126: 37-46. Zanzari A.R. (1998). Studio dei gas disciolti in pozzi e sorgenti uinbri appartenenti alla rete di inonitoraggio del progetto PRISMAS. Final Report: 1-9.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeifljnger, Lisse, ISBN 90 2651 824 2

Massflows of the subsurface hydrosphere: Global and regional cycles V.P.Zverev Institute of Environmental Geosience RAS, Moscow, Russia

ABSTRACT: Subsurface water massflows form three principal geochemical cycles of global circulation: hydrogeological, lithogenical and geological. Quantitative model of this global cycle of subsurface massflows is developed. Quantitative estimation of subsurface waters on regional levels are made for hydrogeological cycle for examples of several regions of Russia and for lithogenical one for example of Caspian Sedimentary Basin.

1INTRODUCTION Massflows of the subsurface hydrosphere are a principal factor of realization of water-rock interaction processes in the Earth Crust. Those massflows appearing under the influence of the Earth's gravitation field and heatflow from the Earth are continuous during development processes in the lithosphere.

2 GLOBAL CYCLES Subsurface water global circulation includes several cycles beginning with those of comparatively short duration in zone of active water exchange, to longterm cycles including deep horizons of the Earth Crust. Nowadays so-called hydrogeological circulation of subsurface water is mostly well investigated. It includes the massflow of free subsurface water from regions of infiltration to places of discharge into river basins, and directly into seas and oceans through coastline. Lithogenic cycle of subsurface water circulation means the burial of ocean and sea water with sediments in sedimentation processes, during their immersion into the Earth Crust in the process of sedimentary basins evolution, and their discharge in a free condition under changes of the system thermodynamic parameters. The quantitative evaluation of massflow of different natural water types participating in lithogenic

circulation is based on water dynamic process in the frames of separate Earth Crust blocks and water reservoirs. Such evolution is characterized by the development process in time commensurable with the period of its realization (Zverev 1993, 1994). Results of this estimation shown in Table 1,allow to define the relative role of separate massflow of subsurface water in lithogenic cycle where linked waters became free physically and chemically during the stages of hypergenesis, diagenesis, catagenesis, metagenesis, and metamorphism. The geological cycle is of great importance in the process of Earth evolution, connected with convective flow in the upper Earth mantle and lithosphere plates. The geological cycle of subsurface water circulation is formed by two important massflows. They are: the hydrothermal discharge in the limits of rift valleys of middle-oceanic ridges which is developed practically along the whole distance (60,000 km), and where the hydration of mantle rocks is realised. Some part of the water from hydrothermal activity in the limits of island arcs and active margins of continents is formed by the dehydration of the main magma and sedimentary rocks. Hydrothermal discharge in the limits of island arcs and active continental margins consists of: dehydration of sedimentary rocks of the first seismic layer (0.42~1O'~g/y), and deserpentinization rocks of the 2nd and 3rd seismic layers of oceanic crust (0.39~10'~g/y)in a processes of their plunging during subduction and rising up stream of subsurface water of meteoric origin (3.19~10'~g/y)(Zverev 1994, 1999).

123

Massflows of subsurface waters

Lithogenical

Geological

Hydrothermal

Mountain rock series of the Earth'Crust drawn in water circulation

Predominate condition of subsurfacewaters

Mass of subsurface water 1 0 ~ ~

continentalcrust. Long-term water exchange zone of continentalcrust. physically linked Sedimentary rocks of continental and subcontinental crust (clays and clay states). Sedimentary rocks of 1st seismic layer of oceanic crust. chemically linked Granite-metamorphic zone of continental crust. 2nd and 3rd seismic layers of oceanic crust. free, System of middle-ocean ridges. System of island arcs and active steam-watermixture continentalmargins.

The results of quantitative estimation of modern natural water circulation (Table 1) demonstrate that among the main massflows of subsurface waters the dominant role belongs to waters participating in the hydrogeological cycle. Their mass is more than: 3 orders higher then the flows participating in lithogenical cycle, and by 4 orders - in geological. The exception is hydrothermal discharge of the middle-oceanic ridges which is one of the main constituents of the geological cycle and still is 2 orders less than the hydrogeological flow.

Territory

Subsurface water exchange intensity g/y

0.065

0.49~10'~

0.051

4 . 4 2 1015 ~

0.098

0.42~10'~

0.17

0.041~10'~

0.052

0.39~10'~ 0.18~10'~ 4.00~ 1015

Area, 106km2

Subsurfacewater massflows g/Y g/y.km2

3.2 Lithogenical cycle

3 REGIONAL CYCLES 3.1 Hydrogeological cycle Quantitative estimation of subsurface water massflows at regional levels was made for the hydrogeological cycle considering as example the territory of Russia (Zverev 1999), the Volga river basin (Zverev et al. 1994), the platform Moscow artesian basin, and the mountain-belt south-west Caucasus (Zverev 1983) (Table 2). This estimation shows that the modulus of subsurface massflows varies from 52x109g/y-km2 to >4O0x1O9 g/y-km2. For all the continents this modulus is - 7 5 ~109g/y*km2. The value of modulus of massflows subsurface water is principally dependent on climatic and geological-structural conditions of the region.

124

Quantitative estimation of subsurface water on regional levels for lithogenical cycle was made for the Caspian Sedimentary Basin for 185 million year of its evolution. The Caspian Basin, corresponding to the depression filled with water, is 1200 km long and about 320 km wide. Its area is 374,000 km2. The northern, shallow part of the basin, with a water depth of 5 m, occupies an area of about 80,000 km2. The middle part represents an asymmetric depression with the area of 138,000 km2 and a maximum water depth of 788 m. The southern part of the Caspian Sea is a deep basin (with a maximum depth of 1025 m) and steep western and southern coastal zones. Its area is 156,000 km2. The total water mass in the Caspian Sea reaches 0 . 7 8 ~ 1 g. 0~~ Within the area presently occupied by the Caspian Sea, three principal geological-structural elements can be identified in the northern, central, western and southern parts.

Parameters

Area, 10' kmL Average thickness of sedimentary cover, km Volume of sediments, 1021cm3 Mass of sediments, 1021g Mass of water captured during sedimentation, 1O2O g Mass of free and physically linked water of the sedimentary cover, 1020g Mass of free and physically linked water released from the sedimentary cover, 1020g

Precaspian Syncline

43.2 10.5

Skythian Turanian Plate

Alpine folded Alpine folded zone of middle zone of the part of the southern part of Caspian Basin the Caspian Basin

145.6 6.0

41.6 12.5

143.6 21.5

All Caspian sedimentary Basin 374 -

0.454

0.874

0.52

3.087

4.9

1.135 3.76

2.185 7.75

1.3 4.32

7.717 24.89

12.3 7.4

0.512

1.253

0.315

5.316

40.7

3.248

6.497

4.005

19.574

33.3

They are respectively as follows: the southern portion of the Precaspian syncline; epivariscan Skythian-Turanian plate; and the zone of the Alpine orogeny. It makes sense to divide the latter into the northwestern part adjoining the eastern termination of the Greater Caucasus and the southern one, representing a megatrough underlain by basaltic basement. These structural subdivisions were used for the approximate estimation of the subsurface water mass in the sedimentary cover of the Caspian Basin, where the thickness of sedimentary deposits varies within a wide range: from 5-6 km on the Skythian-Turanian Plate to 30 km and more (Guliev et al. 1988) in the southern megatrough. Rocks of variable ages making up the consolidated basement are taken as the lower surface of the sedimentary deposits. The estimation of the free and physically bound subsurface water content was done assuming that all pores in the rocks are saturated with water. Variations of porosity in the main rock types were defined with the use of the analysis of the change of porosity in covers from Eastern Precaspian area adjacent to the Caspian Basin. The chemically bound water content of the sedimentary cover of the Caspian Basin was estimated using the previously obtained data (Zverev 1993) on the water weight percentage in the main rock types and its average values for platform and geosynclines. Thus, we estimate that the sedimentary cover of the Caspian Basin contains around 1 1 . 9 ~ 1 g0 ~of~ chemically and physically bound and free subsurface 125

water, of which the latter account for almost 7 . 4 ~ 1 0 ~g,' that is, by an order of magnitude greater than the mass of water in the Caspian Sea ( 0 . 7 8 ~ 1 0 ~ ~ g). It should be emphasized that the significantly greater part of these water ( 5 . 3 ~ 1 0 ~ 'g) is concentrated in the Southern Caspian Basin (Zverev et al. 1998). The history of the Caspian Basin is closely related to the evolution of its marine histories. The entire continuous sedimentary cover of the Southern Caspian Basin was formed in the presence of the liquid phase (oceanic and sea water). Sedimentation represents the conversion of mobilized matter which captures a considerable amount of physically linked water, which constitutes 72 and 40 % in the clay and in clastic and carbonaceous rock, respectively (Zverev & Kostikova 1999). As a result, we have evaluated the amount of water trapped in the sedimentary cover of Caspian Basin and water release of free and physically linked water from the sedimentary cover (Table 3). Results show that about 4 0 . 7 ~ 1 0g ~of~ free and physically linked water was accumulated during the 185-million year - long evolution of the basin. About 3 3 . 3 ~ 1 g0 ~of subsurface water was returned in the course of the basin development and about 6 . 2 ~ 1 0 ~g 'having been released only in the Southern Caspian during the last 5 million years. Thus the average rate of the free and physically linked water from sedimentary cover of the Southern Caspian for ~ the last 5 million years constitute 1 . 2 6 ~ 1 0 'g/y (Zverev & Kostikova 1999). The evolution history of the sedimentary basin shows that, in some

periods, the actual rate varied and differed significantly from the average values.

Zverev, V.P. & I.A. Kostikova 1999. Hydrogeological features of sedimentary Cover of the Southern Caspian Megabasin. Environmental Geoscience. 2: 140 - 146.

4 CONCLUSION Massflows of the subsurface water are a principal factor of realization of water-rock interaction in the Earth Crust. Three principal geochemical cycles of subsurface water global circulation are distinguished: hydrogeological, lithogenical and geological cycles. The subsurface water exchange on hydrogeological cycle costitutes 1 0 . 1 6 ~ 1 0g/y, ~~ lithogenical - 4 . 8 4 ~ 1 0 ’ ~g/y and geological 0 . 4 3 ~ 1 0 g/y. ’~ Quantitative estimation of subsurface water massflow of hydrogeological cycles on regional scales showed that the rates of subsurface massflows ~ changed from 52x109 to 4 0 0 ~ 1 0g/y*km2. The study of the subsurface balance of Caspian sedimentary basin showed that about 4 0 . 7 ~ 1 0g ~of~ free and ph sically linked water was accumulated and 3 3 . 3 ~ 1TO0 g was returned over 185 million years of evolution.

ACKNOWLEDGMENTS This work was supported by the State Program in Science and Technology “Global Changes in the Environment and Climate”, project no. 1.1.2.

REFERENCES Guliev, I.S., Pavlenkova, N.I. & M.M. Radgapov 1988. The Zone of Regional Decongoliolation in the Sedimentary Cover of the Southern Caspian Depression. Litol. Polezn. Iskop. 5: 130 - 136. Zverev, V.P. 1983. Rol podzemnyh vod v migracii chemitcheskich elementov (Role of subsurface water in migration of chemical elements). Moscow, Nedra. Zverev, V.P. 1993. Gidrogeochimia osadochnogo processa (Hydrogeochemistry of the Sedimentary Process). Moscow, Nauka. Zverev, V.P. 1994. Quantitative estimation of massflows of subsurface waters in the Earth‘s crust and hydrogeochemical balance of the Earth‘s surface. Mineralogical magazine. 58A: 1008 - 1009. Zverev, V.P. 1999. Massopotoki podzemnoy Gidrosfery (Massflows of the subsurface Hydrosphere). Moscow, Nauka. Zverev, V.P., Zolotych, E.O. & N.V. Kiseleva 1994. Subsurface chemical runof in the river Volga basin. Geoekologia. 3: 42 - 51. Zverev V.P., Varvanina, 0.Yu. & I.A. Kostikova 1998. Quantitative Estimation of the Groundwater Content of Caspian sediments. Environmental Geoscience. 1: 355 359.

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Model 1ing water-rock interaction

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

H,S -controlling reactions in clastic hydrocarbon reservoirs from the Norwegian Shelf and US Gulf Coast P.Aagaard & J.S.Jahren Department of Geology, University of Oslo, P. OBox 1047, Blindern, N-0316 Oslo, N o w a y

S .N.Ehrenberg STATOIL, Forusbeen 50, N-4033 Stavanger, Norway

ABSTRACT: The concentrations of H2S in clastic hydrocarbon reservoirs from the Norwegian Shelf (Hiland et al. 1999) and the US Gulf Coast (Smith 1997) exhibit a steady increase with burial depth and temperature. Various potential mineral reactions which may control the H2S level have been explored and compared with the reported H2S abundances. These buffers were calculated based on the general geochemical control on burial diagenesis discussed previously by Aagaard & Egeberg (1998). The pyrite-magnetite buffer predicts too low H2S concentrations, while the pyrite-kaolinite-chlorite buffer on the other hand gives too high levels. Instead it appears that the observed H2S concentrations are buffered by iron sulfide-carbonate assemblages within the reported variation of the FeC03 component. This behaviour contrasts the anhydrite containing reservoirs which may produce high levels of H2S by reactions with hydrocarbons. 1 INTRODUCTION

2 COMPUTATIONAL APPROACH AND RESULTS

The occurences of hydrogen sulfide (HlS) in hydrocarbon reservoirs has received considerable attention in deep carbonate systems (Orr 1977, Worden & Smalley 1996, Wade et al. 1989, Dakhnova et al. 1993), where high concentration levels may reduce the economic viability. Controlling reactions appear to involve anhydrite and hydrocarbon species (specially methane). In clastic hydrocarbon reservoirs hydrogen sulfide levels are normally much lower and less is known about the origin and controlling mechanisms. The aim of the present communication is to present new data of H2S from the Norwegian Shelf (HBland et al. 1999) and the US Gulf Coast (Smith 1997) and analyse potential mineral gas reactions which may buffer the H2S fugacity. The geochemical characteristics of formation waters from hydrocarbon reservoirs on the Norwegian shelf have been reviewed and discussed previously (see Aagaard & Egeberg 1998). However, the equilibrium status of the iron containing phases, has not yet been analysed to any extent. This is partly due to uncertainty in thermodynamic data of some iron mineral components, but not least due to previous missing information about the fluid composition. We will here combine the new H2S data with our general system understanding of sandstone diagenesis and formation water constraints (Bjmlykke et al. 1995, Aagaard et al. 1992, Aagaard & Egeberg 1998).

2.1 Computational approach We have tested a variety of mineral gas reactions to map out potenial H2S buffers. The computations applied the thermodynamic data base and the SUPCRT92 computer program of Helgeson and coworkers (Johnson et al. 1992) and also utilised the carbon dioxide trend of Smith & Ehrenberg (1989) from the Norwegian shelf.

2.2 Redox conditions The redox condition under burial diagenesis also needs to be evaluated and quantified. We have followed the approach and arguments of Helgeson et al. (1993) and Shock (1994), where the oxygen (and hydrogen) fugacity is assumed to be buffered by the acetic/propionic acid equilibrium. This assumption appears reasonable as homogeneous aqueous species reactions normally reach equilibrium faster that those involving solids. All calculations were done along an average North Sea P,T-gradient, taken from Smith & Ehrenberg (1989). The organic acid data of Barth (1991) and Barth & Riis (1992) from North Sea oil field brines were used. The resulting 02fugacity are depicted in Figure 1, following a trend between the hematite-magnetite and magnetitepyrite-pyrrhotite buffers. A corresponding H2fugacity was also established. 129

Giggenbach (1980), in his work on gas abundances and distribution in hydrothennal systems in New Zealand, advocated control by equilibria between pyrite, aluminosilicates and iron containing aluminumsilicates to buffer the hydrogen sulfide partial pressure in the low temperature range. Giggenbach suggested either chlorite or epidote to be the potential candidates. Thus in the diagenetic regime authigenic chlorite, which is frequently observed (Curtis et al. 1985, Jahren & Aagaard 1989, 1992, Hillier & Velde 199l), may be the iron containing aluminumsilicates involved in H2S reactions. Authigenic chlorite also exhibits several signs of dissolutiodprecipitation behavior, such as grain coarsening (Jahren 1991). Carbontes are another potential group of reactive iron-containing minerals. Diagenetic calcite and dolomite are frequently ferroan (Boles 1978, Reksten 1990a, b, Milliken 1998, Morad et al. 1998). We have tested various mineral buffers for our diagenetic systems, but the ones that embracketed the observed H2S partial pressure range were the pyrite-kaolinite-chlorite, the pyrite-Fecarbonates and the pyrite-magnetite buffers. These mineral buffers can be expressed by the following equilibrium reactions:

Figure 1. Oxygen fugacity in hydrocarbon reservoirs from the Norwegian Shelf calculated from organic acid data reported by Barth (1 99 1) and Barth & Riis (1 992). The data plot between the hematite-magnetite buffer (HM) and the magnetite-pyritepyrrhotite buffer (MPyPo).

Figure 2. The fugacity stability diagram of the Fe-S-0 system at 150°C and 750 bar. The marked rectangle depicts the fH2 calculated from the propionidacetic acid buffer and the observed variation in fHHZs.

2.3 Potenial mineral H2S buffers There are several potential mineral buffers which may control the fugacity, but it is always useful1 to start with the iron sulfide / - oxide system. Figure 2 depicts the stability fields of the most common minerals in this system, as a function of H2S- and H2fugacities. Apparently, the variation in H2S fugacity is not effectively controlled by the pyrite-magnetite buffer, but may involve other minerals as well. However, pyrite appears to be a stable mineral phase in these diagenetic systems.

FeS2 + CO2 + H2 + H20 = FeC03 + 2H2S

(2)

3FeS2 + 2H2 + 4H20 = Fe304+ 6H2S

(3)

The fugacity of hydrogen and carbon dioxide are given by the above assumptions. For all these reactions, the H2S fugacity is expressed as a function of the corresponding equilibrium constant and the activity of mineral end member components. Gibbs free energy data of chamosite (the iron end member chlorite) were adopted from Stoessel (1984, data set no. 4). In these calculations pyrite and magnetite were assumed pure and at their stable form. Thus, the activity of the chamosite end-member in chlorite and the siderite end-member in carbonates are the only other variables that will effect the H2S fugacity. The resulting H2S buffer lines are depicted in Figure 3 , as isopleths of unit activity of mineral components. 2.4 Comparison with reported P(H2S) In order to compare with the observed H2S partial pressures in clastic reservoirs on the Norwegian Shelf and the US Gulf Coast, the buffer lines of Figure 3 needed to be converted to corresponding pressure lines. This was done by calculating the H2S fugacity coefficient, l?(H2S), according to Reid et al. (1987) along the P,T trend of the Norwegian Shelf

130

actions between pyrite and carbonate minerals may control the H2S level in clastic reservoirs. The three isoplets for this buffer match well with reported iron content of calcite, dolomite/ankerite (Boles 1978, Reksten 1990a, b, Milliken 1998, Morad et al. 1998). Relative to the reported H2S values, magnetite appears to be unstable, whereas the pyritekaolinite-chlorite assemblage will buffer the H2S at too high values. However, as can be seen from equation (l), the low value of H2S will drive the reaction towards the right, stabilizing the chamosite component. This is in accord with formation of diagenetic iron rich chlorites.

3 CONCLUDING REMARKS Figure 3. Potential H2S mineral buffers in burial diagenesis of clastic hydrocarbon reservoirs. The calculations are based on the O2(or H2) fugacity trend of Figure 1 and unit activity of the chamosite and siderite end-members of chlorites and carbonates.

Clastic reservoirs in oil prone sedimentary basins have H2S concentrations which appear to be buffered by iron sulfide-carbonate assemblages. This behaviour contrasts the anhydrite containing reservoirs which may produce high H2S levels by reactions with hydrocarbons.

ACKNOWLEDGEMENT We are grateful to Shell E&P Technology Company (Houston, Texas) for allowing use of data and theory from an unpublished 1980 report by John T. Smith, who presented a model for how mineral-hydrocarbon equilibrium controls the trend of increasing H2S partial pressure with temperature in Gulf Coast clastic sediments.

REFERENCES Aagaard, P. & P.K. Egeberg 1998. Formation waters and diagenetic modifications: General trends exhibited by oil fields from the Norwegian shelf - A model for formation waters in oil prone subsiding basins: In Arehart & Hulston (eds), Water Rock Interaction, Taupo. Aagaard, P., Egeberg, P.K. & J.S. Jahren 1992. North Sea clastic diagenesis and formation water constraints: In Kharaka & Maest (eds), Water Rock Interaction, Park City. Barth, T. 1991. Organic acids and inorganic ions in waters from petroleum reservoirs, Norwegian continental shelf: a multivariate statistical analysis and comparison with American reservoir formation waters. Applied Geochemist v 6 : 1-16. Barth, T. & M. Riis 1992. Interaction between organic acid anions in formation waters and reservoir mineral phases. Org. Geochem. 19: 455-482. Bjlarlykke K., Aagaard P., Egeberg P.K. & S.P. Simmons 1995. Geochemical constraints from formation water analyses from the North Sea and the Gulf Coast Basins on quartz, feldspar and illite precipitation in reservoir rocks: in Cubitt, J.M. and England, W.A (eds), The Geochemistry of Reservoirs, Geol Soc. Spec.Pub1. No 86, p. 33-50. Boles, J.R. 1978. Active ankerite cementation in the subsurface Eocene of Southwest Texas. Contr. Mineral. Petrol. 68: 1322.

Figure 4. H2S mineral buffers expressed as H2S partial pressures versus temperature (isopleths). The unbroken lines correspond to pyrite-siderite buffer with siderite activity of 1, 0.1 and 0.0 1, while the dotted lines refer to pyrite-magnetite(a) and pyrite-kaolinite-chlorite(b).Data points are from Hiland et al. ( 1999) and Smith (1997).

and making the simplifying assumption that the Lewis fugacity rule applies. The convertion to partial pressure is then done by:

P(H2S) f (H2S) / T(H2S)

(4)

The H2S partial pressure mineral buffer lines are depicted in Figure 4 together with reported P(H2S) values from the Norwegian Shelf and US Gulf Coast. Despite some scatter, the H2S values group around the pyrite-siderite buffer, indicating that re131

Curtis C.D., Hughes C.R., Whiteman J.A. & C.K. Whittle 1985. Compositional variations whithin some sedimentary chlorites and some comments on their origin. Min. Mag. 49: 375-3 86. Dakhnova, M.V., Gurieva, S.M. & E.N. Shlutnik 1993. On the distribution of hydrogen sulphide in the carbonate oil and gas fields of the Russian Platfom: in Spencer, A.M. (ed) Generation, Accumulation and Production of Europe s Hydrocarbons III, Spec. Publ. EAPG no 3: 337-342. Egeberg, P.K. & P. Aagaard 1989. Origin and evolution of formation waters from oil fields on the Norwegian Shelf. Applied Geochemistry 4: 131- 142. Giggenbach, W.F. 1980. Geothermal gas equilibria. Geochim. Cosmochim. Acta 44: 2021-2032. Helgeson, H.C., Knox, A.M., Owens, C.E. & E.L. Shock 1993. Petroleum, oil field waters, and authigenic mineral assemblages: Are they in metastable equilibrium in hydrocarbon reservoirs?. Geochim. Cosmochim. Acta 57: 3295-3339. Hillier, S. & B. Velde 1991. Octahedral occupancy and the chemical composition of diagenetic (low-temperature) chlorites. Clay Miner. 26: 149-168. H%land, K., Barrufet, M.A., bnningsen, H.P., & K.K. Meisingset 1999. An empirical correlation between reservoir temperature and the concentration of hydrogen sulfide. SPE 50763: 1-8. Jahren, J.S. 1991. Evidence of Ostwald ripening related recrystallization of chlorites from reservoir rocks offshoe Norway. Clay Miner. 26: 169-178. Jahren, J.S. & P. Aagaard 1989. Compositional variations in diagenetic chlorites and illites, and relationships with formation-water chemistry. Clay Miner. 24: 157-170. Jahren, J.S. & P. Aagaard 1992. Diagenetic illite-chlorite assemblages in arenites. I. Chemical evolution. Clays Clay Min. 40: 540-546. Johnson, J.W., Oelkers, E.H., & H.C. Helgeson 1992. SUPCRT92: A software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species, and rectionss from 1 to 5000 bars and 0" to 1000°C: Computers and Geosciences. Milliken, K.L. 1998. Carbonate diagenesis in non-marine foreland sandstones at the western edge of the Alleghanian overthrust belt, southern Appalachians: in Morad, S. (ed), Spec.Pubs int Ass. Sediment. 26: 87-105. Morad, S., de Ros, L.F., Nystuen, J.P. & M. Bergan 1998. Carbonate diagenesis and porosity evolution in sheet-flood sandstones: evidence from the Middle and Lower Lunde Members (Triassic) in the Snorre Field, Norwegian North Sea: in Morad, S. (ed), Spec.Pubs int Ass. Sediment. 26: 5385. Orr, W.L. 1977. Geologic and geochemical controls on the distribution of hydrogen sulfide in natural gas: in Campos, R. & Goni, J. (eds), Advances in Organic Geochemistry 1975, Enadimsa, Madrid, 571-597. Reid, R.C., Prausnitz, J.M. & Poling, B.E. 1987. The properties of gases & liquids: McGraw-Hill. Reksten, K. 1990. Superstructures in calcite. Amer. Miner. 75: 807-8 12. Reksten, K. 1990. Superstructures in calcian akerites. Physics Chemistry Minerals 17: 266-270. Shock, E.L. 1994. Application of thermodynamic calculations to geochemical processes involving organic acids: in Pittman, E.D. & Lewan, M.D. (eds). Organic acids in geologicalprocesses, 270-3 18. Smith, J.T. 1997. Personal communication. Smith, J.T. & S.N. Ehrenberg 1989. Correlation of carbon dioxide abundance with temperature in clastic hydrocarbon reservoirs: relationship to inorganic chemical equilibrium. Marine and Petroleum Geology 6: 129-135.

Stoessel, R.K. 1984. Regular solution site-mixing model for chlorites. Clays Clay Min. 32: 205-212. Wade, W.J., Hanor, J.S., & R. Sassen 1989. Controls on H2S concentration and hydrocarbon destruction in the Eastern Smackover trend. Trans. Gulf Coast Assoc. Geol. Soc. 34: 309-320. Worden, R.H. & P.C. Smalley 1996. H2S-producing reactions in deep carbonate gas reservoirs: Khuff Formation, Abu Dhabi. Chem. Geology 133: 157-171.

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Quantifying recharge of the Ghussein wells using chemical tracers Nizar S .Abu-Jaber Department of Earth and Environmental Sciences, Yarmouk University, Irbid 21163, Jordan

ABSTRACT: A modified chloride mass balance approach is used to assess ground water recharge in the shallow aquifers of the Ghussein well field in northeastern Jordan. Chemical mass balance of other major solutes (carbonate, sulfate, magnesium, calcium, potassium and sodium) are taken into account as well as alternate sources of solutes other than wet precipitation. These phases and constraints are used as inputs for NETPATH simulations (Plummer et al. 1991). Results of this approach suggest that salts in the surficial deposits are a major contributor of chloride to the recharge water, and thus recharge values are higher than would be expected assuming traditional chloride mass balance assumptions. These results are confirmed by stable isotopic data for the water.

1 INTRODUCTION

The remoteness of the area precludes anthropogenic changes in the geochemistry of the water. The regional water table in the area is over 200m deep, and is considerably older than the waters in the shallow aquifers (Abu-Jaber et a1.1998).

1.1 Study area The Ghussein wells in arid northeastern Jordan (rainfall O are in the stability field for stilbite (Fig. 2c).

We propose that plagioclase dissolves by a reaction that produces stilbite and analcite (or albite or a Na-zeolite). The anorthite component of plagioclase is used up by stilbite formation and Ca in the water cannot further increase.

4 WATER-ROCK REACTIONS Plagioclase dissolution can be written for an average crustal feldspar composition (oligoclase): Na4CaAlgSi14040 + 6 H++ 3 H20 =$ 3 A12Si205(OH)4 + 8 Si02 + Ca+++ 4 Na+

(1)

Progress of reaction (1) increases Ca and Na in proportions controlled by the composition of plagioclase. TDS and pH increase as well. The problem with this formulation of the plagioclase reaction is that it does not specify the anion that is associated with Ca and Na. If the anion is formed from the components from the system defined by equation (1) the main candidate is OH(+complexes). Any H+ withdrawn from the water and invested into the release of Ca and Na must be compensated by the concurrent production of OH-. The dissolution reaction (1) will rapidly lead to very high pH-waters, which are not observed. Furthermore, increasing the pH from 6 in typical low-TDS meteoric water to pH=lO by reaction (1) will increase total Ca + Na concentration by a very small amount only. The water evolution path for the oligoclase hydrolysis reaction (1) is shown on Figure 3a for pure water arld two different initial low-TDS water compositions. Evidently, reaction (1) cannot explain the data array.

Reaction (2) progresses until also N a is taken up by either Na-smectite or Na-zeolite. From then on, plagioclase will be transformed directly into product zeolites according to the reactions for both plagioclase componennts (here written with natrolite as Na-sink): anorthite + 5 quartz + 7 H20

stilbite

(3)

2 albite + 2 H20 3 natrolite + 3 quartz

(4)

The relative cation composition of the water is not affected by the simultaneous transformation of albite and anorthite component of plagioclase into Ca-Na-zeolite. In the crystalline rocks of the continental crust plagioclase is a mineral present in “infinite” amounts. The plagioclase reservoir cannot be exhausted by reactions ( 3 ) and (4). However, water is available in limited amounts in the fracture network of the basement. Eleven moles of water are fixed into the solid reaction product per mole of oligoclase feldspar. TDS i n the residual water increases i n the process described by reactions (3) and (4).

150

Figure 2. a) Stability diagram at 10 bars, 25°C and qtz-saturation, b) 10 bars, 50°C and qtz-saturation, and c) 10 bars 25°C and SI,,,=l for mineral water (filled) and thermal (open) waters of the Black Forest basement. Laumontite has been found very widespred but also stilbite is a common zeolite in Alpine fissure deposits in granite gneiss (Wagner et al. 2000; Armbruster et al. 1995). Zeolites have been found on fractures in granitic and gneissic drill cores samples in major deep drilling projects (e.g. Borchardt & Emmermann, 1993, Moller et al. 1997). The reports demonstrate that zeolites form in granitoid basement rocks from low temperature interaction of water with matrix plagioclase. For C02-rich mineral waters the fluid evolution is essentially controlled by the reaction:

The reaction path of the plagioclase dissolution ( 5 ) is shown for an An20 plagioclase and for different values of log pco, corresponding to CO2 close to the atmospheric CO2 pressure and 10 % of CO2 saturation (Fig. 3b). The water data from the Black Forest area are contained within this range of CO2 pressures. The calcite saturation shown on Figure 3b for the different CO2 pressures indicate that some reaction path may be accompanied by calcite precipitation. Calcite precipitation will, however, not limit the plagioclase dissolution reaction. Production of secondary calcite is a side effect of the irreversible plagioclase dissolution reaction in the presence of C02. The seawater composition (Parkhurst et al. 1980) is shown on Figure 3b with the seawater dilution trend. The Black Forest data are well represented by the seawater dilution line. However, Stober and Bucher (1999b) showed

from the Cl/Br systematics that only the thermal waters contain a significant seawater component. Alternatively, the thermal waters can be represented by the line describing the equilibrium conditions of reaction (6): CaA12Si20g + 6 1 0 2 + 2Na+ = 2NaAlSi30g + Ca++

(6)

The equilibrium conditions of reaction (6) are shown on Figure 3b for plagioclase of x A n d . 2 . The exchange is independent of pH and the waters are well represented by the 50-100°C equilibria for oligoclase. If reaction (6) was efficient, it would tightly control the Ca/Na of natural waters because of the large amount and the restricted composition of plagioclase in basement rocks. On an average, mineral waters are more Ca-rich than thermal waters. Expressed as X A b , mineral waters are close to 0.75 and thermal waters average at 0.93 on an activity basis. Hot basement water is sodium rich and in cold basement water calcium dominates as predicted by the temperature dependence of equilibrium (6). ACKNOWLEDGMENTS The generous supply of unpublished thermodynamic data of Ca-zeolites by Ch. de Capitani (Basel) is gratefully acknowledged. J. Liebermann and E. Perkins provided software for the computation of the phase diagrams.

151

Figure 3. a) Reaction path of plagioclase hydrolysis rxn ( 1): b).Three different processes: seawater dilution (A); plagioclase exchange equilibrium (6) (B); reaction path of plagioclase hydrolysis rxn 5 (C). Data of mineral water (filled) and thermal (open) waters of the Black Forest basement.

REFERENCES Armbruster, Th., Kohler, Th., Meisel, Th., Nagler, Th.F., Gotzinger, M.A. & H.A. Stalder 1996. The zeolite, fluorite, quartz assemblage of the fissure at Gibelsbach, Fiesch (Valais, Switzerland): crystal chemistry, REE patterns, and genetic speculations: Schweiz. Min. Pet. Mitt. 76: 131-146. Berman, R.G. 1988. Internally-Consistent Thermodynamic Data for Minerals in the System: Na20- K 2 0 - CaO- MgO- FeOFe203- A1203- Si02- Ti02- H20- C02: J. Pet. 29:445-522. Borchardt, R. & R. Emmermann 1993. Vein minerals in KTB rocks: KTB Report 2:481488. Bucher, K. & I. Stober 2000. Hydrochemistry of water in the crystalline basement. In: Stober, I. & Bucher, K. (eds.); Hydrochemistry of water in the crystalline basement. Dordrecht: Kluwer Academic Publishers, 141-175. Helgeson, H.C., Kirkham, D.H. & G.C. Flowers Theoretical prediction of the 198 1. thermodynamic be h a v i o ii r of a q U e o i t s

electrolytes at high pressures and temperatures. IV. Calculation of activity coefficients, osmotic coefficients, and apparent molal and standard and relative partial molal properties to 600°C and 5 kb: Am. J. Sci. 281:1249-1516. Moller, P. 1997. Paleo- and recent fluids i n the upper continental crust - Results froin the German Continental deep drilling Program (KTB): J. geoph. Res. 102: 18245-18256. Parkhurst, D.L., Thorstenson, D.C. & L.N. Plummer 1980. PHREEQE - a computer program for geochemical calculations: U.S. G eol o g i c a 1 S ur v e y , W ate r Re s o i i rc e s Investigations, 80-96:2 10 pp. Stober I. & K. Bucher 1999a. Deep groundwater in the crystalline basement of the Black Forest region. Appl. Geoclz. 14: 237-254. Stober, I. & K. Bucher 1999b. Origin of Salinity of Deep Groundwater in Crystalline Rocks: Terra Nova 11:181-185. Wagner, A., Stalder, H.A., Stuker, P. 8: E. Offermann 2000. Arvigo - eine der bekanntesten Mineralfundstellen der Schweiz: Schweizer Strahler, 12:41-58 & 118-133.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Seawater-basalt interaction: field observations and modeling result 0.V.Chudaev Far East Geological Institute, Vladivostok Russia

V.B .Kurnosov Geological Institute, Moscow, Russia

0.V.Avchenko Far East Geological Institute, Vladivostok Russia

N.A.Chepkaya Far East Geological Institute, Vladivostok, Russia

ABSTRACT: Mineralogy and geochemistry of altered basalts from aseismic ridge structures (legs 26, 62, 72, 74, 115, 120 and 121), mid-oceanic ridges (legs 33,37,52 & 53, 91, 104, 106, 111, and 123), as well as from West Pacific trenches were studied. This study confirms the suggestion that basalt alteration occurred during oceanic floor spreading and that smectite is a widespread mineral. Using SELECTOR geochemical modeling software the authors attempted to reconstruct seawaterbasalt interaction in seafloor hydrothermal systems. 1INTRODUCTION Studies of geochemistry of hydrothermal systems in the ocean began over than 20 years ago. The main results in this field were summarized by Von Damm (1995). Experimental studies of seawaterbasalt interaction were carried out by Bischoff and Dickson (1975), Seyfried (1987), Bishoff (1991) and many others. These studies suggest that most of the chemical elements, during seawaterbasalt interaction, are mobile and their mobility is connected with seawater temperature and pressure, waterhock ratio, composition of dissolved gases, residence time, etc. Von Damm (1995) assumed that, during rock alteration, C1, B, Na, Mg, are removed from seawater, while Si, K, Ca, S, Fe, Mn, Li, Sr, Cu, leach from the wall rocks. One of the tools for understanding this process is geochemical modeling software. Reed (1983) calculated seawater-basalt reactions at 300°C and 500 bar and concluded that chlorite, hematite, pyrite, quartz, kaolinite, epidote, paragonite, magnetite, sphalerite, phlogopite, albite, tremolite, and chalcopyrite formed in these conditions. Using chemical data of fluids from EPR 21"N and Guayamas basin Bowers et al. (1985) calculated saturation state for the seafloor hydrothermal fluid. Mixing of the end member fluids with seawater provided mineral assemblages closely approximating those observed in the samples of chimneys from the East Pacific Rise and Guayamas Basin. In spite of progress in studies of the seafloor hydrothermal systems, some aspects of seawaterbasalt interaction are still

uncertain or debatable. For example, the influence of volcanic gases on fluid composition and P, T stability of smectite. We can suggest that C02, HC1, H2S, H2 are predominant in the gas phase. We know exactly that volcanic gases not only input some chemical elements in fluids, but also change their chemical characteristics (pH, Eh and etc). It is known that fluids discharged in hot springs derived from seawater penetrating into oceanic crust and interacting with wall rocks. However, the reactions that occur in the recharge zone are still uncertain. It is agreed that smectite is a low temperature mineral, but one found in high temperature condition (up to 380 "C) in some hot springs. In this paper, basing on mineralogy of altered basalts from West Pacific trenches, aseismic ridge structures (legs 26, 62, 72, 74, 115, 120 and 121), mid-oceanic ridges (legs 33, 37, 52 & 53, 91, 104, 106, 111, and 123), and on the results of SELECTOR geochemical modeling software (Karpov et a1 1997), we attempted to reconstruct the physico-chemical conditions of seawaterbasalt interaction.

2 METHODS Secondary minerals from basalts were studied using traditional methods: X-ray diffraction, electron transmission microscopy, and EPMA analyses. Geochemical modeling of seawater/basalt/gas interaction consists of two main runs, but in the beginning we composed the seawater. The total

153

dissolved solid in water was 35.17'/,,. Some amounts of H and 0 2 were added to balance the system composition. The model of seawater was constructed under open system (with air) and closed system (without air) in conditions T=25'C and P = l bar. In the first run, seawater was calculated at different temperatures and pressures. Some amounts of CO2, H2S, HCl were added at T=350°C and P=350 bar. In the second run, seawater interacted with basalt at different waterhock ratios, temperatures, pressures and amounts of volcanic gases (Fig. 1). Modeling Seawater T=25"C, P=l bar

Run#2

/

I

Seawatedbasalt I (1O:l) T=125"C, P=200bar Seawatedbasalt (1 0:2) T=225'CC,P=250bar

\Rm#1

I

Seawater I F=125"C, P='IOObar/ Seawater T=225"C, P=250bar

throughout Galapagos spreading center (hole 504B). In pillow lavas the minerals of smectite group, as well as calcite and quartz are widespread. The minerals of green schist facies appear in dike complex. In mid-oceanic ridges smectite formation temperature ("C) ranges from tens to first hundreds degrees Celsius. In aseismic zones, basalts are less altered. Sporadic dioctahedral smectite, Fe-oxide and calcite occur in the surface part of pillow lavas, while trioctahedral smectite with minor ammount of illite, chlorite, zeolites, quartz, and sulfides in their inner part. Smectite formation temperatures range from 10 to 120°C (Kurnosov at al. 1997). In altered basalt from the West Pacific dioctahedral smectite is predominant in the ground mass, while. veins and fractures are filled with trioctahedral smectite, mixed-layered minerals of the smectite-chlorite group, as well as calcite, barite, zeolites, and apatite. Isotopic studies revealed 90°C as the upper formation temperature for smectite (Chudaev 1995). Thus, smectite is well-developed in secondary minerals of altered oceanic basalts. Its formation temperature ranges from first tens to first hundreds of OC that proves by computer modeling. Table 1 lists the minerals being formed in the course of basal-seawater system modeling. Smectite was stable up to 325"C, its amount depending on waterhock ratio. Increase of this ratio lead to the decrease of smectite formation. Gypsum appeared at 125OC, than dropped at 225OC, and disappeared at 325OC. Cooling of the system up to 25OC caused a large amount of gypsum to precipitate Table 1. Precipitated minerals during modeling of seawaterbasalt interaction. Minerals, %

r

9

$

' '3

Figure 1. Process modeling scheme.

125OC 200b 225'C 250b 325°C 300b

Sm

Gy

D1

Cc

Amf

Mt

Pr

WIT

80 86 84 91 91 95

13 5 6 0 0 0

0 0 0 0 0 0

0

7 7 7 7 7 0

0 3 2 2 5

0 0 0 0 0 0

100 10 50 5 33 3.3

0

0

0 0 0 0 0

2

3 RESULT AND DISCUSSION

-

3.1 Mineralogy

Alteration of basalts in mid-oceanic ridges is not uniform. A typical secondary mineral of altered basalts is smectite that fills vesicles and fractures. In the areas of high-temperature hydrothermal vents chlorite, epidote, smectite, and sulfides are predominant. Mineralogical zoning is well detected

Y

9

250c

34

65

0

0

0

0

1

100

200b

65

27

5

0

0

0

3

10

ti

Note: Sm-smectite; Gy-gypsum; Amf-amphibole; Mt-magnetite; Pr-pyrite. w/r-water/rock ratio (100, 50.. mean 100, 50.. parts of seawater and 1 part of basalt).

154

some chemical elements depends on water(fluid)basalt ratio, temperature and the amount of gas added to the fluid (Table 3).

(Table 1). Amphibole (tremolite-actinolite) forms (about 7%) at 125OC and remains stable at 325"C, while it disappears with hydrothermal solution cooling (Table 1). The amount of actinolite does not depend on water-rock ratios. Calcite and dolomite mainly participate when gas (C02, HC1, H2S) was added to the solution. Minor amounts of dolomite and quartz formed when the system was cooled. A small amount of magnetite and ilmenite formed in most runs, disappearing in the solution-gas system. Pentlandite appeared at 325OC and remained stable in the same amount (1%)after cooling. Precipitation of pyrite started only after cooling of the hydrothermal fluid.

Table 3. Contents of some chemical components mmol) in seawaterbasalt modeling (run ## 2, )

Initial sea water

-3

h 2 ti

3.2. Seawater composition

200b 225OC 250b 325'C 300b

pH

Na

K

Ca

Mg

C1

SO4 w/r

7.88

483

10

10

53

565

0.3

10.6 11.7 9.9 10.5 9.4 5.4

455 514 464 684 505 909

10 223 11 209 10 60 13 696 11 132 17 1511

54 127 65 320 98 733

520 510 506 572 538 599

18 2.8 13 4.8 14 4.9

991

20

54

3

1129

900

18

22

1.5

960

462 828

11 102 15 3.4

68 40

102 770

100 10 50 5 33 3.3

v1

2

3

Concentrations (mmol) of some main chemical components are shown in Table 2.

ti

Table 2. Contents of some chemical components (mmol) in seawater modeling (run ## 1,Figure 1) Initial sea water

200b 225'C T 3 250b 325k 300b

2

pH

Na

K

Ca

Mg

C1'

SO4

7.88

483

10

10

53

565

0.3

6.7

448

9.5

8.2

47

520

26

6.6

433

9.1

8.2

39

498

26

6.8

401

8

7

27

450

25

Seawater starts as a fluid containing C1 >>SO4 >> alkalinity>>Br>B>F. Ion concentration changes during runs. Only heating of seawater causes pH and chemical element contents to change. PH slowly decreases during runs. Adding a small amount of gas (COz, HC1, H2S) to the fluid at 35OoC and 350bar shifts pH to 5.2, and after fluid cooling pH drops to 2.7 (Table 2). Chlorine slowly decreases during fluid evolution and only after fluid cooling it shifts back close to original values (535 mmol). SO? decreases in runs with gas due to some gypsum precipitation. Concentrations of Mg2', K', Na', and Ca2+slowly decrease with increasing temperature (Table 2). When seawater interacts with basalt the behavior of

% ti

350°C 350b

25OC 200b

5.2

2.8 4.8

10 0.7 27

100 10

Comparison of the initial seawater with evolving fluid shows that pH increases at T=325'C, P=300 bar and reaches 9.4 (fluidbasalt = 33) and 5.4 (fluidbasalt = 3.3). When gas is added, pH shifts to 5.2. When the system is cooled, pH drops up to 2.8 (fluidbasalt = 100) and 4.8 (fluidbasalt = 10). C1 content does not change up to 325"C, increasing up to 1129 mmol when gas (including HCl) is added to the reservoir. After system cooling, C1concentration drops to 102 mmol (fluidbasalt ratio = 100) and increases up to 770 mmol when fluidbasalt ratio is about 10 (Table 3). SO,"- content increases during seawaterbasalt interaction, while after decreases (fluidbasalt ratio = 100). cooling SO-: During runs the content of Na' does not change much when fluidbasalt ratio is high, and increases with increasing temperature when this ratio is low. The Na' and C1- contents correlate well. K' concentration is rather stable. Its content increases during fluidbasalt interaction after adding gas to fluid at low fluidbasalt ratio (Table 3). Ca2' concentration of is higher in fluids than in the initial seawater. Its content depends on temperature and fluidbasalt ratio, as well as Ca-bearing minerals precipitation. Magnesium increases with increasing temperature, and fluidbasalt ratio and its contents drop at 3OO0C, P=200 bar. At that temperature and pressure dolomite and calcite precipitate.

155

REFERENCES

In the proposed model, smectite prevailed and remained stable until temperature dropped 325°C. At the same, time under high P and T parameters, many minerals of green-schist facies (epidote, chlorite, and others) were not obtained. This might have been due to kinetic reasons. Redistribution of chemical elements and precipitation-dissolution of minerals in the model are controlled by T, P, fluid\basalt ratio, which was proven by both experiments and field observations (Bishoff & Dickson 1975, Seyfried 1987, Von Damm 1995). Experimental results, however, revealed that in basalt\seawater interaction Mg precipitated as Mg-OH-Si mineral to form acidic pH. In the model we proposed, acidic pH formed only when gas (C02, HCl, H2S) was added. It, probably agrees better with natural processes. Improvement of the model depends very much on availability of a trustworthy thermodynamic database for smectite, characterized by variable chemical composition and development of its layerstructure. The calculation results are probably, more suitable for recharge zones of the seafloor hydrothermal system. In this model we could calculate of separate process occurring at high temperature and pressure, due to the poor thermodynamic data on the solution near critical fluids, but probably this process plays significant role in geochemistry of hydrothermal fluids (Bishoff 1991).

Bishoff, J.1991. Density of liquids and vapors in boiling NaC1H20 solutions:A PVTX summary from 300-500°C. Am. J. Sci. 291, 309-338. Bishoff, J. & Dickson, F. 1975. Seawater-basalt interaction at 200°C and 500bars: Implication for origin of seafloor heavy metal deposits and regulation of seawater chemistry. Earth Planet. Sci. Letters. 25. 385-397. Bowers, T., Von Damm, K., M., Edmond. 1985. Chemical evolution of mid ocean ridge hot springs. Geochim. Cosmochim. Acta. 49.2239-2252. Chudaev, O., I., Tararin. 1989. Hydrothermal metamorphism in deep-sea trenches of the Western Pacific.. Proceedings of Water-Rock Interaction. Balkema. Rotterdam. 159-162. Chudaev, 0. 1995. The budget of chemical elements in Pacific ocean. Guiots of the Western Pacific and their mineralization. Moscow. Nauka. 326-335. Karpov, I., Chudnenko, K., D., Kulik. 1997. Modeling chemical mass-transfer in geochemical process: Thermodynamic relations, conditions of equilibria and numerical algorithms. Amer. J. Sci. 287. 1-39 Kurnosov, V. 1986. Hydrothermal alterations of basalts in Pacifc ocean and ore deposits. Moskow. Nauka. 251p. Kurnosov, V., Zolotarev B., A., Artamonov. 1997. Alteration Effects in the oceanic crust. Scientific report. JOI/NERC Russian Scientists Support Program. 136p. Reed, M. 1983. Seawater-basalt reaction and origin of Greenstones and related ore deposits. Economic Geology. 78.466-485. Seyfried, W. 1987. Experimental and theoretical constrains on hydrothermal alteration process at mid-ocean ridges. Ann. Rev. Earth Planet. Sci. 15. 317-335. Von Damm K. 1995. Control on the chemistry and temporal variability of seafloor hydrothermal fluids. 1995. Geophysical monograph 91.222-247. Von Damm, K. 1995. Temporal and compositional diversity in seafloor hydrothermal fluids. Reviews of geophysics supplement. Paper number 95RG00283.

4 CONCLUSIONS 1. The hydrothermal alteration of basalts from aseismic ridge structures, mid-oceanic ridges, as well as basalts from the West Pacific trenches confirms suggestions that alteration in the upper part of oceanic crust occurs during the spreading of the oceanic floor and smectite is a widespread mineral. Its formation temperature ranges from first tens to first hundreds degrees C. 2. During most of the runs of seawater (fluid)basalt interaction, smectite was a significant mineral of the calculations. It remained stable up to 325OC, its amount depending on waterhock ratio. 3 Redistribution of chemical elements depending on T, P, fluid basalt ratios, and mineral precipitationdissolution is clearly observed by modeling results.

ACKNOWLEDGEMENTS The financial support from the Russian Foundation for Basic Research (Project 99-05-64487) is hereby acknowledged 156

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Arsenic sulphide precipitation in an active geothermal system: reaction path modelling James S .Cleverley & Liane G.Benning University of Leeds, Woodhouse Lane, Leeds, Ls2 9JT, UK Bruce W.Mountain Wairaki Research Center, Institute for Geological and Nuclear Science, Private Bag 2000, Taupo, New Zealand

Margaret C .Gorringe University of Leeds, Woodhouse Lane, Leeds, Ls2 9JT, U K

ABSTRACT: A sampling profile across the Eastern hydrothermal field in the Uzon Caldera (Kamchatka, Russia) revealed a strong zonal character in the As mineralisation that coincides with changes in both temperature and redox conditions. Four fluid samples, collected along a profile, were used as a basis for geochemical speciation and reaction path modelling to predict the observed mineral zonation. Modelling the changes in concentration of dissolved As between the samples (0.2-8.6 ppm) indicates a strong dependence on redox (log f 0 2 ( ~ ) from -53 to -60) and temperature (95-65'C), whilst pH (2.7-6.2) exhibited little control. A reasonable representation of the observed mineral zonation was modelled using cooling of As-rich (10 ppm) fluids with a fixed pH (5.5) and redox (log f 0 2 ( g ) = -58) from 125 to 25'C. This work has shown that despite assumptions in fluid parameters and modelling approaches, acceptable and useful analogues for natural systenls can be developed. 1 INTRODUCTION

nic concentrations in the fluids range between 0.2 to 8.6 ppm and are highest at the South-Western end of the profile where the temperature at the surface reaches 80°C (U24 -Table 1). The mineralisation is localised at depths between 0.05 to l m and the deposits have a strong zonal character, which correspond to sharp redox and temperature changes. From top to bottom, the following sulphide layers were observed: sulphur + amorphous A s ~ S->~ orpiment (AS&) -> orpiment + realgar (ASS) -> realgar -> realgar + pyrite (Fe&) -> pyrite (e.g., Karpov & Pavlov 1982, Benning & Mountain 1996, Migdisov & Bychkov 1998). Locally, uzonite ( A s ~ S ~alacranite ), (AsgS9) and stibnite (Sb2S3) are found. Migdisov & Bychkov (1998) have shown that these sharp changes can also be correlated with progressive shifts in sulphur chemistry and the formation of a complex set of sulphur ligands (e.g. H2S(aq),S*O?-, SO?-, SO-: and S'(aq)). Here we present the results of a speciation and reactive path modelling study, in which the chemical characteristics of four representative fluids sampled from the active arsenic-antimony sulphide forming environment in the Eastern geothermal field were used as input data. Note that, difficulties in modelling the metal speciation and precipitation reaction paths arise from the lack of thermodynamic data for most of the sulphur species mentioned above as well as the shortage of data for many aqueous arsenic species. In addition, the situation is complicated by the fact that in such a system, particularly at lower

The deposition of metal sulphide phases in active ore-forming environments is strongly controlled by the metal speciation in the hydrothermal fluids as well as by the solubility of the precipitating sulphides. Rapid changes in pH, temperature, and redox potential, greatly affect the solubility and speciation and thus the character of an ore-deposit. The active deposition of arsenic and antimony sulphide minerals occurring in the Uzon Caldera (Kamchatka, Russia) provides a unique opportunity to study and model the relationship between metal sulphide precipitation and extreme shifts in fluid parameters. The geochemistry of the mineralisation was studied along a 42-meter profile over the Eastern geothermal field. Fluid and solid samples were collected along the profile in 2-meter intervals (Benning & Mountain 1996). The temperature, pH and total reduced sulphide for each fluid sample were measured on site. The sampled fluids were acidified and their composition was determined using a combination of ICP-MS for the cations and IC for the anions (Table 1). Note, however, that although the measured pH and total reduced sulphur determinations are average values over the depth of each hole, these values agree well with the detailed potentiometric measurements reported by Migdisov & Bychkov (1998) over a crosscutting profile. The fluids are NaCl brines containing significant concentrations of K, Ca and SO:-, B, As and Fe. Arse157

temperatures, kinetic effects will control the reactions. However, under certain assumptions (discussed below) the reaction path modelling shows that the precipitation patterns observed in the field can be reproduced fairly accurately. 1.1 Arsenic and sulphur species A compilation of literature data for aqueous As species and arsenic sulphide minerals was made and in all cases (except amorphous AS&) literature data included or provided the raw data to calculate AGO, AH’, So, V” and the Maier-Kelly Cp fitting coefficients. This data was added to a database (GEOPIG, 1998) used by SUPCRT92 (Johnson et al. 1992). Equilibrium constants for the new or modified As species with H2AsOi as the base species were calculated and the resulting values were subsequently used as input for the Geochemists Workbench database (GWB, Bethke 1998, LLNL, 1996). Aqueous species added or modified in the GWB database include A~zS3(aq),HAs2S4- and As2S;(from compilations in Zotov et al. 1994). The mineral data for claudetite (As203, mon), arsenolite (As2O3, cub), orpiment (AS&) and realgar (ASS) are from Pokrovski et al. (1996). The experimental log K data of Eary (1992) for following the reaction:

For H3As03(,,), in this study the experimental data of Zakaznova-Iaklovleva at al. (2000) was used, and not the data from the SUPCRT92 database. As a consequence, orpiment is more stable than A S ~ S ~ ( ~ and to overcome this, the stability field for AS&,,,) was calculated by suppressing the precipitation of orpiment and realgar in some runs (see below and Figure 1 & 2). For this modelling study the whole array of sulphur species reported by Migdisov & Bychkov (1998) were neglected because no thermodynamic data were available. Note also that the HzS(,,) ionisation constants in the SUPCRT92 database deviate substantially from the new spectrophotometric determinations of Suleimenov & Seward (1997). However, at the temperatures discussed below (25 - 125°C) the effect is small and can be neglected. 1.2 Estimation of redox conditions The redox potential of the fluids was estimated, using the equilibrium between sulphate and total reduced sulphur;

+ 202(g) = SO42-(aq)+ 2 ~ +

~2s(aq)

was fitted to a temperature dependent polynomial equation to allow prediction of equilibrium constants at the temperatures used in the GWB database. Table 1. Key data’ from the different sample trenches used in the studv. All data in molal unless sDecified. NE sw UZ15 UZll UZ4 Data UZ18 80 95 65 T (“(3 68 5.4 2.7 5.3 6.2 PH

cs reduced

2.0Xio-4 4 . 3 ~ 1 0 ~2 . 8 ~ 1 0 . ~1.1xlo4 SO-: (ppm) 79.3 115.0 76.9 98.4 0.24 8.6 (ppm) 0.88 2.0 CAstaq) CSO;-(,) 1.03~10” 3.78~10” l . l ~ l O - ~1.14~10” Logf02tg) -58.2 -54.1 -59.0 -53.1 Log f H Z ( g ) -6.3 -5.0 -6.1 -7.2 SO?-(,,) 7 . 8 1~0-4 1 0 . 910-4 ~ 7 . 4 1~0-4 6 . 51~0-4 HSO . ~ X ~ O 1- .~6 ~ 1 0 - ~0 . 2 ~ 1 0 ~1.5x10-* H2S (a,) 1 . 9 ~ 1 0 - ~2 . 5 ~ 1 0 ~2 . 7 ~ 1 0 - ~0 . 8 ~ 1 0 - ~ AS&(,,) 5.1xlO-’ 5.4x10-* 1 . 7 ~ 1 0 - ~5 . 4 ~ 1 0 - ~ HAsO2(aq) 3 . 7 ~ 1 0 - ~9 . 7 ~ 1 0 . ~8.9x10-’ 4 . 2 ~ 1 0 - ~ *Note: Rows 1 to 5 are field or laboratory measurements. CSO?. is calculated CS, as sulphate. Rows 7 to 13 are equilibrium speciation distribution of major S and As species in the fluids. The most predominant species are tabulated only.

158

(2)

where the concentration of H&,), S04?- and pH represent measured values. Note, that the estimated redox values for all fluids (Table 1) fall within the range of the potentiometric measurements reported by Migdisov & Bvychkov (1998). In all fluids, was assumed to be the most dominant reduced sulphur species at the measured pH (2.7-6.2). The value for sulphate (Table 1) may include partially oxidised portions of the reduced sulphur species. Therefore, the total sulphur data are maximum values and the calculated redox potential (log f 0 2 ( g ) range: -59 to -53) represents an upper limit. However, a variation of l log unit in the measurements for sulphide andor sulphate, changes the log f 0 ~ ( ~ ) by a maximum 0.5 log units, thus showing that the calculated redox range represents a reasonable estimate. 2 MODELLING RESULTS

2.1 Equilibrium speciation The fluid chemistry was speciated with the ‘REACT’ shell of GWB in a two-stage process. First the redox and the Total Dissolved Solids (TDS), as well as the CS(,,, in the system were estimated and subsequently used as input for the second stage speciation run. The metastable, rather than the full, equilibrium output results were used because precipitation of the supersaturated phases such as quartz and diaspore

Under these conditions, it can be shown (Figures 1 and 2) that in terms of thermodynamic equilibrium, realgar and orpiment are the stable phases while amorphous As2S3 is metastable. However, in order the to be able to plot the boundaries for precipitation for orpiment and realgar were suppressed and pH and log fH?-(g) were varied while the activities of H2S(aq) and A~2S3(aq) were fixed (Figure 2). The log aHlS(aq)(-3.7) in the diagram lies within the range of the average values reported in the fluid analysis (Table 1). However, the log U A S ~ S is ~(~~) significantly increased to enable the calculation of stable solid phase boundaries that were found to be undersaturated in the natural fluids (figure 1 and 2, grey box). At 70°C and log aAS2S3(aq) of -3.7 the concentration of total arsenic is approximately equal to 61 ppm.

are unlikely to occur at the temperatures of interest (25-125°C). Phase diagrams for the system As-S-H20 illustrate the critical phase boundaries for the dominant sulphur and arsenic species. The position of the box in figures 1 and 2 indicates the range of compositions for the fluids from the Uzon Caldera.

2.2 Reaction path modelling

Figure 1. Diagram to illustrate the phase relationships in the system As-S-HzO for variable activities of ASZS3(aq)and H2S(aq) at fixed redox, pH and temperature conditions (approximating UZlS). Solid lines separate arsenic minerals and aqueous species. The dotted line is the approximate position of the metastable boundary (see text for details). pH = 5.5, log fH2(g)= -6 and T = 70°C.

The GWB shell m A C T ' was further used for both forward and reverse reaction path modelling of the fluid compositions listed in Table 1. In this study we attempted to: a) investigate the relationship between the high CA~(,,,urn fluid and the low CA~(,,, fluids at "11 and and b, try to investigate the pattern Of As-S precipitation with depth.

2.3 Predicting the concentration of As in solution

The calculations were completed using the 'ACT2' shell of GWB and the output data from the speciation modelling. Figure 1 shows the relationship between solid and aqueous phases in the system As-SH2O in terms of the activity of AS2S3(aq) and with T, pH and log H2(g)fixed.

Two reaction paths (a and b) were modelled using the fluid chemistry from UZ4 to try to resolve the relationship of the variation of aqueous As between the samples.

Figure 2. Redox-pH phase relationships in the As-S-H20 system. Solid lines separate As minerals and species, dashed lines separate aqueous sulphur species. The dotted line marks the approximate metastable AS~S~(~,,,) boundary, while the dasheddotted line marks the stability field for pyrite. Note: the AS(aq) species are not shown where the solid phases are stable. T = 70"C, log a A s ~ S ~ (-3.7, ~ ~ )log = a H & , ) = -3.7, log a Fe2+= -5.

Figure 3. Prediction of changes in CAS(,~)in solution and the stability of Fe/As solid phases by two different reaction path models for the UZ4 fluid. pH changes from -2.7 to -3.4. Solid line = model a, dashed line = model b. (See text for details).

In model (a) change in redox from log f 0 ~ =( -53 ~) to -60 (log fH2(g)= -3.8 to -7.3) with freely variable pH and fixed temperature (80°C) was tested. Conversely, in (b) the model was run with temperature changing from 80 down to 60°C. These values were

159

has been modelled using the chemical information derived from solid and fluid samples in the Uzon Caldera (Kamchatka, Russia). Despite the dearth of thermodynamic data for many aqueous arsenic and sulphur species, that could play important roles in such systems, it was possible to reproduce, to fairly high accuracy, the distribution and precipitation sequence observed in the field (Figure 4). However, it must be considered that with all reaction path models, the results are based on equilibrium thermodynamic relationships between the components in the system. Shortfalls in available data will have a bearing on the results, although overall the model system represents a good estimate of the natural geothermal system at Uzon Caldera.

chosen to overlap with the ranges observed in the field. The actual As concentrations and the estimated log for the fluids from the Uzon Caldera are shown on Figure 3 as black circles. For the first model (a) the predicted As(aq)closely matches that measured for UZ15 (95'C). In model b, where temperature is linearly decreased to 6OoC, the model closely predicts the total As(aq)concentration in both UZ11 (65'C) and UZ18 (68°C). The discrepancy in pH (-2 log units) between the model (3.4) and the actual fluids (5.3, UZlS) does not appear to effect the reaction paths. 2.4 Modelling the mineral zonation Reaction path modelling was also used to investigate the mineral zonation patterns observed with depth. The zonation, pyrite - realgar - orpiment - A~2S3(am), broadly follows decreasing fluid temperature. In order to simulate the complete cooling of a fluid from high temperature, the fluid chemistry of UZ18 was reverse modelled to higher temperatures without allowing precipitation of mineral phases. In addition, the AS(aq)concentration was increased to 10 mg / kg (a reasonable equivalent to the concentration in UZ4). The reaction path modelling of UZI 8 was done by linearly decreasing temperature (120 - 25OC) while pH (rock buffered) and log fH2(g) were fixed (Figure 4).

Figure 4: Reaction path model for the cooling of As-rich fluid from 125 to 25OC with fixed pH and redox. pH = 5.3, log f02(g) = -58, CAs = 10 mg / kg, CFe = 2 mg / kg. Note: That the Y-axis records the concentration of aqueous species in the fluid and the amount of precipitate in equilibrium with the fluid. This model does not predict the occurrence of because of thermodynamic metastability (see section 1.1)

3 CONCLUSIONS The distribution and relative importance of aqueous arsenic species and the precipitation paths of arsenic sulphide mineral, in an active geothennal system,

REFERENCES Benning, L.G. & B. W. Mountain 1996. Metal distribution in modern arsenic mineralization associated with a hot spring envirnment: Uzon Caldera, Kamchatka, Russia. Geochemistry of Crustal Fluids. ESF. Austria, December 6-1 1. Bethke, C.M. 1998. Geochemists Workbench version 3.0. University of Illinois. Eary, L.E. 1992. The solubility of amorphous AS& from 25 to 90°C. Geochmica et Cosmochimica Acta. 56:2267-2280. GEOPIG 1998. Slop98.dat, htt~://7.onvark.wustl.edu/peoDie/, Washington University. Johnson, J.W., Oelkers, E.H. & H.C. Helgeson 1992. SUPCRT92: A software package for calculating the standard molal thermodynmic properties of minerals, gases, aqueous species and reactions from 1 to 5000 bars and 0" to 1000°C. Computers and Geoscience. 18:899-947. Karpov, G.A. & A.L. Pavlov 1982. Zoning of mineral deposits in the discharge areas of recent hydrotherms. In S.I. Naboko (ed.), Hydrothermal mineral forming solutions in the areas of active volcanism: 233-237. New Delhi: 0x0nian Press. LLNL, 1996. Thermo96.dat, ftp://s 122.es.llnl.pov, Lawrence Livermore National Laboratory Migdisov, A.A. & A.Y. Bychkov 1998. The behaviour of metals and sulphur during the formation of hydrothermal mercury-antimony-arsenic mineralization, Uzon caldera, Kamchatka, Russia. Journal Of Volcanology And Geothermal Research. 84:153-171. Pokrovski, G., Gout, R., Schott, J., Zotov, A. & J. Harrichoury 1996. Thermodynamic properties and stochiometry of As(II1) hydroxide complexes at hydrothermal conditions. Geochimica et Cosmochimica Acta. 60:737-749 Suleimenov, O.M. & T.M. Seward 1997. A spectrophotometric study of hydrogen sulphide ionisation in aqueous solutions to 350°C. Geochimica et Cosmochimica Acta. 61 :51875 198 Zakaznova-Iakovleva, V.P., Seward, T.M. & O.M. Suleimenov 2000. Spectrophotometric determination of the first ionisation constant of H3As03 from 25 to 300°C. In P.R. Tremaine, P.G. Hill, D.E. Irish & P.V. Balakrishnan (ed.), Hydrothermal systems: Physics and chemistry meeting the needs of industry: 694-695. Ottawa. Zotov, A.L., Kudrin, A.V., Levin, K.A., Shikina, N.D. & L.N. Var'yash 1994. Experimental studies of the solubility and complexing of selected ore elements (Au, Ag, Cu, MO, As, Sb, Hg) in aqueous solutions. In K.I. Shrnulovich, B.W.D. Yardley & G.G. Gonchar (ed.) Fluids in the crust. London:Chapman & Hall

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Mineral growth in rocks: interacting stress and lanetics in vein growth, replacement, and water-rock interaction R .C.Fletcher Department of Geological Sciences, University of Colorado, Boulder CO 80309-0399, USA

E .Merino Department of Geological Sciences, Indiana University, Bloomington IN 47405, USA

ABSTRACT: Growth of crystals in rocks generates a local stress, or force of crystallization, which may drive pressure solution, deformation, or cracking of the host rock. Several novel feedbacks involving this local stress may account for replacement, certain deformation textures, earthquake triggering, and ore-body selflocalization.Because the force of crystallization may modify (a) the rates of local phenomena such as mineral dissolution, mineral growth, andor rock deformation, and (b) transport parameters such as permeability, porosity, and rock viscosity, and because it is itself modified by those phenomena and parameters in several potential feedbacks, models of water-rock interaction should be extended in particular cases to incorporate both the force of crystallization and its relevant kinetic consequences and feedbacks. We find equations giving the interrelated force of crystallization and growth rate for the widespread cases of replacement and vein growth. Driven by supersaturation,veins may grow not as cement in previously open fractures but by pushing the host rock apart as they grow, and promoting crack propagation.

of Bruton and Helgeson (1983), who studied the effects of fluid pressures different from lithostatic on water-mineral equilibria, and of Dahlen ( 1992), who clarified the relation between macroscopic and microscopic (e.g., non-local and local) stresses and their relation to water-mineral equilibria. In short, because it is local, the growth-driven stress is able to enter into feedback with its own local consequences, kinetic or rheological. This is not possible for non-local tectonic or gravitational stresses, which cannot be modified by their local consequences. Our purpose here is (a) to suggest new possible feedbacks involving the force of crystallization that may help explain the elusive problems of volume preservation in replacement, earthquake triggering, and ore body self-localization, and (b) to calculate the force of crystallization and model its kinetic and rheoiogical consequences in two common occurrences of mineral growth in rocks, replacement and vein formation (Fletcher & Merino, 2000).

1 INTRODUCTION Growth of crystals or crystal aggregates in rocks necessarily generates a local stress, one long called a bit inaccurately - force of crystallization. Through this local stress, the growing crystal or crystal aggregate makes room for itself within the surrounding rock in three possible ways: by dissolving, by displacing, or by fracturing its surrounding matrix. The continued growth of a crystal in a rock presupposes some supersaturation and needs sufficient transport of intergranular species to and from the growth site. Because the force of crystallization is local, it naturally can interact with other local phenomena involved in its own genesis - specifically, with the very kinetics of the crystal growth that generates the stress, with the kinetics of pressure solution driven by it, with local transport properties, andor with the rheology of local deformation or fracturing of the surrounding rock matrix. It is for this reason that water-rock-interaction calculations -- which utilize mineral growth and dissolution rates, and permeabilities and porosities - would become more realistic if they incorporated the force of crystallization. Reaction-transport models now in use (e.g., Steefel & Lasaga 1994) do not provide for stress. Incorporation of force-of-crystallization models into geochemical water-rock reaction-transport calculations would warrant close consideration of the work

2 FEEDBACKS INVOLVING THE FORCE OF CRYSTALLIZATION 1) Mineral replucement. When a growing crystal pressure-dissolves the surrounding matrix, the force of-crystallization stress acts not only on the adjacent 161

has a much larger rate constant than that of the matrix mineral; see Eq. (9) below.

matrix grains, accelerating their dissolution, but on the growing crystal itself - slowing down its own growth rate. Thus, the volumetric growth rate and dissolution rate quickly come to be automatically equalized: this explains why mineral replacement characteristically preserves volume - a longununderstood phenomenon. See Figure 7 in Nahon & Merino (1997). Also: Maliva & Siever (1988), Merino et a1 (1993), Fletcher & Merino (1997, 2000), and Merino & Dewers (1998). Note that for this constant-volume replacement feedback to work it is essential that the kinetics of mineral growth and pressure solution be functions of the local stress - a factor ignored in current kinetic laws. The widespread occurrence of replacement in rocks of many kinds (references in Merino & Dewers, 1998) attests to the existence of that dependence in rocks.

3 MINERAL GROWTH ACCOMMODATED BY

REPLACEMENT In this section we model the growth of a spherical crystal of mineral A, of current radius a, in host rock made of mineral B. The volume of A is accommodated by dissolution of B. The stress far from the A/B interface, beyond a “mineralized zone,” M Z , of radius R (with R >> a), is assumed to be a homogeneous hydrostatic stress, GO. The macroscopic homogeneous hydrostatic stress within the MZ is OR, which differs from GO because mineral growth in the M Z results in a macroscopic dilation. The current stress at the AB interface is o,(a) = o n . For simplicity we assume that both A and B have the same Young’s modulus E and Poisson’s ratio v. We suppose that transport of aqueous species to and from the site of deposition and dissolution occurs along intergranular fluid films and does not limit growth or dissolution. The chemical potentials of the two components at the A/B interface are

2 ) Earthquake triggering. In the third of the responses to force of crystallization listed in the Introduction, growth of a crystal aggregate causes local fracturing or cracking of the rock around it. The 10cal fractures caused by nearby aggregates growing simultaneously in a rock may cooperatively interact with each other to become a larger, through-going fracture. If the rock already had sufficient stored elastic energy to begin with, that fracture could trigger its release as an earthquake.

Eqs. (1) ignore minor terms in stress multiplied by elastic strain (Kamb 1959). Growth of the A crystal is driven by a supersaturation ApA = RTlnfiA in the A component relative to the equilibrium value at the far-field normal stress, GO.Component B is assumed to be just saturated, f i = ~ 1, also at the reference normal stress, 00. VoA and VoA are the specific volumes of the minerals in an unstressed state. At the AA3 interface the equilibrating chemical potentials of the two components are

3 ) Ore b o 4 self-localization. The local fiacturing and cracking caused by growth of a crystal aggregate at a site in a rock (again, the third response listed) would improve the local permeability around the aggregate. This permeability increase would attract a larger flux of mineralizing solution to the site in question. The larger flux would feed further crystal growth, which would further fracture a new shell of rock, which would again increase the local permeability, and so on. The result would be a large, well-localized mineral body.

4) Deformation of sedimentary laminations and metamorphic schistosity by growth of crystals or crystal aggregates. Both sedimentary laminations wrapping around a concretion and schistosity wrapping around a garnet porphyroblast could be interpreted as produced by the force of crystallization developed by the concretion or the garnet. This suggestion was already made for porphyroblasts by Misch (1971) and others later, and it was criticized on the basis that the force of crystallization could not be large enough to deform the rock locally. The equations we develop below allow us to estimate both the growth-driven stress and its kinetic consequences quantitatively: the stress could in fact be large enough to deform rocks in some cases, especially at high temperature and supersaturation andor where kAkB >> 1, that is, where the growing crystal

We assume that the rates of growth or dissolution are linear in the chemical-potential difference, (4) (5) where kA and kB are kinetic rate constants, and a’ is the change in the radial position of the interface due to the (negative) growth of host mineral B. Since the only mechanism of accommodation considered here is replacement, we require (see Figure 7, Nahon & Merino 1997: Merino & Dewers 1998) dddt + da’/dt = 0

162

(6)

since the rates of A growth and B dissolution become mutually equal. A small elastic accommodation of the host around the A crystal also occurs, but is negligible. Combining (1) through (6), we obtain

The second relation in (7) yields the force of crystallization, or normal stress, at the AA3 contact as a function of the specified supersaturation

and, substituting (8) back in the first relation in (7), the growth rate of the A crystal turns out to be f k ~ v o ~ ) ] dddt = kAApA[kBVoB/(kAVoA

(84

Since o n < GO,or more compressive, the force of crystallization given by (8) is positive. If the supersaturation is constant during the growth of mineral A in the MZ, the interfacial normal stress must be constant as well, as long as replacement is the only mechanism of accommodation. The normal stress difference given by eq. (8) is maximum if kA/kB >> 1 (that is, if the host mineral B has a small lnetic constant relative to that of the growing A crystal), and minimum if kA/kB >l as in “quartz in calcite”, the force-of-xln and growth rate are smallest (eqs. 10,8a). For k$k,3 km)depth

Figure 3. The relationship between Si concentration and run duration in the case of quartz dissolution at (a) 30 MPa and (b) 60 MPa at various temperatures.

The dissolution rate of quartz was defined as the average reaction rate, ranging from time=O to 24 hours. Figure 3 indicates that the dissolution reaction of quartz depends on both temperature and pressure, although we argue pressure dependence

self-sealing at elevated temperatures (i.e. >350"C), evident in natural systems, has major implications for development and energy utilization of supercritical DSGR. 210

under supercritical condition. In Figure 4,we show dissolution rate as a function of specific volume of supercritical water, which is the reciprocal of water density. With increasing specific volume, dissolution rate decreases exponentially, thus specific volume is a suitable thermodynamic variable with which to demonstrate the capability of the fluid as a solvent. In Figure 5, we highlight the dissolution of granite,

Figure 6 . Logarithmic diagram showing the relationship between specific volume and dissolution rate of quartz based on the kinetic equation described in the text.

5 DISCUSSION The rate expression of quartz dissolution under suband supercritical conditions could be described as following a nth order kinetic equation: Figure 4. Relationship between specific volume of supercritical water and dissolution rate of quartz at various temperatures.

as represented by fluid Si-concentration in our autoclave experiments. At temperature-pressure conditions less than the critical temperature, high Si-concentrations were determined, whilst in a high temperature/low pressure region low Siconcentrations are indicated. A strong water-rock

Figure 5. Dissolution of granite, showing circles at various temperature and pressure. SVPC: Saturated Vapor Pressure Curve of pure water.

interaction, inferred by the dissolution of granite, was obtained in both sub- and supercritical regions, ranging from 300°C to 450°C. However chemical reaction processes tend to be relatively weaker in the high temperature region.

r=kV“, where Y is the dissolution rate, k is the apparent kinetic constant, V is the specific volume of fluid, and n is apparent reaction order. The apparent reaction order, n, can be evaluated on a log-log diagram, as shown in Figure 6. The slope of the ‘approximation line’ in the region of low specific volume is minus unity, which indicates the dissolution rate is proportional to water density, although the slope changes drastically in the region of higher specific volume. The steep sloping approximation lines in the region of high specific volume, at ‘so-called’ super critical conditions, indicates that the fluid at some elevated T-P condition loses its capability as a solvent, and becomes less effective in dissolving the rock. The ‘turning point’ of the dissolution rate on the logarithmic diagram (Fig. 6) may be plotted on a P-T diagram (Fig. 7), where the supercritical region is defined conventionally as being at T-P conditions of higher temperature and pressure than the critical point. In our granite and quartz dissolution experiments, we demonstrate that the supercritical region can be subdivided into two regions, based on variations in the intensity of the water-rock interactions. In the relatively high-pressure region, at supercritical conditions, the fluid shows high potential as a solvent, somewhat similar in the respect to subcritical liquids. In contrast, a weak water-rock interaction is evident at greater temperatures, within 21 1

ACNOWLEDGMENTS The authors would like to express our gratitude to Prof. Yamasaki, Dept Geoscience and Tech, Tohoku Univ., for his critical discussion. This study was financially supported by Grant-in-Aid for Research for the Future Program (JSPS-RFTF 97P00901). REFERENCES

Figure 7. Liquid-like and vapor-like regions under supercritical conditions. Circles shows Si concentration for granite dissolution in Fig. 5. Open squares indicate the ‘turning point’ represented in Fig.6.

the supercritical region. Thus we conclude that the turning (or inflection) points for the rate of quartz dissolution point to the existence of an apparent phase boundary within the supercritical region, as indicated in Figure 7. In the past, the supercritical region has been regarded as a homogeneous state, and neither classified as a liquid phase nor a vapor phase, and this inference may not be appropriate. Morita et al. (2000) recently described density fluctuation in the supercritical region of pure water, using a high temperature small-angle X-ray scattering technique. Their work supports our view that supercritical region may be subdivided into two ‘phases’, one being a liquid-like region and the other a vapor-like region, at least with respect to the dissolution of granite and quartz.

Bando, M., K. Sekine, G. Bignall, & N. Tsuchiya 2001. Exceedingly rapid emplacement of a Quaternary volcano-plutonic complex: the Takidani pluton and associated volcanics, Japan. J. Vol. Geothem. Res. (submitted). Doi, N., 0. Kato, K. Ikeuchi, R. Komatsu, S. Miyazaki, K. Akaku & T. Uchida 1998. Genesis of the Plutonic-Hydrothermal System aroud Quaternary Granite in the Kakkonda Geothermal System, Japan.. Geothermics 27: 663-690. Fournier, R. 0. 1999. Hydrothermal processes related to movement of fluid from plastic into brittle rock in the magmatic-epithermal environment. Econ. Geol. 94: 1193-1211. Ikeuchi, K., N. Doi, Y. Sakagawa, H. Kamenosono, & T. Uchida 1998. High-temperature measurements in Well WD-la and the thermal structure of the Kakkonda geothermal system, Japan, Geothermics 27: 591-607. Morita, T., K. Kusano, K., Ochiai, H., Saitow, K., & Nishikawa, K. 2000. Study of inhomogeneity of supercritical water by small-angle x-ray scattering. J. Chem. Phys. 112: 4203-4211.

6 CONCLUSIONS A Deep-seated Geothermal Reservoir (DSGR) may be expected beneath a convective hydrothermal system and may have great potential for geothermal energy extraction. Key factors for the utilization and development of DSGR are the creation of a large surface for heat exchange, such as a high density fracture network and/or a porous media, at conditions close to those of the brittle-plastic transition, where weak water-rock interaction might preserve the heat exchange surface without plugging and self-sealing. The supercritical region can be subdivided into two phases, one liquid-like and the other a vapor-like regions. The vapor-like region, being a relatively low pressure region, at supercritical conditions, has the greatest potential to satisfy criteria necessary for development of a supercritical DSGR.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, lSBN 90 2651 824 2

Porewater geochemistry and modeling within Oligocene-Miocene clays of North Central Spain M.J.Turrero, J.Pefia, A.M.Fernhndez, P.G6mez & A.Garral6n CIEMAT, Departanlento de Impacto Ambiental de la Energia, Madrid, Spain

ABSTRACT: The characterization of the interstitial waters of clays collected from a borehole of 650 m depth, and the vertical distribution of dissolved constituents was made in order to identify sources of waters and water-rock interaction processes. The study of the samples included mineralogical analyses of the rock, cation exchange capacity and exchanger population, chemical analyses of porewater samples and integration of results using the PHREEQC geochemical code. The porewaters of the clay samples were extracted at room temperature by squeezing at high pressure (60 MPa). The porewaters analyzed were Na-S042-, with ionic strength ranging from 0.12 to 0.29 mol/l. The evaluation of the fluid composition, mineralogy, water-rock reaction processes and modeling, indicated that oxidation and degassing processes during collection and storage of the samples could occur. Taking this into account, the original porewater of the formation was reconstructed. 1 INTRODUCTION

2 SAMPLES AND METHODS

In clay formations, with low water-rock ratios, the effect of the water-rock reactions is enhanced. The mineralogy and the obtention of porewaters from clays can provide key information to reconstruct fluids controlling reactions taking place in the system. When the samples are analysed in the laboratory, the extent to which water-rock reactions have proceeded depends on the way of collection and conservation of the samples until analyses (Baeyens et al. 1985). The clays investigated are alluvial to lacustrine sediments of Oligocene-Miocene age and occur throughout the Duero Basin (North-Central Spain), in a continental and mainly endorheic environment. In this work emphasis was made in studying the Reference Spanish Clay (AER), a Clay Formation with a thickness around 300 m, occurring at depths below 80 m (early low temperature diagenesis). The studied well is located in the Northeastern area of the basin, in which AER consist of a 106 m thick basal lutite sequence, which is dominated by red and green lutites with some sandstones alternations deposited by fluvio-alluvial processes. The upper part is a 198 m thick lutite to mar1 sequence, with gypsum and carbonates increasing towards the top, deposited in a predominantly lacustrine environment (Turrero et al. 1998).

Porewater samples and analyses were obtained from eighteen cores from a borehole of 650 m depth (ICl), drilled in the Oligocene-Miocene clays of the North-Central Spain. The samples were isolated from the atmosphere by means of PVC tubes sealed with paraffin to minimize oxidation or degassing, and were stored in a room with high relative humidity and constant temperature until analyses of both rock and porewaters. 2.1 Porewater extraction and cation exchange determination The porewater samples were extracted using the compression technique described in Cuevas et al. (1997), similar to that developped by Peters et al. (1992) and Entwisle & Reeder (1993). The porewater was extracted by applying a constant pressure of 64 MPa onto the sample, under laboratory conditions (25"C), and avoiding contact with atmospheric air by means of a close extraction circuit. The accuracy of the porewater chemical data was assessed through charge balance calculations. The moisture content of the individual samples was calculated from the ratio between water weight loss after heating sample to 110°C for 24 hours, and the weight of the dried clay, expressed as percentage. The Chapman displacement method was used to determine the exchangeable cations by means of

213

successive washing with ammonium acetate 1N at pH=8.2, after flushing the soluble salts (Thomas 1982). To determine the cation exchange capacity, the exchange sites of the sample were saturated with sodium by means of successive washing with sodium acetate 1N at pH=8.2. The adsorbed sodium was displaced by successive extractions with ammonium acetate 1N at pH = 8.2 (Rhoades 1982).

(TDS ranging from 1000 to 10000 mg/l). pH ranges from neutral to slightly alcaline (6.9 to 8.1). The ionic strength is around 0.2 moVl for all the samples. Table 1. Characteristics of the sample C.290 (depth 286.2286.4 m) belonging to the basal sequence of AER.

Porewater comDosition Exchanper DoDulation [MI meq/IOOg moVL Element c14.5 x IO-’ NaX 3.48 0.262 so:CaX2 12.38 0.465 9.3 x IO-’ Alkalinity 8.8 x 10-4 MgX2 2.06 0.077 ca” 1.9 x 10-’ KX 1.6 0.12 Mg” CEC 19.57 1.5 x 10-’ Mineralogical composition Na’ 1.7 x 10-’ K’ 8.1 x 10-4 Quartz 21% Sr’’ 3% 1.2 x 10-4 K-feldspar 2.2 x 10-3 Phyllosilicate 60% Fe,,, Ba 6.5 x 10’7 Calcite 14% 10.53 pH Dolomite 3% mol/L 8.1

2.2 Chemical analyses Total dissolved Ca, Fe, Mg, Ba, Na, K and Sr were determined by ICP-AES on a Perkin-Elmer Elan 5000 spectrometer. Major anions Cl- and SO-: were analyzed using a Dionex ion chromatograph. The alkalinity measurements were made on a Metrohm 682 titroprocessor. pH was measured using an Orion EA 920 pH meter. The conductivity was determined using an ORION 115 conductivimeter. 2.3 Geochemical modeling The geochemical data were interpreted with the geochemical code PHREEQC (Parkhurst & Appelo 1999) and the WATEQ4F modified thermodynamic database (Ball & Nordstrom 1991). 3 RESULTS

3.2 Conceptual model Geochemical equilibrium calculations performed using PHREEQC indicate that the water closely approaches saturation with respect to the sulfate minerals celestite, anhydrite, gypsum and barite. However, an excess of SO‘: concentration was measured in orewaters, related to the total of Ca2+,Ba2’ and Sr concentration. Even considering that sulfate could be associated to sodium or magnesium, its concentration is higher than expected. Hence, the excess of SO‘: concentration is probably derived from sulfide (mainly pyrite) oxidation reactions. The sulfide oxidation is assumed to be controlled by natural oxidation by the supply of atmospheric oxygen into the samples during the storage period. Like this, the system exchan ed 0 2 and CO2 with atmosphere, and sulfide and Fe were oxidized to sulfate and Fe3+,precipitated as Fe(OH)3 (supersaturated as indicated in the calculations made with PHREEQC). Calcium required for gypsum precipitation is assumed to be present from calcite dissolution, triggered by the H+ produced by pyrite oxidation, buffering this acid production. Furthermore, carbonate minerals analyses indicate that the carbonate minerals are being dissolved because of porewaters undersaturated conditions; SEM images provide evidences on calcite partially dissolved and with corroded edges. The oxidation of pyrite is described as (Appelo & Postma 1996):

zp +

3.1 Rock and porewater composition The composition of the clays is briefly described in Pelayo and C6zar (1999). Mineralogical composition of the basal sequence of AER consist of 10-30 % quartz, 6-30 % calcite, 1-4 % dolomite, 50-70 % clay minerals (0-2 % smectite, 85-90 % illite and 1012 % chloritekaolinite) and accessory feldspars, gypsum, barite, celestite, ilmenite, iron oxyhydroxides and pyrite; the upper part of AER consist of 2-9 % quart, 6-38% calcite, 0-8 % dolomite, 4565 % clay minerals (45-60 % illite, 30-50% smectite and 3-7 % chloritekaolinite), 2-4 % gypsum and accessory feldspars, barite, celestite, iron oxyhydroxides, ilmenite and pyrite. The composition of all the porewaters is given in Turrero et al. (2000). Table 1 shows a representative porewater composition, exchanger population and mineralogy from the basal sequence of AER (sample C.290). This sample was selected to do the reconstruction of the original water in the formation, based on water-rock reaction processes and modeling. In all the samples analyzed, the charge balances were better than & 10%. Most of the waters are sodium-sulfate type waters, with high concentration of chloride, calcium and magnesium. According to Hanor (1987) the waters can be classified as brackish

F+

FeSz(s) + 15/402+ 7/2H20 = Fe(OH)3(s) + 4H’

+ 2S042-

Carbonate dissolution and acid neutralization can be represented by the following reaction: 214

Exchanger population was calculated considering that sodium is conservative (balanced with chloride), except for the exchanger population. Magnesium and potassium were mantained as measured (Table 1). Calcium was calculated to mantain the total cation exchange capacity in 19.57 meq/100g, since the total CEC of the clays not varies, only the exchanger population can change. The porewater was equilibrated with calcite, by using the calcium concentration calculated from the initial exchanger population. The pC02 was determined considering equilibrium with calcite at the initial pH. 02(g) and C02(g) were added to the water during simulation, until atmospheric values were reached. Table 3 shows the calculated original water. Results of modeling, with the variations from this original water until the measured one (Table l), are depicted in Figures 1 , 2 and 3.

CaC03(s)+H'= Ca2'+ 2HC03The excess of sodium concentration with respect to chloride is attributed to cation exchange processes. High potassium concentrations are derived from silicates transformation and equilibrium with the exchanger population. Magnesium is consumed in the process of chlorite formation (there is a positive correlation between [Mg] and % wt. of chlorite). Calcium is controlled by calcite dissolution and cation exchange.

4 MODELING THE CHANGES OF THE POREWATER CHEMISTRY PHREEQC was used to simulate the changes of the water chemistry, assuming that the initial water chemistry in the formation was Na-C1 and evolved, as a consequence of an oxidation process during storage, towards a Na-SO2- type water. For modeling purposes, it was supposed that the oxygen difusivity was constant and depends only on the moisture content (w)of the clay. The moisture content represents a single measurement made when samples were squeezed, after the storage. For the sample C.290 w=13.3%. The values of the exchanger population (meqA00g) and the minerals (% wt.) were calculated to moles of exchanger or mineral in contact with a liter of solution (mol/L) since PHREEQC code uses this unit. The selectivity constants (Table 2) were calculated using the Gaines-Thomas convention (Gaines & Thomas 1953).

Table 3. Characteristics of the calculated original porewater composition for the sample C.290. All the concentrations are expresed in molesL. C14.5 x 10' pH 8.1

so,"-

Alkalinity Ca2' Mg2+ Na' K'

Table 2. Selectivity constants of I" + iX-= IX; (considering logKNdu = 0) for sample C.290. The concentration of I" was calculated with PHREEQC.

Free specie (p) log K c a log K M g log KK

1.808 1.174 2.038

4.1 Hypothesis considered for calculations and results It is assumed that originally Cl- was the dominant anion, and its concentration was similar to that measured in the samples, since it is a conservative element. The excess of SO:- concentration is assumed as derived from sulfide (mainly pyrite) oxidation reactions. Hence, the system was assumed to be originally in equilibrium with pyrite. Since there is not available data of pyrite, the pyrite content considered initially (0.08 %wt.) to equilibrate the porewater, was enough to produce, by means of an oxidation process, the SO:- concentration that is actually measured.

5.7 10" Pe -3.86 2.7 x 10" Exchanger pop. (mol/L) 6.7 x 10-4 NaX 0.363 CaX2 0.375 5.5 x 10*4 5.6~ 10-2 MgX2 0.072 1.6 x 10'4 Kx 0.1 13

Figure 2 represents the variation, after adding 02(g) and C02(g) (final pOz=10-0.7atm. and P C O ~ = ~ Oatm.), - ~ . ~of the ionic content in the water from the Na-Cl initial calculated water until the measured Na-SO?- water. Figure 3 depicts the inverse behaviour of sodium and calcium concentration. Oxidation promotes a decreasing of sodium concentration in the exchanger, at the same time that calcium concentration increases. The pyrite oxidation reaction increases the sulfate content in the water, that forms complexes with calcium, CaS02. Hence, Ca2+ decreases and the exchanger equilibrium is modified. Figure 4 shows the variations of pyrite, Fe(OH)3(a), Ca2', SO?- and CaS02.

215

5 CONCLUSIONS The evaluation of mineralogy and porewater composition of AER, as well as the definition of water-rock reaction processes and modeling, pointed out problems of oxidation and degassing processes, likely during collection and storage of samples. Geochemical modeling, used in this work to quantify geochemical processes, has served as a powerful tool to reconstruct the original water of the formation, tak-

REFERENCES

Figure 1. Variation of ionic water composition from the initial Na-CI until that measured Na-SO?-.

Figure 2. Variation of the exchanger ComPosition from the hYpothetic initial until the measured one.

Figure 3. Variation of the reactive mineralogy through the oxidation process.

ing into account this water-rock interaction processes derived from observations and analyses.

ACKNOWLEDGEMENTS This research was funded by ENRESA within the frame of the third R&D plan.

216

Appelo C.A.J. & D. Postma 1996. Geochemistry, groundwater and pollution. Balkema, Rotterdam. Baeyens B., Maes A., Cremers A. & P.N. Henrion 1985. In situ physico-chemical characterization of Boom Clay. Radioac. Waste Man. Nucl. Fuel Cycle, 6: 391-408. Ball J.W. & D.K. Nordstrom 1991. User's manual for WATEQ4F, with revised thermodynamic database and test cases for calculating speciation of major, trace, and redox elements in natural waters. USGS, Open File Report 91183,189 pp. Cuevas J., Villar M.V., Fernandez A.M., Gdmez P. & P.L. Martin 1997. Pore waters extracted from compacted bentonite subjected to simultaneous heating and hydration. Applied Geochemistry 12: 473-481. Entwisle D.C. & S. Reeder 1993. New apparatus for pore fluid extraction from mudrocks for geochemical analysis. In: Geochemistry of Clay-Pore Fluid Interactions. Manning Hall & Hughes (eds.). Chapter fifteen: 365-388. Gaines G.L. & H.C. Thomas 1953. Adsorption studies on clay minerals. 11. A formulation of the thermodynamics of exchange adsorption. J. Chem. Phys. 21: 714-718. Hanor J.S. 1987. Origin and Migration of subsurface sedimentary brines. SEPM short course no 21,247 pp. Parkhurst D.L. & C.A.J. Appelo 1999. PHREEQC (Version 2): A computer program for speciation, reaction-path, 1D transport, and inverse geochemical calculations. USGS, Water-Resources Investigation Report. Pelayo M. & J. C6zar 1999. Estudio mineral6gico y geoquimiCO de las arcillas del Oligoceno-Mioceno del borde oriental del Macizo IbCrico (sondeo IC-1). CIEMAT/DIAE/54221/5/99. Peters C.A., Yang Y.C., Higgins J.D. & P.A. Burger 1992. A preliminary study of the chemistry of porewater extracted from tuff by one-dimensional compression. Water-rock Interaction. Kharaka & Maest (eds.): 741-745. Rhoades J. D. 1982. Cation Exchange Capacity. In: Methods of Soil Analysis, Part 2. Agronomy Monograph no 8 (2nd Edition). ASA-SSSA, 677. WI 537 11. USA. Thomas G.W. 1982. Exchange cations. In: Methods of Soil Analysis, Part 2. Agronomy Monograph no 9 (2nd Edition). ASA-SSSA, 677. WI 5371 1. USA. Turrero M.J., Peiia J., G6mez P. & A. Garral6n 1998. Origen y caracteristicas de las muestras de 10s materiales arcillosos de la Cuenca del Duero para su estudio en el context0 del Proyecto Mar. CIEMAT/DIAE/54221/2/98,21 pp. Turrero M.J., Peiia J., Fernandez A.M., G6mez P. & A. Garra16n 2000. Modelizaci6n Hidrogeoquimica de la Arcilla Espaiiola de Referencia (AER). CIEMAT/DIAE/54221/2//OO, 70 PP.

Water-Rock Interaction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

pH calculation through the use of alkalinity in geochemical modeling of hydrothermal systems M .P.Verma Geotermia, Instituto de Investigaciones Electricas, Apartado Postal 1-475, Cuernavaca 62001, Mor., Mexico

A .H.Truesdell Consultant, 700 Hermosa Way; Menlo Park, CA, USA

ABSTRACT: The calculation of deep reservoir physical and chemical parameters in both liquid and vapor phases is the first step in geochemical modeling of hydrothermal systems. The reservoir temperature is calculated using the quartz solubility regression equation along the liquid-vapor saturation curve, assuming equilibrium between the reservoir liquid and quartz, and no steam loss or gain during the ascent of the fluid to the surface through the production well. Knowing the reservoir temperature, the chemical parameters (except pH) in vapor and liquid phases are calculated using the conservation of enthalpy and mass. The approach is extended to calculate the reservoir pH using the conservation of alkalinity. 1 INTRODUCTION Modeling of chemical processes occurring in a geothermal system is basically the evaluation of reservoir fluid-mineral equilibrium-state from consideration of the locations, appearance and chemistry of surface natural manifestations and the flow, pressure, enthalpy and chemistry of well discharges. Routine measurements are the concentration of dissolved species and pH in separated and flashed water from weirbox and the contents of gaseous species in vapor collected at the pressure (temperature) condition of the separator. The chemical composition of reservoir fluid is reconstructed as a mixture of separated liquid and vapor. Thus it requires first the calculation of temperature and chemical composition of liquid (including pH) and possible vapor present in the reservoir, in order to construct a chemical equilibrium (speciation) model of reservoir fluids. The total discharge concentrations of dissolved species (like Na’, K’, Cl-,...) are calculated using mass and enthalpy balance equations (Henley et al., 1984). The reservoir temperature and vapor-fraction are estimated through chemical and gas geothermometers and the distribution of gaseous species (Giggenbach 1980). Verma (1997) presented a two-phase flow approach to calculate the reservoir concentration of species dissolved in the liquid phase other than pH. Reed & Spycher (1984) pointed out that the previous methods for calculating pH at high temperature from the analyses at low temperature, without the use of equilibrium among minerals 217

(Truesdell & Jones 1974; Merino 1979; Arnorsson et al. 1982) were based on the estimate of “total ionizable hydrogen”, which they considered an illdefined quantity compared to the abundance of hydrogen ion. This requires the assumption that only chloride analyses were incorrect. Truesdell & Singers (1974) and Glover (1982) contributed considerable to the development of this method. Some of the problems encountered in these approaches may be avoided by use of alkalinity. Alkalinity is a conservative quantity in some chemical reactions (see below) and is independent of changes in pressure and temperature. For example, if two waters are mixed together, their alkalinity will be defined by a conservative mass balance, provided that no pertinent mineral phase (e.g., calcite) precipitates upon mixing. Similarly, if a solution is heated or cooled, changes in the dissociation constants of weak acids and bases will produce or consume H+ or OH- ions. Thus the pH of the solution will generally change with temperature, but not its alkalinity. This article presents a technique to calculate pH through alkalinity taking into account the effects of heating, boiling, dilution and mixing of two or more fluids.

2 DEFINITION OF ALKALINITY A base-neutralizing capacity (BNC) or acidneutralizing capacity (ANC) is the equivalent sum of all the acids or bases that can be titrated with a strong base or acid to a preselected equivalence point (Stumm & Morgan 1981). The BNC and ANC

are more commonly known as alkalinity and acidity, respectively. Both of these terms are defined for certain pertinent equivalence points (EPs) for the system. Acidity is the negative of alkalinity for the same reference EP. In carbonate systems there are three equivalence points called the H,COS EP ,

HCOYEP and CO,'-EP. Alkalinity could be defined with respect to either EP. However, for geothermal fluids the alkalinity (or acidity) with respect to the H,COlEP is most useful. For the carbonate system alkalinity is defined as: alk= [HCO,]+2[CO;-]+[OH-]-[H+] (1) The term in brackets is the molal concentration of the species. Because the geothermal fluid has also other weak acids and bases, the alkalinity is defined here as: afk= [ O K ] +[ H C S ] +2[CG-]+[B(ON)i]

where the a's identify the ionization fractions (Stumm & Morgan 1981) and CT is the total dissolved concentration of the subscripted constituent, i.e., carbonic acid (car), boric acid (B), silicic acid (Si), hydrogen sulfide (S) and ammonia (N), respectively. Chemical speciation can be reconstructed introducing in Eqn. (2) pH, alkalinity and total dissolved concentrations of relevant constituents. It is important to point out here that we are interested in the dissolution-exsolution of NH3, but not of NH,' or its salts like NH4C1. Therefore we defined the alkalinity with respect to the NH,EP

concentration of CZ- in their datasets in order to get initial charge balance condition. In other words one has to justify that the analysis of Cl- was only incorrect in the datasets, considered by Reed & Spycher (1984). Contrary, it may be possible that the initial pH (or H' concentration) or a cation concentration (for example Nu') or other anion (HCO;, or SO,"-) is incorrectly analyzed. If we adjust the initial (analytical) concentration of H' or HCO, , SO,"-, the results of pH calculation will be quite different. Although the adjustment of initial concentration of Na' or CZ- does not affect significantly the final pH value, it affects the chemical speciation of the solution. Theoretically, a solution should be electrically neutral, but the electro-neutrality condition is rarely satisfied, even in best quality analyses. Thus the alkalinity approach is safer for the pH calculation of hydrothermal fluids. This approach is a continuation of work initiated by Merino (1975, 1979).

3 RESERVOIR PARAMETERS CALCULATION 3.1 At the well separator Generally the liquid separated in the separator is flashed in the silencer at atmospheric pressure and the water sample is collected from the weirbox. Some fraction of vapor with non-condensable gases is lost in the atmosphere. The chemical composition of separated water (sep) can be back calculated from the analyzed composition of the water sample flashed to the atmosphere (atm), by using the following procedure, which is based on mass and energy balances (Henley et al., 1984). The fraction of vapor lost at the weirbox is:

instead of NH,'EP in Eqn. (2). The procedure of writing the alkalinity expression for different types of reactions in a system is explained by Stumm & Morgan (198 1). Thus the alkalinity defined here does not charge upon dissolution or exsolution of CO;! (HiCO3 ) and other gases, such as H2S and NH3. On the other hand, the addition or removal of CaC03 or other carbonate minerals, and Ca(0H)z or other hydroxides, will increase or decrease alkalinity . There are three types of equations for an aqueous solution: mass balance, charge balance and proton balance. But out of the three equations two are independent and the third can be derived as an algebraic sum of the other two equations. Reed & Spycher (1984) used the total moles H' as conservative quantity in order to calculate pH with temperature change. They had to adjust the

(3)

The chemical concentrations of dissolved species @a+, K', .. . ) in the separated water are given by: Cl.iiep

= C,.arm * (1 - Y a m )

(4)

Most dissolved gases like C02, H2S, NH3, CH4, etc. are lost from water samples during the sampling to analysis time. The concentration of these gases can be measured more accurately in the vapor phase. Therefore, the concentration of these gases in the liquid phase is calculated from their distribution coefficients at the separation pressure:

B = C,/C,

(5)

The calculation of pH is not as simple as that for the 21a

Table I . Physical-chemical parameters of geothermal fluid at various positions in the well M-19A at Cerro Prieto. The values are reported up to 3 decimal points for sake of comparison. The actual accuracy depends on the analytical error (modified after Verma 2000). Parameter

(2 5°C)*

Weirbox

(1 0O0C)**

1 IOO.00 142.72

Separator

Corrected data at SeDarator 7.55 7.55 168.06 168.06 142.72 142.72 63.20 63.20 0.3 1 Liquid Phase (mmolkg water)

***

Pressure (Bar) Temperature ("C)

1 25 97.8

Na' K+ Ca" Mg'i Li'

320.577 42.457 10.928 0.0 16 28.8 14

2 19.678 29.094 7.489 0.01 1 19.745

2 19.678 29.094 7.489 0.01 1 19.745

1.332 1.318 0.014 13.448 13.410 0.038 0.958 0.105 0.852 7.42E-4 389.248 0.187 7.27 0.906

0.9 13 0.903 0.01 1 9.2 15 9.136 0.080 0.656 0.127 0.529 4.02E-4 266.735 0.128 7.04 1 0.621

0.9 13 0.903 0.01 1 9.2 15 9.136 0.080 0.656 0.127 0.529 4.02E-4 266.735 0.128 7.04 1 0.62 1

219.678 29.094 7.489 0.01 1 19.745 1.040 0.131 0.908 0.243 0.102 0.141 0.913 0.90 1 0.0 12 9.2 15 9.065 0.150 0.770 0.32 1 0.449 1.61E-4 266.735 0.128 7.075 0.62 1

Wellhead

Reservoir

35.00 242.30 187.83 18.09 0.09

37.67 246.82 190.04 15.88 0.08

166.920 22.107 5.690 0.009 15.003 2.284 0.125 2.158 1.707 1.568 0.139 0.693 0.692 0.00 1 7.002 6.978 0.024 7.150 6.7 17 0.444 8.09E-6 202.675 0.098 6.479 0.472

164.978 21.850 5.624 0.008 14.829 2.375 0.128 2.247 2.022 1.885 0.137 0.686 0.684 0.00 1 6.92 1 6.900 0.02 I 8.827 8.391 0.436 6.54E-6 200.3 18 0.096 6.432 0.466

15.698 1.344 0.820 0.589 0.105 0.196

17.503 1.459 0.936 0.67 1 0.120 0.198

0.32 1" 0.102 2.30E-3 1.26E-3 1.44E-4 0.908 Vapor phase (mmol/mol steam)

* @

a

** ***

4.8454 0.466 0.235 0.169 0.030 0.136

4.845 0.466 0.235 0.169 0.030 0.136

Analytical data for liquid phase Analytical data for vapor phase Concentration obtained through gas distribution (in mmol gaskg water) Just before flashing in the weirbox at the atmospheric conditions Just after vapor separation, the liquid phase concentration were calculated from the separated water

dissolved chemical species, because the dissociation constants of water and weak acids and bases change with pressure and/or temperature. Similarly the dissolution or exsolution of gases like CO2 also change the pH of the solution. Therefore, we use an

approach based on the alkalinity as defined in the above section. The alkalinity of separated water in the separator is obtained by Eqn. (4).Here a decision is needed for the total concentration of carbonic species, because the concentration of dissolved CO;, 219

can be calculated from the analysis of both liquid and vapor phases. This will be discussed in the next section.

calculating the reservoir temperature. The vapor fraction in the reservoir is obtained through an enthalpy balance. Based on these vapor fraction and reservoir temperature, chemical constituents are redistributed between coexisting vapor and liquid phases. Throughout the approach the alkalinity is considered as a conservative entity.

3.2 In the reservoir In order to calculate the deep reservoir composition, the temperature and vapor fraction are required. Verma (200 1) derived the following regression equation for quartz solubility along the water-vapor saturation curve for the temperature range 0-374°C: log sioz(ppm)= -1 175.7/T(K) + 4.88 (6) An iteration process, which considers conservation of mass and enthalpy and equilibrium between water and quartz in the reservoir, is used here to estimate temperature and vapor fraction in the reservoir. Once these two parameters are obtained, the distribution of gaseous species between the vapor and liquid phases is calculated using the approach of Giggenbach (1980) and the pH is computed using the procedure defined above, which is based on conservation of alkalinity.

5 CONCLUSIONS Since alkalinity is a conservative entity during dissolution or exsolution of gas species like CO2, HzS, NH3, etc., it is a powerful tool for calculating the pH of geothermal reservoir liquids, including the effects of boiling and mixing. This approach assumes chemical equilibrium between liquid and vapor phases in the geothermal reservoir fluid. ACKNOWLEDGEMENTS Dr. Luigi Marini read critically the earlier version of the manuscript.

4 A CASE STUDY: CERRO PRIETO

REFERENCES

In Table 1, the results of stepwise calculations for well M-19A are presented. The accuracy of these results depends on the quality of analytical data, whose discussion is beyond the aims of this work. It is possible to calculate the concentration of all carbonic species knowing pH and the concentration of one of such species. In this work, the concentrations of all carbonic species are calculated starting from pH and HC03-. Similarly the speciation of boric and silicic species and total alkalinity are calculated. Then this water is heated up to 100°C to get the concentrations of the flashed water at the weirbox. In the following step, the water is diluted for the vapor lost to the atmosphere at the weirbox. In this way, the concentration of the separated water just before flashing at 100°C is obtained. The water is again heated up to the pressure and temperature conditions of the separator. Alternatively, the concentration of undissociated carbonic acid can be calculated from gas analysis. Both values may have some errors due to nonequilibrium between vapor and liquid in the separator and re-equilibration in flashed water due to loss or gain of volatile species. It is assumed that the residence time of geothermal fluids in the Cerro Prieto reservoir is sufficiently high to allow attainment of chemical equilibrium between the liquid phase and quartz. Therefore the quartz regression equation is used for

Giggenbach, W.F. 1980. Geothermal gas equilibria. Geochim. Cosrnochim.Acta, . 4 4 : 2021-2032. Glover, R.B. 1982. Calculation of the chemistry of some geothermal environments. New Zealand D.S.1.R Chemistry Division Report CD 2323. Henley, R.W., Truesdell A.H. & P.B. Barton 1984. Fluidmineral equilibria in hydrothermal systems. Reviews in Economic Geology, 268p. Merino, E. 1975. Diagenesis in Tertiary sandstones from Ketleman North Dome, California-11. Interstitial solution: distribution of aqueous species at 100°C and chemical relation to the diagenetic mineralogy. Geochim. Cosmochim. Acta, 39: 1629-1645. Merino, E. 1979. Internal consistency of water analysis and uncertainty of the calculated distribution of aqueous species at 25°C. Geochim. Cosmochim. Acta, 43: 1533-1542. Reed, M. & N.Spycher 1984. Calculation of pH and mineral equilibria in hydrothermal waters with application to geothermometry and studies of boiling and dilution. Geochim. Cosmochim. Acta, 48: 1479-1492. Truesdell, A.H. & B.F. Jones 1974. WATEQ a computer program for calculating chemical equilibria of natural waters. U.S.G.S. J. Research, 2: 233-248. Truesdell, A.H. & W. Singers 1974. The calculation of aqueous chemistry in hot-water geothermal systems. U.S. Geol. Survey J. Res., 2: 27 1-278. Stumm, W. & J.J. Morgan 1981. Aquatic chemistry. Wiley, New York. Verma, M.P. 1997. Thermodynamic classification of vapor and liquid dominated reservoir and fluid geochemical parameter calculations. Geoflsica Internacional, 36, 18 1-1 89. Verma, M.P. 2000. pH calculation through the use of alkalinity in modelling of hydrothermal systems. Proc. 30"' Geotherm. Workshop. Stanford, 166-170. Verma, M.P. 200 1 Silica solubility geothermometers for hydrothermal systems. This Volume.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Scale versus detail in water-rock investigations 1: A process-oriented framework for studies of natural systems Richard B .Wanty, Byron R.Berger & Michele L.Tuttle U S . Geological Survey, M.S. 973 Denver Federal Center, Denver, CO 80225, USA

ABSTRACT: Studies of water-rock interaction rarely explicitly address the antithetical properties of scale and detail. Explicit treatments of these properties usually rely on statistical methods and focus independently on geological, hydrological, or geochemical data. While these approaches have the power to draw conclusions from large data sets, they lack predictive capability when data sets from other areas are considered, or when perturbations (e.g., climate changes) are imposed on existing data. However, when processes and properties of natural systems are understood, such a predictive capability may be gained. We propose a framework for considering the interaction of scale and detail of spatial and temporal properties of natural systems and processes, applied to mineralized rocks. Weathering of ore-related minerals, especially sulfides, provides the opportunity to trace flow along specific fracture sets, and demonstrates the isolation of some sets from others. Case studies presented in this paper, and its second part (by Berger et al. this volume), demonstrate scale-dependent phenomena in natural systems. Although our approach is qualitative, it allows for an efficient design of field and lab studies, and results in a fully integrated study of natural systems, as it includes geological, hydrological, and geochemical factors. 1 INTRODUCTION An assumption implicit in all environmental studies is that the collected samples represent the system under study at some spatial or temporal scale. Few studies ever test this assumption. Similarly, if a new area is to be studied, there exists no reliable protocol for defining a sampling strategy that will guarantee a representative sample set. Statistical methods are available to generate spatially unbiased (objective) sampling networks (cf. Miesch 1976), which have the advantage of producing unbiased geochemical baselines. However, such methods put little weight on important geologic variables such as lithology and structure. As a result, statistically designed, objective sampling grids provide no predictive capability for areas exceeding the study boundaries, or for other sampling times within the study area. In a coupled hydrologic-geochemical processoriented approach, the researcher’s objective is to test hypotheses related to the relative importance of various properties of the system. Therefore, sampling strategies will necessarily be subjective, as greater sample density will be required in geologically or hydrologically complex areas, or in areas with strong geologic or geochemical gradients. The problem the researcher faces, then, is to correctly recognize these complex areas, based on existing in-

formation or preliminary geologic, hydrologic, and geochemical reconnaissance studies. Prior geologic knowledge (usually in the form of geologic maps or reports) of a study area is a great advantage. This research strategy introduces the dilemma of scale of studies versus detail. In general, scales of study can be increased at the expense of detail, or vice versa. Given limited fiscal resources, a delicate balance exists between scale and detail. A budget might allow for the collection and analysis of a finite number of samples, so the spatial density of those samples is critically important. This paper presents a discussion and examples focusing on fractured, mineralized rocks, but the philosophical approach is general and broadly applicable. The presence of mineralized rocks documents a paleohydrologic system that could still be active today (it should be noted, however, that the driving forces behind fluid flow may be different; the former thermally driven flow system might now be gravity driven). Further, mineralized rocks may behave as localized sources of solutes in the weathering environment, providing a suite of “natural” tracers whose migration and attenuation in streams and ground water can be followed. Because of our focus on a specific mineral deposit type, the hydrologic regimes we will discuss in this and the following paper (Berger et al. this vol221

ume) are primarily fracture controlled. Conceptual and numerical models of fracture-controlled flow are still in development. In general, stochastic or deterministic models of fracture flow can be constructed using either a continuum approach or a discrete fracture network approach (Hsieh 1998). The model fracture networks rarely provide an exact match to a field situation (nor are they intended to); rather, they are constructed to match field data (e.g., aquifer tests) or observations. These models seem to have limited accuracy in that they fit the overall features of a flow system without matching the fine detail present in most aquifers. More importantly, though, because these models are usually based on statistical representations of field data, no reliable predictive capability is expected if the study area is expanded. From the analysis of spatial scales of various natural properties and processes presented in this paper, it should be possible to design more efficient field studies, as well as to assess the degree to which collected samples represent a natural system. An outgrowth of this approach is to suggest a strategy for field sampling that locates areas where geochemical and hydrologic gradients exist.

the overall geologic systematics, more reasonable model fracture networks might be constructed at a number of spatial scales, and the results of a truly geologic-integrated study might enjoy greater predictive capability. Further, an understanding of the geologic systematics might facilitate a more efficient field sampling strategy. 3 STUDY DESIGN AND OBSERVATIONS AT VARIOUS SPATIAL SCALES

Figure 1 may be used as a guide for organizing investigations of water-rock systems. By determining which processes and properties may be important in a field study, the appropriate sampling strategy may be developed. For example, a study of regional ground-water chemistry and flow in a fractured intrusive rock might include an area up to several tens of kilometers, but to assess flow in individual fracture sets in that area, there must be detailed areas within which several samples are collected within meters or tens of meters of each other.

2 SPATIAL SCALES OF NATURAL PROPERkm

TIES AND PROCESSES Figure 1 shows a number of properties and processes found in water-rock systems, and the scales at which we observe them. The list is not comprehensive, but demonstrates the wide range of spatial scales over which commonly studied natural system properties and processes occur. Some 16 orders of magnitude of distance are shown on the scale bar. The approximate spatial extent over which each property or process is relevant is shown by the horizontal bar. At the top of the figure the line labeled ‘deposit drainage’ demonstrates that the spatial extent of drainage from mineral deposits, from chemical genesis to physical transport, may span most of the spatial range shown in the figure. The focus of this paper is on mineral deposits, but the concepts presented are applicable to a wide variety of systems. Therefore, the ‘deposit drainage’ line could alternatively be labeled ‘water-rock interactions.’ Figure 1 is organized so that the properties, processes, and observations are shown from top to bottom in the rough order of geology 3 chemistry 3 hydrology 3 ecosystems. Each horizontal line in the figure is meant to show the actual process or property, but in fact, these lines also may be considered to show the spatial extent or relevance of conceptual or numerical models of the properties. The geologic properties and processes are shown at the top because conceptual models of geologic environments provide an overall context within which hydrogeochemical systems can be studied. By understanding

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.-

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optical & outcrop observations electron-beam remote sensing methods microscopy whole-rock microbeam analyses analyses water samples from wells, streams, & springs hydrologic measurements

Figure 1. Relevant spatial scales for properties, processes, and observations in systems of interest in water-rock investigations.

222

To assure that the best set of representative samples are collected, geologic, hydrologic, and geochemical investigations must be conducted simultaneously. Observations by the geologist determine the sites where the hydrologists and geochemists focus their attention. Also, the latter may find an important site (e.g. a spring, a site where water flow or chemistry changes abruptly, etc.) to which the geologist should focus attention. There must be continuous interaction among all members of the field party. This interdisciplinary and iterative approach carries through all phases of the study. According to Deaton & Winebrake (2000), complex environmental systems are more readily understood by considering the whole system first, then focusing on details. Thus, the areas selected for detailed study are chosen within the context of the larger system. One strategy for conducting field work, which fits within this approach, is to use geologic parameters (bounding faults, lithologic terranes, etc.) to delimit a study area, then determine hydrologic boundaries from topographic maps. If possible, the field party should walk the length of surface-water drainages, monitor flow, some chemical parameters such as conductivity or pH, and pay attention to rock outcrops or rocks in the streambed. If changes in any of these properties are observed, appropriate samples should be collected. An example is given in the next section. From the point of view of project management, each item in Figure 1 represents a person’s expertise, a facility, or piece of equipment. Thus, Figure 1 may prove to be a useful planning and budgeting tool.

trated waters were found in the north. This separation is readily apparent at a sample density of 2 km-2, but still apparent at sample densities slightly below 1 km-2.This trend is likely due to lithochemical variations in the Osgood intrusive rocks.

4 CASE STUDY ILLUSRATING MULTISCALE BEHAVIOR Figure 2 shows an area in the Osgood Mountains of north-central Nevada, USA, where our geologic, hydrologic and geochemical studies identified multiple scales of variability. The following discussion is keyed to the locations marked A - E on the map.

4.1 Variations at 10’s of kilometers The Osgood intrusive rocks were emplaced in local extensional zones created by strike-slip displacement on the NW-trending faults shown on the map. A third NW fault is inferred from poor exposures at the surface along the SW margin of the southern intrusive body. These large-scale geologic structures resulted from regional tectonic stresses. The fracture network that resulted from those stresses provides the pathways for present-day ground-water flow. Locations of surface-water samples collected from the Osgood Mountains are shown by the round symbols in Figure 2. More dilute waters were generally found in the south, and somewhat more concen-

Figure 2. Map of study area showing several scale-dependent phenomena. Shaded area shows outcrop of granodiorite intrusions; patterned areas show zones of sulfide alteration. Heavy dashed lines indicate right-lateral strike slip faults that controlled emplacement of intrusions (geologic map base from Hotz & Willden 1964; alteration zones from Neuerberg, 1966).

4.2 Variations at 100’s to 1000’s of meters Between points A & B along Granite Creek, changes were observed in the chemistry of surface water. Upstream, near point A, the conductivity values were less than or equal to 210 ps. At point B the conductivity increased to 240 ps and flow increased by more than 5 times. The creek follows an alteration zone, which is defined by an abundance of sul-

223

fide minerals in the rocks (Neuerberg 1966). No tributaries to Granite Creek exist other than those shown on the figure, so any chemical changes from A to B must be explained by mixing of the upstream samples at A with ground-water discharge to produce the observed chemistry at B. The small tributaries near A were more dilute than the main stream at A, and so do not account for the increase in solute load from A to B. Therefore, weathering of the altered rocks in the ground-water environment more than doubled the concentrations of C1, B, K, Ba, Mg, Na, Sr, and Mn. Ground water discharging along this several-km reach of stream bore the chemical signature of the rock alteration. The spring sample collected at point E is unusual from a geomorphological point of view. The spring is located on a ridge crest, rather than in a valley. The localization of the spring along the ridge is controlled by a zone of EW-trending fractures in the rock that extend to the west for at least several km. Although these fractures are obviously hydraulically conductive, their connection to crosscutting fractures in the area must be limited, or else the spring would not be found on a ridge top. 4.3 Variations at 1's to 10's ofmeters The fault at point C is one of the principal bounding faults for the Osgood intrusive rocks. In today's stress regime, it is also an hydraulically conductive feature. The fault is regionally extensive (many km), but the hydrologic effects on the creek are localized within a very narrow zone. As the creek crossed the fault, flow increased by more than a factor of 30, and conductivity decreased from more than 300 ps to about 250 ps. There were two springs at point D, approximately 20 meters apart from each other. Despite their proximity to each other, one had a conductivity of 120 ps, while the other had a conductivity of 280 ps. Because of the thin soil cover, ground-water flow is predominantly in bedrock, and the locations of these springs are structurally controlled. Therefore, the difference in water chemistry is attributed to either differences in residence times of the spring waters in the ground, or to local variations in lithochemistry. 4.4 Summary of scale-dependent phenomena The average sample density in the Osgood Mountains study area was approximately 2 km-*. Many of these samples were collected while walking along the drainages and monitoring conductivity and temperature, and measuring hydraulic heads of ground water beneath streambeds (Wanty & Winter 2000). At the same time, the geologists were nearby observing fracture orientation, density, and offsets (if any). The continuous interaction between geologists and chemists led to many of the observations de-

scribed in this section. With little prior knowledge of the hydrology of the Osgood Mountains, we collected several important pieces of data that will help unravel the hydrology and chemistry of ground and surface waters in the region. Many of the features described in this paper would have been missed without geologic context or perhaps with a lesser sample density. It should be noted, however, that the sample density was not predetermined. Rather, samples were collected based on observations of geologic, hydrologic, and geochemical parameters as field work progressed.

5 CONCLUSIONS Consideration of scale-dependent properties and processes should be an integral part of the planning and execution of all water-rock interaction studies. With this approach, appropriate sample densities may be chosen, and appropriate chemical or physical parameters to measure also can be determined. Questions as to whether the results of a study are truly representative of the system are also best answered in the context of scale dependency. Although not discussed in this paper, temporal scales of variation are equally as important. Environmental systems can vary hourly, daily, seasonally, annually, decadally, etc., so temporal scales of variation should be considered. In dry climate regions, especially, seasonal variations in precipitation may lead to dramatic variations in system hydrology and chemistry. Temporal scales of variation also might be important in problems involving long-term climate change, radioactive waste disposal, and resource assessment, to name a few. REFERENCES Berger, B.R., Wanty, R.B., & Tuttle, M.L., this volume. Deaton, M.L. & Winebrake, J.J. 2000. Dynamic Modeling of Environmental Systems. Springer, New York. Hotz, P.E. & Willden, R. 1964. Geology and mineral deposits of the Osgood Mountains Quadrangle, Humboldt County, Nevada. US Geological Survey Professional Paper 43 1. Hsieh, P., 1998. Scale effects in fluid flow through fractured geologic media, in Scale Dependence and Scale Invariance in Hydrology, Sposito, G., ed., Cambridge University Press, Cambridge. p. 335-353. Miesch, A.T., 1976, Geochemical survey of Missouri--methods of sampling, laboratory analysis, and statistical reduction of data. US Geological Survey Professional Paper 954-A. Neuerberg, G.J. 1966. Distribution of selected accessory minerals in the Osgood Mountains stock, Humboldt county, Nevada. US Geological Survey Miscellaneous Geologic Investigations Map 1-47 1. Wanty, R.B. & Winter, T.C. 2000. A simple device for measuring differences in hydraulic head between surface water and shallow ground water. US Geological Survey Fact Sheet FS-077-00.

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Thermodynamics, kinetics and experimental geochemistry

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Equation of state for aqueous non electrolytes Nikolai N.Akinfiev Chemistry Department, Moscow State Geological Prospecting Academy, Moscow, Russia

ABSTRACT: A new approach for describing thermodynamic properties of aqueous non electrolytes at infinite dilution (chemical potential, entropy, molar volume, and apparent molar heat capacity) in a wide T-P range (0 - 600°C, 1 - 2000 bar) is proposed. It is based on the equation of state for the solvent (H20) given by Hill (I 990) and requires only three empirical parameters that are independent on temperature and pressure and can be estimated from the known standard state properties of the solute. The proposed approach could also be used to estimate the concentration dependencies of the non-electrolytes' thermodynamic properties. 1 INTRODUCTION Proper thermodynamic description of neutral (uncharged) aqueous species in a wide pressure - temperature range is an important task with many practical applications in geochemistry and technology. Nevertheless, such a description is now far less accurate than those of ionic species. For instance, the well known HKF equation of state given by Tanger & Helgeson (1988) gives excellent predictions for thermodynamic properties of ions in the wide range of temperatures (0-600°C) and pressures (1-5000 bar). However, extension of the model to neutral aqueous species (Shock et al., 1989) does not provide the correct behaviour for nonelectrolytes solutes in the nearcritical and supercritical regions. As reported by Plyasunov (1 99 1), calculation of chemical potentials of the dissolved gases based on the HKF equation diverges from experimental observations at T > 4OO0C, and the discrepancy grows as the temperature increases. In our previous study (Akinfiev 1997) the Redlich & Kwong (1949) equation of state was extended for description of binary H20 - gas systems in the low temperature (down to 1OOOC) and infinite dilution conditions. The approach was essentially based on the usage of the "exact" multiparametric unified equation of state for H2O given by Hill (1990). Now expanding on this method, we propose a simpler and more accurate low-parametric approach applicable to a wide temperature - pressure range that encomposses the H20 critical region.

2 DEVELOPMENT OF A NEW EQUATION OF STATE 2.1 Basic relations Chemical potential of the dissolved species at infinite dilution p2 can be derived using an approach similar to the virial equation adopted for non ideal gases (Mason & Spurling 1970): ~ 2 ( P . T ) = ~ 2 , R ( T ) + R T ( - I n )+in.fi ( N , , +2Pl . A B ) (1)

Here plg(T) stands for chemical potential of the pure gaseous component at temperature T and standard pressure P = 1 bar, R is the gas constant, N,, = 1000/M1 55.51 mol/kg, MI, .fi and p1 stand for molar mass (glmol), fugacity (bar) and density (g/cm3) of the pure solvent (H20) at P-T conditions of interest, respectively, and ALI characterises the difference in the short-range interaction energy between solute and solvent molecules. Applicability of equation (1) was tested using the numerous experimental data on temperature dependencies of Henry's constants kH for the dissolved gases. It was found that eligible k ~ ( 7 Jdescription could be provided with the use of an empirical approximation to AB

where a and b are adjustable parameters of the model. Since the P derivatives of (2) are non negligible, an additional empirical parameter 6 (the socalled scaling factor) is introduced to describe the 227

5, a, and b by the least squares method. Results are given in Table 1. The goodness of the fitting procedure may be appreciated in Figures 1 - 3 - The accuracy of temperature dependencies of V2 and Cp2 characterise the prognostic capacity of the proposed equation of state since the corresponding experimental data were not used in the estimation procedure.

difference between intrinsic volumes of H20 and the dissolved molecule (Plyasunov et al. 2000). The ultimate master equation for p2 can be now written as: ~ 2 ( P , T ) = p 2 , g ( T ) - R T 1 " N H+' (1-c)RTInfi +RTcIn

(3)

RTpl -2AB Here RV = 83.14 is the gas constant expressed in "volume" units (cm3.bar.K-'-mol-'). Differentiating (3) with respect to temperature and pressure, expressions for partial entropy (Sd, heat capacity (Cp2),and volume (V2) of the dissolved component are obtained:

R

Table 1. Adjustable parameters of the proposed equation of state (2,3) for a number of dissolved gaseous species. Gas a B 4 cm'/g cm3.K0.'/g Ar 0.0730 -8.5139 11.9210 C ~ H ~ -0.48 15 - I 8.2 120 19.8058 C6H6 -0.6050 - 1 8.0662 20.98 19 - I I .8462 14.86 15 -0.1 I31 CH4 -0.0850 -8.832 1 CO2 1 1.2684 H2 0.3090 -8.4596 10.830 I H2S -0.2020 - 1 1.4803 12.7158 N2 -0.0320 -1 1.5380 14.6278 0 2 0.0260 -9.7540 12.941 1

4.5 and close to equilibrium. As pH decreased during stage (B) the rate of zeolite dissolution increased, and as a result A I concentration increased, The silicon concentration is probably limited by the solubility of a silica rich phase, with Al/Si stoichiometric ratio lower than that of NaPl (0.6). As a result the N/Si ratio increases during stage (B). Due to the disappearance of the NaOH and the resulting decrease in pH, the solution becomes undersaturated with respect to this silica rich phase, and both this phase and the zeolite are dissolved during stage (C). Therefore, the Al/Si ratio decreases and approaches the stoichiometric ratio of NaPl at the beginning of stage (D). The increase in concentration during stage (C) is due to the increase in dissolution rate as the pH drops. Figure 2 focuses on the change in Al and Si concentrations with time during stage (D). During this stage (D), both A I and Si concentrations slowly decrease with time. However, the Al/Si ratio remains constant throughout this stage and is equal to 0.6, i.e., to the stoichiometric Al/Si ratio of NaP 1. The slope of the change in Al with time (0.09*0.01) is by a factor of 0.6 shallower than that of Si (0.15*0.02). This congruent dissolution indicates that the change in concentration during stage (D) is dominated by zeolite dissolution. Due to the slow change in concentrations with time, the presumption can be made that at each time the concentration is constant within error, and therefore it is possible to calculate the dissolution rate using equation (1). Figure 2 shows that although the Al and silicon concentrations change with time, the calculated dissolution rates are constant throughout this stage. The decrease in Al and Si with time is a result of the fast change in zeolite mass during this stage (Fig. lb). The observation that the dissolution rate remains constant throughout the stage, even though a significant amount of zeolite was dissolved, indicates that the (specific) reactive surface area did not decrease during this stage. An alternative explanation may be that the decrease in reactive surface area was balanced by an increase in dissolution rate due to the decrease in pH. During stage (E) both Al and Si concentrations decreased faster than in stage (D) and the Al / Si ratio became lower with time. Due to the dissolution of zeolite during previous stages, the change in concentration at stage (E) is dominated by dissolution of 249

Figure 2. Changes in AI and Si concentrations and in dissolution rate during the stoichionietric period (Stage D) of a dissolution experiment at 25°C.

accessory phases that dissolved slower than zeolite. Only quartz and mullite were identified in X-ray diffraction patterns of the powder recovered after the experiment (Fig. 3), indicating that NaP 1 has been totally dissolved. Figure l b shows the estimated accumulated loss of mass during the experiment, that was calculated based on the outflux of aluminurn and silicon, and assuming that only NaPl is dissolving. This estimation shows that about 40% of the initial mass was lost before the beginning of stage (E). The complete dissolution of NaPl may imply that during different stages in the experiment, different population of zeolite was dissolved, i.e., that the zeolite grains that dissolved early in the experiment have different surface properties and therefore dissolved faster than grains that dissolved towards the end of stage (D). Unfortunately, due to the complete

Figure 3. XRD patterns of the raw sarnpIe and afterdissolution experiments showing the major reflections of NaPl (only in the raw material), niullite and quartz (in both samples).

dissolution of the zeolite, measurement of the final surface area (and any other property of the remaining powder) becomes meaningless. Consequently, the calculated dissolution rates were normalized to the product of the initial surface area (18.1 rn2 g-')

and the estimated remaining mass. The average dissolution rate obtained at steady state (stage D) is 5.5*0.4~10-'~ mol m-2s-'. 3.2 Dissolution of raw Nap1 at 50 "C The variations of pH and Al and Si output concentrations as a function of time at 50°C (Fig. 4a) are similar to that at 25°C (Fig. la). Due to the higher temperature, the experiment approaches stoichiometric dissolution aRer less than 200 h (Fig. 4b). Figure 5 focuses on the change in Al and Si concentrations with time during the stoichiometric period. The shaded area indicates an intermediate non-stoichiometric disruption. We do not have a good explanation for this disruption. As for the 25" experiment, the calculated dissolution rates are constant with time, regardless of the changes in pH and in Al and Si concentrations. The average dissolution rate obtained at 50°C is 8.9&0.3~10-'~ mol m'2 s-'. 4 SUMMARY AND CONCLUSIONS Preliminary dissolution experiments of raw NaP 1 zeolite were conducted at 25 and 50 "C and input pH 3. Even though the sample is a mixture of crystalline and amorphous phases, a congruent dissolution of Nap1 is observed during significant periods of the experiment. During these stoichiometric periods it is possible to determine the zeolite dissolution rate.

Figure 4. Dissolution of raw NaPl sample at 50 "C. Initial mass is 0.47 g; the input pH solution is 3. Variation of (a) Si and A1 output concentrations (ClM) and output pH and (b) AIoutput / Siout(,ut ratio and accumulated loss of mass (%) versus time. The shaded area corresponds to a stirring stage, which is not discussed in the present communication.

REFERENCES Querol, x.,AlasheY, A., Lk~ez-Soler, A., Plan% F., Andrks, J.M., Pedro-Ferrer, R..J. 8~RuiZ c. 1997. A fast method for recycling fly ash: microwave-assistedzeolite synthesis. Environmental Science & Technology 3 1; 2527-2533 Ragnarsdottir, K.V. 1993. Dissolution kinetics of heulandite at pH 2-12 and 25 "C. Geochimica and Cosmochirnica Acta 57: 2439-2449. Saaltink, M.W., Ayora C. & Carreras, J. 1998. A mathematical formulation for reactive transport that eliminates mineral concentration. Water Resources Research 34: 1649-1656.

Figure 5. Changes in AI and Si output concentrations and in dissolution rate during the stoichiometric period of a dissolution experiment at 50°C. The shaded area indicates an intermediate non-stoichiometric disruption.

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Wafer-Rock Interaction 2001, Cidu (ed.), 02001 Swefs & Zeitlinger, Lisse, ISBN 90 2651 824 2

Dissolution rate of apophyllite. The effects of pH and implications for underground water storage L .C .Cav6 University of Cape Town, Cape Town, South Africa

M .V.Fey University of Stellenbosch, Stellenbosch, South AJi-ica

D .K.Nordstrom U S . Geological Survey, Boulder, Colorado, U S A .

ABSTRACT: The dissolution rate of apophyllite, KCa4Si8020(F,OH).SH20, was determined in solutions ranging from pH 2 to 10 at 25°C. The samples used were from South Africa, Mexico, India and the USA and had a large range of OH-F and N b - K substitution. Dissolution is non-stoichiometric in acid solution, with release of Ca and F occurring at 4 and 3 times respectively the rate of Si at pH 2. SEM images confirm preferential acid leaching of the interlayer regions. Dissolution stoichiometry arproaches congruency in the neutral pH range. The average rate of fluorapophyllite dissolution is 3.5 x 10- moVm2/s at pH 2 decreasing to 3 x 10 mol/m2/s at pH 4 and 2 x 10 mol/m2/s at pH 7 and 10. Dissolution rates above pH 4 are almost 3 times faster for samples with a high proportion of OH substitution for F. The slow rate of fluorapophyllite dissolution is advantageous for an underground water storage scheme in South Africa. Ammonian fluorapophyllite is a major secondary phase in a mineralised breccia pipe, intended as a reservoir for drinking water supplies. Concern that the injection of oxygenated water could lead to rapid dissolution of fluorapophyllite and high concentrations of dissolved fluoride, appears to be unfounded.

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1 INTRODUCTION

2 MATERIALS AND METHODS

The apophyllite group, K C ~ ~ S ~ ~ O ~ O ( F , O H ) . S A H ~sequential O, batch method (Amrhein & Suarez 1992) was used for the dissolution rate experiments. consists of a series of hydrous sheet silicates, usually Four apophyllite samples (Table 1) were crushed, found in zones of low temperature hydrothermal alground and sieved to between 53 and 106 pm. The teration. Our research into the rate of dissolution of apophyllites was stimulated by a feasibility study for samples were sonicated in methanol for 15 minutes to dislodge ultrafine particles, washed thoroughly artificial groundwater recharge in South Africa. with double distilled water and dried at 102°C. SpeAmmonian fluorapophyllite is a major mineral cific surface area, determined by multipoint nitrogen phase in a hydrothermally altered breccia pipe in the BET analysis, ranged from 0.28 to 0.34 m2/g. PowKaroo region. The high permeability of the fractured der X-ray diffraction analysis was used to confirm rocks in the breccia pipe make this site a promising target for underground water storage for a nearby mineral purity. Apophyllite samples were weighed out in 1.3 g town, but the presence of a fluoride-bearing mineral portions into sixteen 50 ml polypropylene reaction poses a risk for fluoride contamination of the stored vessels, four for each sample. Stock solutions of pH water. Introducing oxygenated, neutral pH, surface 2, 4, 7 and 10 were prepared from double-distilled water into a reducing, alkaline groundwater is likely to shift the natural equilibrium conditions in the water by adding HC1 or NaOH. Aliquots of 40 ml of these solutions were added to the reaction vessels, so breccia pipe, with the result that oxidation of sulthat each apophyllite sample was reacted under phide minerals, NH4' and organic carbon in the subsurface could lower pH and impair the water quality fourdifferent pH conditions. The samples were agi(Cave 1999). An understanding of the dissolution tated at 60 cpm on a reciprocating shaker immersed rate of apophyllite under a range of pH conditions in a water bath, which was maintained at 25°C for will help assess the fluoride contamination risk, esthe duration of the experiment. At one-week interpecially since the mineral solubility is W o w n . vals, 35 ml of the supernatant liquid was carefully

251

withdrawn for analysis, after allowing the solid to settle, and replaced with fresh pH-adjusted solutions.

with no detectable K, Ca or F remaining. The Si/Ca and Si/F ratios at pH 4,7 and 10 approached mineral stoichiometry. (Fig. 1).

Table 1. Composition of apophyllite samples used in dissolution rate experiments. KF

JL PT GJ

Origin Calvinia, South Africa Jalgaon, India Paterson, NJ, USA Guanajuato, Mexico

Composition KOs(NH4)osCa,&8020F.8H20 KCQSi802o Fo s(OH)o2.8H20 KCa4Si8020F0 s(OH)os.8H20 KOs(NH4)o 2Ca4Si8020F0 s(OH)os.8H20

Samples were returned to the shaker and the experiment continued for a total of 737 hours. The solutions were analysed for pH and filtered through 0.2 pm filters for analysis of Ca, Si and K by ICP-AES and F by ion selective electrode. H2SO4 was added as a preservative to prevent degradation of NH4, determined by ion chromatography. A second experiment was conducted following the same method for a total time of 649 hours, using only the apophyllite sample from South Africa, KF, under a wider range of pH conditions. NaCl was used as a background electrolyte to maintain a uniform ionic strength of 0.01M. Scanning electron microscopy (SEM) and energy-dispersive X-ray analysis (EDX) were used to examine the mineral surface morphology and bulk composition before and after the second dissolution experiment. Dissolution rates with respect to Si, Ca and F were determined by accumulating the number of moles that had been dissolved from the apophyllite mineral at the end of each sequential rinse. The removal of species, caused by periodic changing of the reaction solution, was taken into account in the calculations. Reaction rates were taken as the slope for the linear portion of a plot of moles of product released per unit surface area against time, assuming a zero-order reaction rate with respect to products. 3 RESULTS AND DISCUSSION

3.1 Reaction stoichiometry All apophyllite samples were found to dissolve nonstoichiometrically at pH 2. The amount of F, Ca and NH4 leached from the apophyllite samples was three to four times the equivalent stoichiometric amount of Si released from the mineral. Acid leaching of apophyllite produces a crystalline silica hydrate residue (Frondel 1979, Sogo et al. 1998) by selectively leaching out the Ca, F or OH and K or NH4 ions from between the silicate sheets. SEM-EDX analysis of the sample KF after 4 weeks’ dissolution at pH 2 revealed only Si and 0 in the outer layers of the mineral residue penetrated by the X-ray beam,

Figure la). Release of F and Si in apophyllite dissolution experiments at pH 4 (solid symbols), pH 7 (shaded symbols) and pH 10 (open symbols). b.) Stoichiometric release of Si and Ca in apophyllite dissolution experiments at pH 4, 7 and 10. Dashed lines show the mineral stoichiometry.

The molar volume of apophyllite (Colville & Anderson 1971) and the excess dissolved Ca over Si was used to calculate the thickness of the leached layer that would form at the mineral surface at pH 2, assuming that Ca is uniformly leached out from the surface. The calculated layer thickness ranged from approximately 350 nm after the first week of reaction to 1000 nm at the end of the experiment. The actual depth of leaching should be greater than this, since leaching takes place preferentially along the crystal planes parallel to the silicate sheets. The accumulation of solutes with time displayed linear behaviour at pH 2 for the first three weeks of the experiment. After the fourth rinse, however, Ca, K and NH4 concentrations were lower than anticipated by a zero-order rate law with respect to the reaction products. This result could indicate that the leached layer had become so thick that diffusion of species into and out of the leached layer, rather than surface reactions, had become the rate-limiting step.

252

3.2 Influence of mineral composition Overall dissolution rates for the apophyllite samples, presented in Table 2, were determined from the average rate of accumulation of Si, Ca and F in solution. At pH 2, where dissolution was nonstoichiometric, the rate of Si accumulation was divided by 8 to give the rate of mineral dissolution. The reasoning was that 8 moles of Si are released per mole of apophyllite only when the silicate layers, which form the framework of the mineral, break down. The rate of fluoride release from the South African sample is also reported, because of its relevance to the fluoride risk in the recharge feasibility study. Table 2. Dissolution rate of apophyllite samples (moVm2/s),assuming zero-order reaction kinetics with respect to Si at pH 2 and Si, Ca and F at pH 4, 7 and 10. The reaction rate with respect to F is also reported for the KF sample. Rates reported for KF are the average for the 2 experiments. Sample KF

JL PT GJ F'-KF

pH2 3.71t0.4 x 10-" 3.1 x 10-l' 3.6 x 10-l' 3.4 x 10-l'

pH 4 2.91t0.4 x 10-" 3.8 x 10-" 8.0 x 10-" 9.5 x 10-"

pH 7 2.261t0.02 x 10-" 3.1 x 10-l' 7.5 x IO-" 9.8 x 10-l'

1.3kO.1 109

3.2rt0.4

2.31t0.1 x~O-~

10-l~

Figure 2. Dissolution rate of South Afiican fluorapophyllite sample, KF, as a function of pH.

0.5 (Wollast & Chou 1985), phlogopite and biotite, n 0.4 (Sverdrup 1990) and wollastonite, n = 0.4 (Rimstidt & Dove 1986). This is generally interpreted by transition state theory as a proton-promoted surface reaction involving an activated complex comprising one hydrogen ion and two molecules of the dissolving mineral. Above a transition point, between pH 4 and 5 , the dependence on pH changes to a much lower reaction order, with the dissolution rate reaching a minimum near pH 10. In the neutral pH region the dissolution rate is typically independent of pH for silicate minerals (Drever 1994). Reaction rates for fluorapophyllite dissolution increased slightly with increasing pH in the alkaline region above pH 10. The slope of 0.13 shown in the alkaline region on Figure 2 is lower than the slope of 0.4 reported by Brady and Walther (1992) for silica, kaolinite, albite and forsterite at 25°C. Dissolution reactions in this region are dominated by deprotonation of silica tetrahedra on the mineral surface.

pH 10 1.95&0.02 x 10-" 2.8 x 10-l' 7.0 x 10-" 9.8 x 10-"

I

=

2.11t0.3 IO-~I

Apophyllite dissolution at pH 2 is rapid in comparison to the higher pH values, and is similar for all four mineral samples. At pH 4 and higher, dissolution rates are less dependent on pH, but vary with mineral composition. The low-fluorine samples, GJ and PT, dissolve at rates up to three times faster than the fluorapophyllites, KF and JL. Substitution of NH4 for K does not seem to have as significant an effect as OH in destabilizing the crystal lattice and increasing dissolution rates. 3.3 pH dependence of dissolution rates Figure 2 illustrates the relationship between solution pH and the rate of dissolution of the KF apophyllite sample from South Africa, investigated in the second experiment. Far from equilibrium, the dissolution rates of most silicate minerals show a pH-rate relationship similar to Figure 2 (Drever 1994). In the acid range, between pH 2 and pH 4, the pHrate relationship for the dissolution of fluorapophyllite may be described by the expression: RateH = kH (aH+)n i.e. log RateH = log kH - n.pH (1) Linear regression through our data points in the acid region gives a value of 0.52 for n, the reaction order with respect to hydrogen ion activity. This is consistent with other studies of silicate hydrolysis, which often show reaction orders for hydrogen near 0.5 in the acid range. Examples include albite, n =

3.4 Limitations of batch experiments Batch experiments are a simple and effective method of measuring reaction rates, especially when these are slow, as is the case for silicate dissolution. Some limitations of batch tests, however, include: - Reaction rates cannot be measured directly, so concentration time data is fitted to the integrated form of a presumed rate law. - Accumulation of solutes in the reaction vessels can lead to supersaturation with respect to secondary mineral products and the concentration time pattern changes each time a new phase forms (Rimstidt & Dove 1986). The absence of A1 in apophyllite reduces the risk of precipitation in comparison with aluminosilicates. Changing the solution in the sequential rinse experiments also helps to prevent the build up of high solute con-

253

Table 3. Comparison of dissolution rates and experimental conditions for other phyllosilicates and low grade metamorphic minerals. Mineral

Rate

pH

muscovite kaolinite prehnite epidote chrysotile

movm'/s 2.4 x 10-" 9.8 x lO-I4 1.3 x 10-l' 3.0 x 10"' 2 x 10-l'

4.6-5.1 6.93 1.4 1.4 8

-

Ionic strength mol/kg variable 0.05 0.05 0.05 0.1

BET surface area m'/g 5.84 11.2 0.381 0.328 48.5

centrations. No overgrowths or precipitates were observed by SEM investigation and calculated saturation indices for common insoluble solids were generally below zero. Calcite supersaturation was calculated for solutions above pH 9 and may have influenced the rate reported for pH 10. The saturation status of apophyllite in the experiments is not constant and cannot be calculated, due to the lack of data for the mineral solubility. Our experiments, therefore, do not take into account the dependence of the reaction rate on chemical affinity.

4 CONCLUSIONS Dissolution of apophyllite takes place at rates that are comparable to those reported for other silicate minerals at 25°C (Table 3). Dissolution rates in the neutral pH range are faster than those of other phyllosilicates e.g. muscovite and kaolinite, possibly because of increased strain in the four-membered silicate rings, when compared with the hexagonal arrangement of the other layer silicates. It is also possible that true steady state rates were not reached in our batch experiments. Apophyllite dissolution rates were also found to be faster than those reported for some other low grade metamorphic minerals, such as prehnite, epidote and chrysotile, at comparable pH values. Hydroxyapophyllite will probably dissolve faster than fluor-apophyllite, but substitution of ammonium in the mineral does not appear to cause a significant increase in dissolution rate. Field dissolution rates are often considerably slower than those predicted in the laboratory (Brantley 1992). Furthermore, fast dissolution and non-stoichiometric leaching of fluoride from apophyllite are only observed under acid conditions below pH 4. Hence, in the neutral to alkaline pH conditions that are anticipated during underground water storage in the breccia pipe, excessive release of fluoride by apophyllite dissolution is unlikely. ACKNOWLEDGEMENTS The U.S. Geological Survey Volunteer Program made it possible for this work to be largely carried out in the laboratories of the USGS, Boulder, Colorado. Lisa Cave is grateful to the University of Cape Town funding her travel to the U.S.A. and

Chemical affinity kcal/mol not reported notreported 62 80 not reported

Mineral/ solution ratio Irjl 25 2 10- 15 10- 15 0.9, 5, 10

Reference Lin & Clemency 1981 Carroll & Walther 1990 Rose 1991 Rose 1991 Bales & Morgan 1985

to Blaine McCleskey and JoAnn Holloway of the USGS for help with the analysis of rock and water samples. David Rutherford (USGS) and Susanna Vasic (UCT) carried out the surface area measurements and Dane Gerneke (UCT) assisted with SEM-EDX analyses. Additional samples of apophyllite were provided by Dan Kile (USGS) and Gordon Brown (Stanford University). Kathy Nagy (University of Colorado) and Alex Blum (USGS) contributed helpful discussions on mineral dissolution.

REFERENCES Amrhein, C. & D.L. Suarez 1992. Some factors affecting the dissolution kinetics of anorthite at 25°C. Geochim. Cosmochim. Acta. 56: 1815- 1826. Bales, R.C. & J.J. Morgan. 1985. Dissolution kinetics of chrysotile at pH 7 to 10. Geochim. Cosmochim. Acta. 49:228 1-2288. Brady, P.V. & J.V. Walther. 1992. Surface chemistry and silicate dissolution at elevated temperatures. Am. Jnl Sci. 292:639-658. Brantley, S.L. 1992, Kinetics of dissolution and precipitation Experimental and field results. In Kharaka & Maest (eds). Water-Rock Interaction - Proc. 7 ~Int. . Symp. Water-Rock Interaction:3-6, R0tterdam:Balkema. Carroll, S.A. & J.V. Walther. 1990. Kaolinite dissolution rates at 25"C, 60°C and 80°C. Am. JnI Sci. 290:797-810. Cave, L.C. 1999. Geochemistry of art$cial groundwater recharge into the Kopoasfontein breccia pipe near Calvinia, Karoo. MSc thesis. Cape Town: University of Cape Town. Colville, A.A. & C.P. Anderson 1971. Refinement of the crystal structure of apophyllite. I. X-ray diffraction and physical properties. Amer. Mineral. 56:1222-1233. Drever, J.I. 1994. The effect of land plants on weathering rates of silicate minerals. Geochim. Cosmochim. Acta. 58:23252332. Frondel, C. 1979. Crystalline silica hydrates from leached silicates. Amer. Mineral. 64:799-804. Lin, F-Ch. & C.V. Clemency 1981. The kinetics of dissolution of muscovites at 25 "C and 1 atm CO2 partial pressure. Geochim. Cosmochim. Acta. 45:57 1-576. Rimstidt, J.D. & P.M. Dove. 1986. MineraVsolution reaction rates in a mixed flow reactor: Wollastonite hydrolysis. Geochim. Cosmochim. Acta. 50:2509-25 16. Rose, N.M. 1991. Dissolution rates of prehnite, epodote, and albite. Geochim. Cosmochim. Acta. 55:3273-3286. Sogo, Y., T. Kumazawa & A. Yamazaki. 1998. [Methylene blue adsorption for layered alumino-silicate prepared from apophyllite]. Abstract in English. JnI C l q Sci. Soc. Japan. 38(2):47-53. Sverdrup, H.U. 1990. The kinetics of base cation release due to chemical weathering. Lund, Sweden: Lund University Press. Wollast, R. & L. Chou. 1985. Kinetic study of the dissolution of albite with a continuous flow-through fluidized bed reactor. In J.I. Drever (ed.), The chemistry of weathering. NATO Adv. Study Inst. Ser., Ser.C. 149:75-96. Hingham: D.Reide1.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Experimental Study on Mixture Corrosion Efiects in Littoral Karst Area, coastal Liaodong Peninsula, China Honghan Chen & Shengzhang Zou & Erping Bi China University of Geosciences, Beijing, 100083, P R . China

ABSTRACT: Seven groups of experiments on the mixture corrosion effect have been carried out in H20-CO2 closed system using rocks from the Daweijia littoral karst system which shows an unique dynamics process of karstification. Basic conclusions as follows: 1.the basic law of corrosion process in transitional zone of seawater-freshwater in littoral karst areas is identical with that of in the fresh water, 2.the influence of carbonate rock structure on its specific corrosion ratio is relative significant, 3.the specific corrosion ratio of the carbonate rock in seawater-freshwater transitional zone is faster than that in freshwater or seawater.

1 INTRODUCTION

2 THEORY

The mixture corrosion effects were found firstly by Buneyew in 1932(Bogli 1980) and were explained by curves of solubility of CaCO3 in water solution in 1964(Bogli 1980). Since then, this explanatory method was widely accepted and applied. But in the case of high ionic content as in seawater, Bogli mixture corrosion theory is no longer applicable (Qian & Hu 1996). Mixture corrosion effects are the most important chemical kinetics process in littoral karst areas, but these effects were rarely studied in China. In some countries, the studies have been carried out since 1960s (Bogli 1960, Back et al. 1986, Stoessell et al. 1989, Morse et al. 1997), however these formers' studies were carried out in artificial synthetic solutions (Bogli 1980) or in seawater mixed with artificial synthetic solutions (Morse et al. 1997), and only a few factors, such as T (temperature), S and Pc02 (Morse & He1993) or T and Mg : Ca (Morse et al. 1997), were considered respectively. So there are some discrepancies among the former conclusions. Furthermore, we must consider that compact Palaeozoic and Proterozoic carbonate rock widely along the coast of Liaodong peninsula, show higher hardness and lower porosity. And the neotectonism is mainly represented an upward movement in this area. So the kinetics process of the karst formation is different from the other littoral karst areas in the world.

In general, the corrosion capacity of dissolving CaC03 will be increased when seawater is mixed with fresh water, although the fresh water is saturated with CaC03. This mainly results from the ionic strength effect, which causes an increase in ionic complex and ion-pair formation and a decrease of ionic activity. At the same time, with the increasing of solution salinity, Ca2' activity coefficient would decrease too. Thus, the mixture solution will have stronger corrosive capacity in comparison to freshwater or seawater. In the seawater-freshwater transitional zone of littoral karst aquifer, karstification of carbonate rock is usually more relevant than that in freshwater aquifer. This process was firstly recognized by Mandel (Mandel 1964), and confirmed by Schmorak and Mercado (Schmorak & Mercado 1969). However, because of complex ionic sorts and high ionic concentration, mixture corrosion effect in seawater-freshwater is more complex and affected by other factors (complex ion, ion-pair and commonion effects, foreign ions, ionic strength and temperature, etc.). In 1987, Buhmann et al. studied the effect of foreign ions on the dissolution kinetics of calcite in the pure system CaC03-H20-C02 for plane water films. The aim of this study is to build the dynamical model of carbonate rock dissolution and precipitation on the basis of field observation on site and experimental study of carbonate rock dissolution and precipitation in transitional zone of seawater and fresh water. Various factors will be considered in the experiments.

This study intends to consider the mixture corrosion effects in littoral karst areas in China, and explore a new way to solve the sea-salt water intrusion in littoral karst areas.

255

3 EXPERIMENT 3. I Experimental set-up

Daweijia (in Dalian City) has been taken as the typical area. Based on detailed field investigation, the kinetic laboratory experiment of dissolution and precipitation of the carbonate rock was carried out, with local seawater and rock. In order to find out the specific corrosion ratio of carbonate rocks in seawaterfresh water transitional zone, four groups static experiments have been carried out in closed C02-H20 equilibrium system. In these experiments, rock specimens (3cm long, lcm wide, and lcm thick) were shaped, and their surface polished. Bottles of 125ml were filled with 125ml solution (four groups of experimental solution are: (1) seawater, (2) freshwater, (3) 20%seawater-8O%freshwater, and (4) 40%seawter-60%freshwater), ( 5 ) 50%0seawater5O%freshwater, (6) 60Seawater-40%freshwater, (7) 80%seawter-20%freshwater. The experimented samples refer to: algalgranular-micrite-dolomite limestone (No. l), micritic limestone (N0.2, NO.^), granular-crystalline limestone p 0 . 4 , NO.^), bioclastic limestone NO.^), arenitic-microspar limestone NO.^), micriticgranular-calcian dolomite (N0.8). A pure marble (No. 18) was, at the same time, carried out for comparison. These rock specimens were hung at a copper wire and soaked in the experimental solution. Each solution was added with CO2 (in order to shorten the experimental period) to pH=4.77 before these experiments were carried out. During the experiment time, pH value, temperature and conductivity were measured continuously. The temperature was controlled between 16.5-1 7.5"C. The experiment was lasting about 288hours. Ionic concentrations were measured at the beginning and at the end of each experiment.

4 DISCUSSION AND CONCLUSIONS

3.2 Experimental results

The results (Table 1) indicate that corrosive content of each specimen is different in different solution. __

Total corrosion content (mg)

Unit area corrosion ratio

The solution showing the largest corrosion content is the mixture solution containing 80% seawater, the second one is the mixture solution which containing 60%seawater, and the smallest one is pure seawater. This behavior agreed with the mixture corrosion theoretical prediction. In this paper, 3 typical lithologies (No.1, No.4, No.8) and pure marble (No. 18) have been selected to illustrate mixture corrosion curves at different times (Figures 1-3). The curves in Figures 1 show that the increasing rate of pH value is the fastest one in seawater, the second one in freshwater, and the slowest one in solution mixed with 80% seawater. That is to say, acid environment will be lasting longest in mixture solution mixed with 80% seawater, so the corrosion content will be the largest in it, too. We will continue to study this phenomenon in our next work. At the same time, mixture corrosion effect is largely influenced by lithologic characteristics. The rocks of different lithology have different specific corrosion ratio in the same solution in Figures 2-3. It was found that the rate of No.8 is the slowest one, and as it is shown No.1 is the fastest one. This can be also found in Table 1. It is worth attending that when pH values increases to about 5.5 or 6.5, the specific corrosion ratios of all samples change. When pH is lower than 5.5, the rate is the very fast; when pH is higher than 6.5, the rate became slower and slower. How to explain the phenomenon is very important. It will contribute us to solve the problem of mixture corrosion effects. Of course, temperature and Pc02 are also the main influencing factors. During the course of experiment, as Morse et al. said (Morse & He1993), with the higher Pc02 decrease gradually, the mixture corrosion rate decreases too.

It is well known that mixture corrosion effect is mainly influenced by composition of mixture solu-

Table 1. Total corrosion content (mg) & corrosion ratio (unit area) of all samples Sample number No.1 No.2 No.3 No.4 No.5 No.6 seawater 60.40 56.20 50.10 43.90 50.50 60.80 53.10 54.70 53.00 59.50 63.40 61.50 freshwater 68.90 70.00 56.10 66.50 20%-seawater 76.30 71.80 97.40 73.20 78.80 40%-seawater 72.80 87.20 83.70 50%-seawater 83.90 73.90 72.60 74.70 85.10 82.30 78.60 73.30 75.70 60%-seawater 84.00 81.70 81.60 84.50 89.80 85.00 89.30 82.60 85.60 SO%-seawater seawater 1.2955 1.2387 1.0502 0.9223 1.0514 1.3083 freshwater 1.1113 1.1524 1.1512 1.2653 1.3362 1.2840 20%-seawater 1.4344 1.481 1 1.1984 1 SO3 1 1S619 1S360 40%-seawater 1.5249 2.1042 1.5290 1.6535 1.8016 1.7740 1.6643 1.4782 1.4498 1.5183 1.7801 1.6734 50%-seawater 60%-seawater 1.7444 1.5497 1.6745 1.4831 1.61 1 1 1.6741 SO%-seawater 1.7739 1.6721 1.7825 1.6792 1.6212 1.6607

* Using corrosion ratio of fresh water as standard 256

No.7

No.8

No.18

45.60 70.60 81.10 83.00 71.90 84.10 99.30 0.9187 1.4348 1.6952 1.6995 1.4460 1.7078 2.0316

18.50 18.80 30.00 33.50 22.90 28.40 27.50 0.4012 0.4042 0.6022 0.7085 0.4476 0.5622 0.5394

51.20 46.70 60.50 62.00 61.50 64.60 66.70 1.0571 1.0000 1.0668 1.2800 1.2365 1.2962 1.3162

tion. Although our researches have not finished at last, initial conclusions can be drawn as follows: (1) The basic law of corrosion process between the seawater-fresh water transitional zone in littoral karst areas and the fresh water in inland karst areas are identical, i.e., the rocks with different lithologic characteristic have different specific corrosion rate (the specific corrosion rate of pure limestone faster than that of dolomite). (2) Carbonate rock structure affects significantly the corrosion rate of the rock, e.g., the specific corrosion rate of coarse crystalline limestone is faster than that of fine crystalline limestone. (3) The specific corrosion rate of the carbonate rock in seawater-freshwater transitional zone is faster than that in fresh water or seawater. On the basis of above-mentioned static experiments, four groups of dynamic mixture corrosion experiments are going to succeed in the open Co2H20 non-equilibrium system. The dissolution and precipitation rates of the carbonate rocks in different concentration of solution and the different carbonate rocks in the same solution will be calculated respectively by inverse geochemical modeling after these experiments have been done. Our final aim is modeling hydrochemical kinetics process in littoral karst, and to evaluate the effects of Ca , Mg” , Na’ , HCO; ,C1- ,SO:- ,PcOz and temperature on the dissolution and precipitation rates of the carbonate rocks.

Figure 2. Time-pH curves of specimens in 40%-seawater

*+

Figure 3. Time-pH curves of specimens in 40%-seawater

REFERENCES

Figure 1. Time-pH curves of No.4 specimen

ACKNOWLEDGEMENTS This research was supported by the National Natural Science Foundation, P.R.C.(49832005). The authors thank Prof. Zhong Zuoxin, Jiang Jingchen, and Tang Minggao for the guide in this study, and Dr. Yu Yongbo for the test of ionic analysis during the experiments.

Bogli, A. 1980. Karst Hydrology and Physical Speleology. Springer-Verlag. Qian, H. & Hu, J.G. 1996. Bogli mixture corrosion theory and problems encountered in its practical application. Carsologica Sinica 15( 1): 367-375. Bogli, A. 1960. Solution limestone and karren formation translated form Zeitschr. Geomorphologic Supplement (2): 4-21. Back, W. & Hanshaw, B.B. & Herman, J.S. & Nicholas, J. 1986. Differential dissolution of a pleistocence reef in the groundwater mixture zone of coastal Yucatan, Mexico. Geology. V. 14: 137-140. Stoessell, R.K. & Ward, W.C. &Ford, B.H. & Schuffert, J.D. 1989. Water chemistry and CaC03 dissolution in the saline part of an open-flow mixture zone, coastal Yucatan Peninsula, Mexico. Geological Society of America Bulletin. v. 101: 159-169. Morse, J.W. &. Wang, Q.W. & Tsio, M.Y. 1997. Influences of temperature and Mg:Ca ration on ration on CaC03 precipitates from seawater. Geology. v. 25: 85-87. Morse, J.W., & He, S. 1993. Influences of T, S and Pco2 on the pseudohomogeneous nucleation of calcium carbonate

257

from seawater: Implications for whiting formation. Marine Chemistry. v .4 1: 291 -298. Mandel, S. 1964. The mechanism of sea-water intrusion into calcareous aquifers. General Assembly of Brekeley. Int Assoc. Scienttfic Hydrologv.V.64: 127-130. Schmorak, S. & Mercado, A.1969. Upconing of freshwaterseawater interface below pumping wells: field study. Water Resources Research. V.5: 129-13 1 1. Buhmann, D. & Dreybrodt, W. 1987. Calcite dissolution kinetics in the system H20-C02-CaC03with participation of foreign ions. Chemical Geology. V.64: 89-102

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Water-Rock Interaction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Reactivity of pyrite surfaces: Combining XPS and speciation in solution M.Descostes, C.Beaucaire* & H.Pitsch* Laboratoire d'e'tudes des Inte'ractions Roche Eau, DCCIDESDISESD, CEA, 91191 Gif sur Yvette, France Laboratoire de Ge'ochimie des E a u , Universite'Paris 7 & Institut Physique du Globe de Paris, France * Present adress: IPSMIDPREISERGD, CEA-FAR,BP no6,92265 Fontenay aux Roses, France

F.Mercier UMR 8587 "Analyse et Environnement, CEAICNRSIUniversite'd 'Evry Val-d'Essonne, France

PZuddas Laboratoire de Ge'ochimie des Eaux, Universite' Paris 7 & Institut Physique du Globe de Paris, France

ABSTRACT: Pyrite oxidation process was followed combining surface and solution investigation techniques. Dissolution batch experiments in various pH media were carried out focusing on sulfur and iron speciation. In ions and Fe (11) predominant indicate a congruent acidic conditions, sulfur essentially present as SO-: dissolution. By contrast, increasing pH leads to an oxidation of iron to Fe(II1) with precipitation of ferric hydroxide, and a more complex sulfur speciation. Carbonate medium tends to stabilize ferrous iron with the formation of carbonate complexes. XPS analysis confirms iron precipitation and provides evidence for an oxidized sulfur component at 163.50 eV leading us to consider a first oxidative step at dissolving surface before production of first dissolved sulfoxyanion (thiosulfate). The pyrite layers concerned by the first solid state oxidation step increases with pH. studies illustrates the fact that oxidation mechanism is not yet well understood. In oxidative conditions, oxidation and dissolution reactions are not two welldistinguished processes. Oxidation can occur either at solid state and then knows dissolution, either after dissolution in solution, or according to both mechanisms. Finally, one may mention the incognizance of the aqueous solution chemistry and solid chemistry approaches (Luther 1997). We developed a multi-techniques approach during dissolution experiments over a pH range between 1 and 9. Iron and sulfur speciations in solution were followed while mineral surfaces were analyzed by X-ray photoelectron spectroscopy (XPS) in order to investigate iron and sulfur chemical and redox environments.

1 INTRODUCTION Pyrite (FeS2) is the more abundant sulfide at the Earth surface and easily undergoes through oxidative dissolution that causes acidity production and serious environmental damages (Singer & Stumm, 1970). At the pyrite surface, the behaviour of S and Fe depends upon both pH and redox conditions of the interacting solution (Goldhaber, 1983). In fact, ferrous iron can oxidize to ferric iron, while sulfur can be found in several oxidation states. The overall reaction of pyrite oxidation by oxygen is usually expressed by reaction 1 FeS2 + 7/2 0

2

+ H2O + Fe2++ 2 SO-:

+ 2 H+ (1)

Sulfate is the thermodynamically stable phase of S in oxidizing conditions. However, because of the high oxidation state of sulfur (+VI), sulfate can not be produced in a single oxidation step. In fact, elementary redox reactions are limited by a maximum transfer of two electrons (Basolo & Pierson 1958). Thus, sulfur can be found in several redox compounds with an oxidation state between (I) and (+VI). Electrochemical (Ahlberg et al. 1990, Bailey & Peters 1976) and surface (Sasaki et al. 1995, Zhu et al. 1994) studies have shown the formation of elemental sulfur on pyrite surface during oxidation in acidic and alkaline solutions. Sulfoxyanions such as tetrathionate (S40:-), thiosulfate ( S Z O ~ ~and ' ) sulfite (S032-) were also detected in basic solutions (Goldhaber 1983). Nevertheless, the large number of pyrite oxidation

2 EXPERIMENTAL AND METHODS Sulfur speciation and analysis was performed by capillary electrophoresis and ionic chromatography. Iron was analyzed by graphite furnace atomic absorption spectrometry while speciation was obtained by spectrophotometry (Viollier et al. 2000). Solid characterization was carried out by XPS (Descostes et al. 2000). Cubic samples of pyrite from Logroiio, Spain were first dipped with concentrated hydrochloric (37%) during several hours to eliminate any oxidation products present at the mineral surface. The pyrite was then introduced in a glove box (p(H20) and p(02) both inferior to 1 259

vpm) and rinsed with acetone. The mineral was ground in an agate mortar and sifted with ethanol (grain sizes in the 150 - 250 pm fraction). Pyrite was then washed in an ultra-sonic bath to remove any fine particles adhering to the grains surface. These two operations were repeated until the ethanol after ultrasonic-bath was clear. Samples were kept in glove box for drying until experiments. Batch experiment reactors were performed in run as batch experiments in a glass electrochemical cell with saturation in atmospheric oxygen (20%). Temperature was regulated through a heating-bath circulator at 25.0 k 0.1”C. Agitation by a Teflon@ stirring bar guaranteed a continuously homogeneous solution. The water to solid ratio was of 150 mL.g-’. Time course begun with pyrite introduction in solution. Dissolution experiments were carried out in different media depending on pH: hydrochloric acid (HCl 0.1 and 10-2 mol/L), perchloric acid (1 Om2, 10” and 10e6mol/L) and carbonated solution ([HC03-] = 104, 10” and 1,12.10-2mol/L) chosen to represent a clayey groundwater. Solution samples were filtered and immediately analyzed for sulfur and iron. The final samples were kept in globe box before XPS analysis. XPS determinations were performed on two samples, 6 and 25 hours interacted respectively.

Figure 1. Evolution of the ratio [SO~-]/[Fe],,,,, in acid media.

3 RESULTS 3.1 Dissolution

As observed in figure 1, in acidic conditions, dissolution of pyrite can be considered as congruent with an elementary ratio S/Fe close to 2. Neither iron nor sulfur is controlled by low solubility solids such as ferric hydroxide, ferrous sulfate and rhombohedric sulfur. At pH 5 2, S is essentially present as SO-: ions and Fe (11) is predominant (Fig. 2). As soon as pH exceeds 3 Fe (111) is predominant and is controlled by weak solubility ferric hydroxide. In these conditions, concentration remains below 1 pmol/L indicating precipitation (Fig. 3). The iron deficit seems to be linked with its oxidation state, with precipitation of ferric hydroxide. Iron oxidation into ferric species is active in a pH range 3-7. As shown in figure 2, for high HCO3- concentration (Plot F [HCO3-]=1.12.10-2 mol/L), Fe (11) becomes predominant, which can be related to the formation of carbonate complexes such as FeC03’, FeHC03’ (Criaud & Fouillac 1986). At the same time, s u l k speciation becomes more complex with thiosulfate apparition above pH=2 (Fig. 4). Total sulfur production is quite similar and does not seem to be dependant upon pH conditions.

Figure 2. Temporal evolutions of iron speciation in various media (A: HC104 IO-’ mol/L, B: HC104 10” mol/L, C: HC104 10-6 moVL, D: HC03- 104 moVL, E: HC03- 10” mol/L, F: HCO; 1.12.10-2moVL; 2q = 10 %).

Figure 3: Temporal evolutions of total iron concentration (square: HC104 10-* moVL, open triangle: HC1o4 10-6moVL, open circle: HC03- 10-4 moVL, open diamond: HC03- 10” mol/L, open square: HC03- 1.12. 10-2moVL).

260

Figure 4. Temporal evolutions of total s u l k and sulfoxyanion concentrations in various media (diamond: HC104 1O-* mol/L, triangle: HC03- 10” moVL, square and circle: HC03- 1.12.10’* moUL; open symbols stands for [SItorat- [SO,*-]).

Figure 6. S2p photoelectron peaks of pyrites oxidized in various media.

4 DISCUSSION 3.2 XPS analysis Fe2p3/2 photoelectron peak is constituted by two main components respectively at 707.15 (Fe&) and near 71 1.OO eV with a charge transfer satellite at the tail of the peak (Fig. 5). Its intensity raises with dissolution time and pH: very weak at pH=2, while pyrite component becomes in the minority. This second component is then dependant upon iron precipitation

XPS and solution chemistry data indicate that sulfur chemistry is controlled by pH. Speciation seems to become more complex when pH increases. In fact, kinetics of sulfur oxidation process is more rapid in acidic conditions and illustrate why thiosulfate is not detected. As observed by XPS analysis, component at 163.50 eV is related to an oxidized sulfur compound with oxidation state close to 0 (Fig. 7), leading us to consider an oxidation step at the surface according to reaction (2): FeS2 + ?40

2 = Fe-S-S-0

(2)

Figure 5. Fe2p photoelectron peaks of pyrites oxidized in various media.

S u l k S2p photoelectron (Fig. 6) is more complex with four components at 162.30 (Fe&), near 163.50 and 165.00 (polysulfides) and 169.00 eV (sulfates). The same chemical trend is observed with Fe2p peak: pH and time seem to govern sulfix chemistry. The three oxidized components intensity increases with pH medium and time dissolution. In acidic medium, the spectra are almost identical. Deconvolution is needed to go further in the interpretation.

Figure 7. Evolution of BE (in eV) of the S2p3,, photoelectron peak in function of the oxidation state of S. S*S03 and SS*S03 are used to distinguish the two different chemical environments for sulfur in S203*-,modified from Descostes et al. 2000.

The intensity of a photoelectron peak is directly function of the probed volume, i.e. for a same surface, function of the thickness analyzed. If this oxidized component was not clearly identified in

261

acidic solution, one may mention that the thickness concerned is too weak. Pyrite oxidation kinetic is then too rapid to allow observation of the reaction step. We only access to the overall reaction (1) expressed before. Furthermore, sulfur concentrations plotted as function of time (Fig. 4) are slightly different. Total sulfur knows a parabolic trend meanwhile sulfoxyanion concentration follows a linear tendency synonym of a surface reaction controlled kinetic. Therefore, we propose an oxidation mechanism with a first oxidation step at solid state. Thiosulfate is the first dissolved sulfoxyanion as proposed by several authors (Goldhaber 1983, Luther 1987, Moses et al. 1987). Its oxidation in solution into sulfite and then sulfate permits to write the overall oxidation reaction (1):

5 CONCLUSION On the basis of these first results, we have shown that Pyrite oxidation proceeds via several reaction steps controlled by sulfur chemistry with an oxidation at solid state before thiosulfate dissolution. Deconvolution of XPS data is needed to go further in the reaction mechanism. Sulfoxyanion dissolution is followed by several oxidation steps into sulfate. Kinetics of sulfur aqueous chemistry is governed by pH, with an oxidation faster in acid medium. Iron oxidation state is also controlled by acidobasic conditions but seems to be stabilized into ferrous form by carbonate complex. In acid conditions (pH below 3), iron remains ferrous and is oxidized into Fe(II1) as pH increases with oxihydroxide precipitation and mineral surface coating. Fe(II1) is known to be an oxidant kinetically more efficient than oxygen in strongly acid medium (Singer & Stumm 1970). XPS analysis didn’t allow to underline solid state oxidation in very acid conditions. Atomic Force Microscopy (AFM) study is in progress to quantify and verify the depth concerned by this solid state oxidation proposed as the first reaction step.

REFERENCES Ahlberg E., Forssberg, K.S.E. & X. Wang 1990. The surface oxidation of pyrite in alkaline solution. Journal of Applied Electroehemistry 20: 1033-1039. Bailey L.K. & E. Peters 1976. Decomposition of pyrite in acids by pressure leaching and anodization: The case for an electrochemical mechanism. Can. Met. Quart. 15: 333-344. Basolo F. & R.G. Pearson 1958. Mechanisms of inorganic reactions: A study of metals complexes in solution. Wiley, New York. Criaud A. & C. Fouillac 1986. Etude des eaux thermominerales carbogazeuses du Massif Central: Potentiel d’oxydorkduction et comportement du fer. Geochimica et Cosmochimica Acta 50: 525-533. Descostes M., Mercier, F., Thromat, N., Beaucaire, C. & M. Gautier-Soyer 2000. Use of XPS to the determination of chemical environment and oxidation state of iron and sulfur samples: Constitution of a data basis in binding energies for Fe and S reference compounds and applications to the evidence of surface species of an oxidized pyrite in a carbonate medium. Applied Surface Science 165: 288-302. Goldhaber, M.B. 1983. Experimental study of metastable sulfur oxyanion formation during pyrite oxidation at pH 6-9 and 30°C. American Journal ofscience 238: 193-217. Luther, 111 G.W. 1987. Pyrite oxidation and reduction: Molecular orbital theory considerations. Geochimica et Cosmochimica Acta 5 1: 3 193-3199. Luther, I11 G.W. 1997. Comment on “Confirmation of a sulfurrich layer on pyrite after oxidative dissolution by Fe(II1) ions around pH 2” by Sasaki K., M. Tsunekawa, S. Tanaka & H. Konno Geochimica et Cosmochimica Acta 6 1: 3269-327 1. Moses, C.O., Nordstrom, D.K., Herman, J.S. & A.L. Mills 1987. Aqueous pyrite oxidation by dissolved oxygen and ferric iron. Geochimica et Cosmochimica Acta 51: 15611571. Sasaki, K., Tsunekawa, M., Tanaka, S. & H. Konno 1995. Confirmation of a sulfur-rich layer on pyrite after oxidative dissolution by Fe(II1) ions around pH 2. Geochimica et Cosmochimica Acta 59: 3155-3158. Singer, P.C. & W. Stumm 1970. Acid mine drainage: The ratelimiting step. Science 167: 1121-1123. Viollier, E., Inglett, P.W., Hunter, K., Roychoudhury, A.N. & P. Van Cappellen 2000. The ferrozine method revisited: Fe(II)/Fe(III) determination in natural waters. Applied Geochemistry 15: 785-790. Zhu X., Li, J. & M.E. Wadsworth 1994. Characterization of surface layers formed during pyrite oxidation. Colloids and Surface A: Physicochemical and Engineering Aspects 93: 201-2 10.

ACKNOWLEDGMENTS The support of the “Agence Nationale pour la gestion des Dechets Radioactifs” through grant FT 00- 1-066 is gratefully acknowledged.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Probing the Electrical Double-Layer Structure at the Rutile-Water Interface with X-Ray Standing Waves P.Fenter. L.Cheng & S.Rihs Environmental Research Division, Argoizne National Laboratory, Argonne, Illinois 60439 USA

M .Machesky Illinois State Water Survey, 2204 GrifJith Drive, Champaign, IL 61820-7495

M .J .Bedzyk Deparment of Materials Science and Engineering, Northwestern University, Evanston IL 60208

N.C.Sturchio Earth and Environmental Sciences, University of Illinois, Chicago, Illinois 60607 USA

ABSTRACT: We demonstrate that the X-ray standing wave (XSW) technique is a powerful probe of the electrical double-layer (EDL) structure. Measurements were made of Sr adsorption at the rutile (1 10)-water interface from aqueous solutions. Our results show that Bragg XSW, using small-period standing waves generated by Bragg diffraction from the substrate, precisely probes the location of ions within the condensed layer, and the in situ partitioning of ions between the condensed and diffuse layers. Such measurements can provide important constraints for the development and verification of theoretical models that describe ion adsorption at the solid-water interface. 1 INTRODUCTION The development of mineral surface charge, and the associated distribution of solute ions at solid-water interfaces, generally referred to as the electrical double-layer (EDL), is fundamentally important for a diverse range of natural and industrial processes. Various physical and chemical models have been used to explain and predict the properties of the EDL (Stumm 1992, Dzombak 1990). A conventional schematic diagram of EDL structure, Figure 1, shows ions distributed between the so-called condensed (or "Stem") and diffuse (or "GouyChapman") layers, to balance a fixed charge of a mineral surface. However, there has been a lack of quantitative molecular-scale experimental data that can independently test available EDL models (Brown et al. 1999, Westall & Hohl 1980) and consequently our current understanding of the EDL is incomplete. Here we demonstrate that the Bragg X-ray standing wave (XSW) technique provides a direct element-specific probe of the in situ EDL structure (Fenter et al. 2000) by precisely measuring the location of condensed layer ions and the partitioning of ions between the condensed and diffuse layers. We demonstrate this capability by investigating Sr ion adsorption at the rutile (1 10)-water interface.

cally stable over a broad range of pH values (Machesky et al. 1994). Chemomechanically-polished synthetic single crystal rutile (1 10) substrates were cleaned ultrasonically in methanol to remove any surface organic contamination followed by three ultrasonic baths in nanopure (-18 MWcm) deionized water. Experimental solutions were prepared by dissolving reagent grade RbC1, RbOH, or Sr(N03)2 in nanopure deionized water, and the pH was adjusted using HNO3 andor NaOH. The aqueous speciation of Rb and Sr in the solutions was calculated by using the Geochemists Workbench; [Rblaq was exclusively present as Rb+ in all solutions, and > 98.5% of [%Iaq was present as Sr2+. The sample was held in a Kel-F "thin-film" cell using an 8 pm thick Kapton window, in which a thin solution layer is held against the sample surface through capillary action during XSW measurement.

2 EXPERIMENTAL METHODS The rutile (1 10) surface was used because it has been studied extensively and is known to be chemi-

263

The solutions were manually injected into the cell so as to expand the solution layer confined by the Kapton film during reaction to a macroscopic thickness (-1 mm). A pump was then used to apply a negative pressure to reduce and maintain the thickness of the solution layer to - 2 pm (Bedzyk et al. 1990). Between in situ measurements, the sample surface was cleaned by exposure to 1000 pM nitric acid solution. X-ray fluorescence measurements confirmed that the surface was free of any adsorbed Sr or Rb ion with a sensitivity limit of 5 x 104 ions/A2.

-

3 DESCRIPTION OF THE XSW TECHNIQUE The Bragg diffraction XSW technique has been described previously (Zegenhagen 1993). Briefly, the substrate Bragg reflection and X-ray fluorescence from EDL constituents are simultaneously measured as the sample is rotated through the Bragg reflection. Due to the nearly perfect reflectivity of the substrate Bragg reflection, an XSW field is generated by a coherent superposition of the incident and reflected Xray beams during Bragg diffraction (Batterman 1969). The XSW has a period that is equal to the dspacing of the diffraction planes, and the position of the XSW anti-nodes shifts inward by d/2 relative to the diffraction planes as the incident angle, 8, is scanned through the Bragg reflection (see Fig. 2). Consequently the fluorescence signal is modulated as the sample is rotated through the substrate Bragg reflection. The atomic position and distribution (characterized by the coherent position, P, and coherent fraction, J; respectively) are determined by monitoring the modulation of the fluorescence yield of a specific atomic species as the crystal is rotated through the Bragg condition. The fluorescence yield, normalized to the off-Bragg yield, varies as:

where the reflectivity R(8) and the XSW phase v(8) are derived from dynamical X-ray diffraction theory. Each set of XSW data (corresponding to a particular substrate reflection) are fully characterized by two model-independent parameters: the coherent position, PH, and the coherent fraction, fH. These parameters are obtained from aX2-fit of Eq. l to the XSW data. Uncertainties in fH and PH are typically smaller than *0.03, on the basis of counting statistics and a sensitivity analysis of the fitting procedure. In this study, H corresponds to the rutile (1 10) Bragg diffraction condition, and thus we denote fllo and Pl10 simply asfand P. Since the (110) reflection is normal to the surface, this provides a direct measure of the vertical EDL structure.

4 RESULTS Both ex situ and in situ XSW measurements of Sr adsorption to rutile (1 10) are shown in Figure 3. As expected based upon Eq. 1, the Sr fluorescent yield shows a clear enhancement near the rutile (1 10) Bragg peak position, whose shape is distinct from that of the substrate Bragg reflection. The modulation of the fluorescence yield depends sensitively upon the solution conditions. For instance, the angular variation of the ex situ Sr fluorescence data appear to be shifted to larger angles with respect to the TiOZ reflectivity, and the normalized fluorescent yield, YsJYoB, is observed to be less than 1 on the small angle side of the Bragg peak. (YsJYoB = 1 corresponds to the fluorescent yield in the absence of a reflected X-ray beam and is noted for each data set as a dashed horizontal line). The in situ data also show a distinct, but smaller, modulation. These are unambiguous indications of the standing wave effect which corresponds to the interference term proportional tofH cos[v(B)- 2n PHI in Eq. 1. We analyze these data using Eq. 1, from which we derive the coherent position, P, and coherent fraction,f; for each set of data. We find that P = 0.91 for the ex situ data and P = 0.84 for the in situ data. However, the coherent fraction, J for the ex situ samples is roughly twice the size of that found in situ. To understand the significance of these parameters, we note that there are three distinct contributions to the fluorescent signal in the standing wave

Figure 2. A schematic of the XSW experiment showing the XSW field superimposed on the EDL.

264

data: condensed layer, diffuse layer, and the bulk solution. For condensed layer ions in a unique site, the coherent fraction would have the value, f, = 1, and the coherent position, P,, will be determined by the height of these ions in units of the substrate lattice spacing (dllo = 3.25 A). The actual coherent fraction is typically smaller than this ideal value due to non-specifically adsorbed ions, atomic vibrations, multiple binding sites, etc, all of which tend to reduce the measured coherent fraction. The diffuse layer contribution to the coherent fraction is negligible, e.g., fd 0, as long as the Debye length is large compared with the Bragg plane spacing. For instance, the ionic strengths used in the present measurements are 5 4240 pM corresponding to Debye lengths of > 50 A. The diffuse layer will only contribute significantly to the coherent fraction when the Debye length is comparable to the Bragg plane spacing. Ions in the bulk solution have fbulk = 0 since their locations are, by definition, random with respect to the substrate lattice. Based upon these considerations only the condensed layer ions contribute to the coherent position. Therefore the measured coherent fraction, J; represents the fraction of ions located in the condensed layer. In other words, the measured coherent position and fractions are given by:

-

where P, is the coherent position of the condensed layer ions, and 8 is the coverage of each layer projected onto the surface plane. The most precise results are obtained when the coherent fraction is large. Based upon this simple expression, it is easy to see that this is achieved by minimizing &lk. Our measurements are therefore performed under conditions where we minimize the bulk solution layer thickness (estimated thickness 2 pm). Under these conditions can be calculated as ebulk(ML) = 2.3~10"' [SrIaq(pM),where the coverage is expressed in units of monolayers monolayer (ML) is defined here as 1 ML = 5 . 2 ~ 1 0 ions/cm2] and the Sr ion concentration is expressed in pM. Since typical ion coverages in these measurements are 0.25 to 0.5 ML, we can expect to be able to use this technique to probe solution ion concentrations as high as 0.001 M. We can, however, probe the EDL structure at significantly higher ionic strengths, particularly if the ionic strength is determined by an ion whose X-ray fluorescence line does not interfere with the fluorescence line due to the ion of interest. When the solution concentration is small with respect to the sum of the condensed and diffuse layer coverages (i.e., 8, + e d >> @,ulk) the measured coherent fraction becomes independent of the exact solution thickness. In this regime, the measured coherent fraction is equal to the product of the doublelayer partition coefficient, X = 8, /(& + &) and the coherent fraction of the condensed layer ios, f,. In this way, systematic measurements of the coherent fraction as a function of solution parameters can provide direct insight into the partitioning of ions between the condensed and diffuse ion layers. The modulation of the in situ Sr fluorescence as a function of angle is substantially weaker than that in the ex situ data (Fig. 3). This indicates that a substantial fraction of Sr ions for the in situ measurements are found in the diffuse layer. We also note that we see no significant adsorption of Sr after exposure at pH = 3.2. In situ measurements also reveal a negligible coherent fraction at pH 3.2 indicating a random Sr ion location (i.e., Sr ions are not found in the condensed layer). These results suggest that the non-zero f values derived from the in situ measurements at pH 11 reflect the in situ EDL structure of Sr ions. Our data show that the coherent fraction increases for both in situ and ex situ data when the samples are exposed to the highest solution concentrations, [%Iaq. This implies that the fraction of double layer ions found in the condensed layer increases with [Sr],. The coherent position measured in situ is -7% smaller than that measured ex situ corresponding to a difference of the ion heights under the respective conditions of 0.23 A. The overall similarity between the in situ and ex-situ results, coupled with the high

-

k

-

-

-

Figure 3. XSW data for Sr adsorption to rutile (1 10) at pH 10. The substrate 110 Bragg peak is shown at the bottom (open circles), and Sr fluorescence for in situ (open symbols) and ex situ conditions (closed symbols) are shown. Each set of fluorescence data is offset vertically by 0.5.

265

reproducibility of these measurements suggests that the Sr ion location is determined primarily by the ion-substrate interaction, and further suggests a welldefined adsorption site. It also provides confidence in the derived Sr ion location for the in situ measurements in spite of the relatively small coherent fractions found in those measurements. However the small but finite difference between these results appears to be significant especially as it is found systematically over a broad range of solution ion concentrations. This suggests that the adsorbed Sr location exhibits some sensitivity to its environment but needs to be confirmed with further measurements. We have demonstrated the capability to directly probe important aspects of the EDL at mineral-fluid interfaces using the Bragg XSW technique for both in situ and ex situ measurements. Aspects of the EDL structure that can be directly probed in this manner include the location of ions in the condensed layer and the partitioning of ions between the condensed and diffuse layers as a function of pH and solution ion concentration. These results suggest that these aspects of the ion distribution near a mineral-water interface can now be measured directly, in situ, to yield a truly atomistic understanding of the EDL structure.

Stumm, W. 1992. Chemistry of the Solid-Water Interface. WiIey-Interscience. Westall, J. & H. Hohl 1980. A comparison Of electrostatic interface. Cozloid Inter-

~~~~j~2r",e,~~~~'ution

Zegenhagen, J. 1993. ray standing waves.

ACKNOWLEDGMENTS This work was supported by U. S. Department of Energy, Office of Basic Energy Sciences, Office of Chemical Sciences, Geosciences and Biosciences. These experiments were performed at beamline 12-ID-D at the Advanced Photon Source which is supported by the Office of Basic Energy Sciences, U.S. Department of Energy, under contract W-3 1- 109-ENG-38 at Argonne National Laboratory.

REFERENCES Batterman, B.W. 1969. Detection of foreign atom sites by their x-ray fluorescence scattering. Phys. Rev. Lett. 22: 703-705. Bedzyk, M. J., Bommarito, G. M., Caffrey, M. & T.L. Penner 1990. Diffuse-double layer at a membrane-aqueous interface measured with x-ray standing waves. Science 248: 5256. Brown Jr., G.E., Heinrich, V.E., Casey, W.H., Clark, D.L., Eggleston, C., Felmy, A., Goodman, D.W., Gratzel, M., Maciel, G., McCarthy, M.I., Nealson, K.H., Sverjensky, D.A., Toney, M.F. & J.M Zachara 1999. Metal oxide surfaces and their interactions with aqueous solutions and microbial organisms. Chem. Rev. 99: 77-174. Dzombak, D.A. & F.M.M. More1 1990. Surface Complexation Modeling: Hydrous Ferric Oxide. Wiley-Interscience. Fenter, P., Cheng, L., Rihs, S., Machesky, M., Bedzyk, M.J. & N. C. Sturchio 1990. Electrical double-layer structure at the rutile-water interface as observed in situ with small-period x-ray standing waves. J. Coll. Int. Sci. 225: 154-165. Machesky, M.L., Palmer, D.L. & D.J. Wesolowski 1994. Hydrogen ion adsorption at the rutile-water interface to 250°C. Geochim. Cosmochim. Acta 58: 5627-5632.

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su$

surface determination with Scj. Reports 18: 199-271.

x-

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Enriched stable isotopes for determining the sorbed element fraction in soils in order to calculate sorption isotherms H .-E.Gabler & A . Bahr Federal Institute for Geosciences and Natural Resources (BGR),Hannover, Germany

ABSTRACT: The fraction of sorbed elements in soils is determined by the use of enriched stable isotopes. For this purpose an isotope dilution mass spectrometric (IDMS) technique has been applied for the elements Cd, Cr, Cu, MO, Ni, Pb, T1, and Zn. The results of the IDMS technique are comparable to the results of conventionally used EDTA extractions for elements forming stable EDTA complexes. For elements for which the EDTA technique fails, IDMS is a valuable alternative, particularly with regard to its easier handling. Adsorption isotherms were determined for T1, demonstrating that IDMS is advantageous for determining the fraction of sorbed elements that form weak EDTA complexes.

1 INTRODUCTION The increasing availability of ICP-MS instruments opens the possibility of using techniques involving nonradioactive isotopic tracers to a growing number of users in the environmental sciences. An isotope dilution (ID) technique using enriched stable isotopes can be employed to determine the “interchangeable fraction of heavy metals in soils” by batch experiments (Gabler et al. 1999). The interchangeable fraction is defined as that portion of an element in a sample that rapidly exchanges with an added spike during a batch experiment. The added spike contains an exactly known quantity of each element of interest enriched in one of the element’s stable isotopes. Thus, this spike is the same as normally used for isotope dilution mass spectrometry (IDMS) (Heumann 1988). In principle, this technique can be applied for all elements that have at least two stable or long-lived isotopes. This is true for many elements (Heumann 1988). To determine sorption isotherms by sorption experiments, it is necessary to first determine the original proportion (SO) of the element that participates in sorption-desorption in the soil (Filius et al. 1998, Springob & Bottcher 1998). In sorption experiments, the sorbed concentrations are then calculated from the difference between the amount of that element added at the beginning of the experiment plus SOand the amount recovered at the end of the experiment. If small amounts are added, SO becomes increasingly important for correct determination of the sorbed amount (Springob & Bottcher 1998), because here SOoften represents a

significant proportion of the sorbed amount. Conventionally, the analysis of So is done by extraction with various solutions, e.g., 0.43 M HN03 (Boekhold et al. 1993) or EDTA (ethylenediamine tetraacetic acid) (Filius et al. 1998), assuming that (i) SO is completely extracted from the soil and transferred to the liquid phase and (ii) none of the element is extracted that does not participate in the sorption equilibrium. Gabler et al. (1999) and Young et al. (2000) recently demonstrated that ID techniques used for the determination of so-called “interchangeable” or “labile” element fractions in soils are robust over a wide range of experimental conditions. The interchangeable or labile element fractions can be regarded as SO.The advantage of the ID techniques is that the labile or interchangeable element portions need not to be transferred completely into solution by a strong extractant, because ID uses isotope ratios (enriched stable isotopes) or equilibrium solution concentrations and isotopic distribution coefficients (radioisotopes). Gabler et al. (1999) use an ID technique using enriched stable isotopes, which requires a mass spectrometer, while Young et al. (2000) use radioisotopes, which require radiometric equipment. The increasing availability of inductively coupled plasma mass spectrometry (ICP-MS), the simplicity of the enriched stable isotope technique, and the fact that this technique can be applied to many elements make this technique advantageous. This study presents the results of recovery experiments with the IDMS technique and compares this technique with the EDTA technique for the determination of So in 45 soil samples for the

267

elements Cd, Cr, Cu, MO, Ni, Pb, T1, and Zn. Adsorption experiments for T1 were carried out and sorption isotherms are calculated on the basis of both IDMS and EDTA for SOdetermination. 2 MATERIALS AND METHODS 2.1 Soil samples Samples of 46 soils (32 sandy soils, one silty soil, eight loamy soils and five clayey soils) were taken from areas in Lower Saxony and MecklenburgVorpommern, Germany, in which the soil pH lies within the silicate and carbonate buffer ranges. The samples were dried at 40 "C and sieved to obtain the l , unfavorable process: a=l , linear process; O [Gd] > [Eu(lI)] (2001). Figure 4 illustrates the main peculiarities of relationships between concentrations of Eu in its speciation at 350°C depending on pH in the fluid, various valence states and Gd concentration have which is saturated relatively to EuF3. In the acid area been revealed as a critical result of thermodynamic the figure shows that EuF2' and EuF2' dominate and modeling of Eu fractionation between mixture of EuC12" and EuC12' contribute to some extent to the solid REE fluorides and hydrothermal fluid. In this total Eu content in the presence of small HCl case, Gd is the nearest to Eu representative of stable concentration (it is added into the model fluid as an 3-valent REE series. Discussing the probable trends acid agent). In alkaline area Eu02- and H E U O ~ ~of Eu fractionation in natural REE - fluorite hydroxide complexes (equal to Eu(OH)4- and Eu depending on redox conditions in fluid, Eu maxima (OH)3O respectively) are the main aqueous species should be found in CaF2 which forms in reduced but both fluoride and hydroxide species can exist in Eu2'-fluid when Eu minima should be typical of intermediate conditions at pH = 4-6. A complete CaF2 precipitated from oxidized fluid or treated by predominance of chloride complexes with this kind of fluid. In other words, according to the concentrations of 10-9-10'4 mol/kg H20 is exist for popular geochemical criterions, if [Eu]/[Gd]solid, i.e. Eu(II) in dilute HCI solutions after EuF3 dissolution the ratio of contents of these REE in any mineral, when concentrations of EuF', EuF2' and other Eu(I1) means Eu/Euv, we can expect that Eu/Eu* > 1 in fluoride species reach only 10-" mol/kg H20. reduced fluid but Eu/Eu* < 1 in oxidized fluid. The temperature decrease up to 100°C shows the 2. The current results correspond to the restricted change of V-shape of solubility curve for EuF3 to form of Eu partition model, taking into account only trough-shape with 104 mol/kg H2O concentration of hydroxide and fluorine types of REE complex E u F ~ complex " (prevalent at pH between 4 and 7.5) formation in fluid. The implication both of other and both additional small branch of positive-charged types of the expected complexes (chloride, sulfate, fluoride complexes at pH < 4 and a small branch of carbonate and bicarbonate) and geochemical data on real composition of mineral-forming fluids is hydroxide complexes (Eu02-, HEuO: ) at pH > 8.

289

REFERENCES

speciationof main existingin fluid in EuF3, depending on pH. ~i~~~~

4,

Barin. I. 1989. Therniocheniical data ofpure substances. VCH. Weinheim: New York. Bau. M. 1991. Rare-earth element mobility during hydrothennal and metamorphic rock interaction and the significance of the oxidation state of europium. Cheni. Geol. 93: 219-230. Greis, 0. & J.M. Hascllke 1982. Rare earth fluorides. In K.A.Gsclmeidner, Jr. & LeRoy Eyring (ed.), Handbook on the ph-vsics and cheniistry ofrare earth: 387-460. Grappin, C., M.Treui1, S.Yanan & J.S.Touray 1979. Le spectre des terres rares de la fluorine en tant que marqueur des proprietes du milieu de depot. Mineral. Deposrta. 14: 297309. Haas, J.R., Shock E.L. & D.S. Sassani 1995. Rare earth elements in hydrothermal systems: estimates of standard partial niolal thennodynamic properties of aqueous complexes of the m e earth elements at high pressures and temperatures. Goch. et (’osmochini. Acta. 59: 4329-4350. Kolonin, G.R. & G.P. Sllironosova 200 1. Thennodynamnic model of dissolution of Ca, REE and Y fluorides in fluid of coinplex composition at temperatures from 100 to 500°C. Proceedings of Joint ISHR and KYTR (h’ochi,Japan) (in press). K~ipriyaiiova, 1.1. M.D.Ryazantseva, 0.A.Kukuslkina & K.A.Kuvshinova 2000. Typoinorpllislri of fluorite from ore different forination types in Sikhote-Alin region (Primorye) Proceeding,y of the Russian ,$,fineralogical ,yociev, 129(3): 39-54.

~ ~ ( 1 1 1aqueous ) with

developlnent is to use’ physi~al-chelnical modeling for the “demarcation” Of ’pace Of occurrence of mineral-forming fluid, where Eu(II1) and Eu(I1) should be stable. We believe that the first results to develop previous data (Sverjensky 1984; Wood 1990; Bau 1991) should be presented by us in the near future. 4. It is obvious that the equilibrium model under discussion is only the first approximation to a very complicated problem of adequate insight into mineral-fluid interaction, concerning REE pattern in fluorite. Taking into account the extremely Complicated way to this equilibrium is reached we need to involve an additional parameter, named conventionally as “water/rock ratio”, to describe a reaction progress (Bau 1991).

Inorgarlic species in geologic fluids: Correlation among standard niolal thermodynamics properties of aqueous ion and hydroxide complexes. Geoch. et (bsniochinz. Acta. 6 1: ‘)()7-950, Sverjensky, D.A. 1984. Europium redos equilibria in aqueous solution. Earth Planet. Sci. Lett. 67: 70-78. Slivarov, Yu.V. 1999. Algorithinization of the numeric equilibrium inodeling geochemical processes. Geokhinzya. 6: 6-16-6.52. Wood, S.A. 1990. The aqueous geoclieinistry of rare earth elenients and yttrium. Part 2. Theoretical predictions of speciation in hydrotliennal solutions to 350°C at satmtion water vapor pressure. (’hem. Geol. 88: 99-12.5.

ACKNOWLEDGEMENTS : This study was jnancial(y supported by RFBR grants 98-05-65299, 01-05-65255 and by the program “Universities of Russia-fundamental researches”, grant 2787.

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Elements transfers in compacted clayey materials under thermal gradient C .Latrille, M .Jullien & C .Pozo Center of Nuclear Energy, Cadarache, France

ABSTRACT: During the Stripa mine experiments (Sweden), the heat-induced transfers of elements in a compacted clay were investigated and related with mineralogical transformations. After four years of experiments, results indicate that Fe, AI and Si moved towards the colder clayey material while Mg and Ca accumulated at the hottest parts of the system. Weathering, dissolution and crystallisation processes were observed at the clay layer scale. Dissolution of kaolinite began from 40°C. The weathering processes led to the formation of largely distributed gels which are able to crystallise depending on the chemical conditions. With increasing temperature, the composition of smectites changed and recorded the chemical environment, comparatively to the parent material. New crystallisation of berthierine-like minerals, saponite and opal were observed in the hottest area. A Si-Al-Fe gel could enable elemental transfers as a support for the migration. The reactivity associating gels-elements-temperature induced sequences of crystallisation which depend on the coupling of Si-Mg and Fe-Al. 1 INTRODUCTION

The clayey material was ground into a fine powder containing fine matrix and millimetric nodules. This material was isostatically compacted at 25 MPa with a dry density of 1.82.

In the Stripa mine (Sweden), a one-scale cylindrical simulated engineered barrier system (EBS) composed of compacted clay was submitted during 4 years to saturation by granitic groundwater and to a thermal gradient. The heat source was a low cast iron steel heater embedded in the compacted clayey materials. This heat source also produces an iron input into the system. The aim of the study is to highlight heat-induced transfers of elements in relation with the mineralogical evolution of the initial clay and the iron input. In a first step, mineral phase markers of the elements transfers in the system are looked for. In a second step, hypothesis on the chemical conditions of the material transformation are exposed.

2.2 Experimental system The experimental system (Fig. 1) was constituted by a cylindrical low cast iron steel heater (51 mm in

2 MATERIAL AND METHODS 2.1 Materials The compacted clayey material used for the experiment was the FoCa7 (Gin et al., 2000) which is characterised in the present study by Analytical Scanning Electron Microscopy and Analytical Transmission Electron Microscopy.

Figure 1. Scheme of the experimental system.

291

diameter) embedded in the compacted clayey materials. This system was disposed in a 350 m deep granitic cylindrical shaft. The clayey experimental system was rehydrated by the granitic groundwater which contains Ca, Mg, S, Si and Na (Grimaud et al. 1990). The rehydration of the clay barrier was carried out during the first few days of the experiment (Pusch et al. 1992) and by a centripetal way. The low cast iron steel heater was thermally raised to 200°C and imposed a thermal gradient in the clayey material between 170°C and 40°C close to the granitic rock over 4 years. The system was then switched off and dismounted after cooling byovercoring .

of 60 nrn, a count time of 60s, a dead time of 10 % and a count number of 1000. Semi-quantitative data were obtained by a Link ISIS software (Oxford Instruments). Elemental compositions of clay aggregates and particles were calculated in phyllosilicate structural formulae expressed with 11 oxygen.

3 RESULTS 3.1 Structural and chemical changes

2.3 Analytical methods The mobility of elements and mineralogical changes were evidenced by multi-scale studies including photon microscopy, analytical electron microscopy (SEM/EDS and TEMEDS) and bulk chemical analysis (ICP/MS). Photon microscopy and bulk chemical analyses permit to define areas of differentiation in the material by their colour and relative elemental concentrations. These methods express the macroscopic differentiation in the clayey material under thermal gradient in a groundwater saturated system. SEM/EDS analyses Polished sections were prepared from the compacted clayey material. Samples were carbon coated and analysed by a Jeol 840 SEM at 15 kV and 0.3 nA, equipped with a Si(Li) diode detector (Oxford Instruments) with a SATW window. Around 100 punctual analyses were carried out on homogeneous clayey areas. Analyses were performed with a depth penetration of 2 pm, a count time of 50s and a dead time of 20%. Semi-quantitative data were obtained by a Link ISIS software (Oxford Instruments). TEM/EDS analyses Samples were taken with a fine needle in the areas differentiated by photon microscopy. These samples were reduced in a homogeneous powder and dispersed in pure water. Suspensions were deposited on copper TEM grids and air dried. Samples were examined by a Jeol 2000 FX TEM at 200 kV and 109 pA, equipped with a Si(Li) diode detector (Oxford Instrument) with a SATW window. Around 100 punctual analyses were carried out on individual particles of each sample. Analyses were operated in convergent mode with a probe size 292

Optical microscopy observations (Fig. 2) indicate a differentiation in the clayey material after thermal treatment. Four areas are distinguished by colour and structural change. The weakly transformed area preserved the parent clayey material colour and structure. In the transitional area, the differentiation is mainly expressed by a lightening of the matrix colour. The well transformed area is characterised by the deconstruction and the tanning of the material. The hardness area (few millimetres) is compact and black. Bulk chemical analyses of each area (Fig. 3) show a large increase of Fe, Ca and Mg close to the low cast iron steel heater. These results show an input of these elements by the corrosion of the low cast iron steel heater (Fe) and their migration through EBS and from the groundwater (Ca, Mg). Conversely, Si and A1 content decrease near the cold area. This induce their mobility along the thermal gradient.

Figure 2. Area differentiation by optical microscopy in the rehydrated clayey material after thermal treatment.

3.2 Mineralogical change and element transfers From the colder area to the hot area, the mineralogical composition changes. The parent clayey material is mainly composed of smectitic

Figure 3. Bulk chemical analyses of Si, AI, Fe, Mg and Ca in the different area.

phases, kaolinite and goethite. Progressively, kaolinite is dissolved (Fig. 4a) and disappears in the well transformed area (Fig. 4b). Concomitantly, a gel containing silica spheroids appears (Fig. 4b). The hardness area is characterised by the formation of berthierine (Fig. 4c), opal and saponite (Fig. 4d) while other smectitic phases and part of the goethite disappear. These mineralogical changes have recorded the chemical environment. Effectively, the elemental transfers are expressed through the structural formulae of the phyllosilicate phases of each area, analysed by SEM/EDS (Table 1). These results show the relative increase of Mg Fe in the area and the hardness area. They confirm the bulk chemical data. The relative content in s i seems to be Constant throughout the thermal gradient and indicates the presence of an individual Si phase, such as quartz. The relative content of A1 in these structural formulae increases close to the colder area and illustrates the disappearance of kaolinite.

(Si3.54 A10.46) (Al0,83 Fe0.31 Mg1,31) 0 1 0 (OH12 Ca0,ZZ (Si3,51AI0.49) (A10,76Feo,73 Mg,,,,) 0 1 0 (OH), Ca0.26 (si3.41 A10.59) (A1135 Fe0.46 MgO.11) Of0(OH):! ca0.19 (Si3.42A10,58)(A11.58Feo,38Mgo.07) 0 1 0 (OH), cao.15

Figure 4. Kaolinite dissolution in the weakly transformed area (a) ; crumbled gel containing Si spheroi’ds and “ghosts” of kaolinite in the well transformed area (b) ; crystallisation of saponite and opale in the hardness area (c); formation of berthierine in the hardness area (d).

thermal gradient is expressed in the beidellite composition. The compositions of gel and kaolinite indicate the remove of Si into a gel during the

in 0,,(OW2

Hardness area

SI

Transitional area Weakly changed area Parent clayey material FoCa 7

The distribution of Si and Fe contained in the mineral component of each area is expressed in the structural formulae of 2: 1 phyllosilicate (Fig.s 5 and 6). components are 2:1 phyllosilicates (rnontmorillonite, beidellite and saponite), kaolinite, berthierine and a gel. Decrease of Si throughout the

Figure 5. Distribution of Si in the structural formulae of mineral phases identified by TEM/EDS.

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REFERENCES Gin, S., Jollivet, P., Mestre, J.P., Jullien, M. & C. Pozo 2000. French SON 68 nuclear glass alteration mecanisms on contact with clay media. Accepted for publication in Applied Geochemistry. Grimaud, D., Beaucaire, C. & G. Michard 1990. Modelling of the evolution of ground waters in a granite system at low temperature: the Stripa groundwaters, Sweden. Applied Geochemistry 5: 5 15-525. Pusch, R., Karnland, O., Lajudie, A., Lechelle, J. & A. Bouchet 1992. Hydrothemal field test with french candidate clay embedding steel heater in th Stripa mine. SKB Technical Report 9302, Stockholm: 85p. Figure 6. Distribution of Fe in the structural formulae of mineral phases identified by TEWEDS.

disappearance of kaolinite and smectite transformation. Important transformation and Si mobility is also shown by the new crystallisation of saponite and berthierine in the hardness area, which are respectively rich and poor in Si. Mineralogical transformations are expressed in the Fe distribution. The Fe octahedral occupancy in the beidellite and gel increases close to the heater. Berthierine which contains high Fe proportion appears. Saponite expresses a high Mg mineral phase. Gel has a composition close to a 2:l phyllosilicate and may be the product of dissolution and precursor of new crystallisation. At local scale, a coupling between Si and Mg, in a part, and between A1 and Fe in other part was evidenced through reactional sequences and relative new crystallisation of saponite and opale; or berthierine.

4 CONCLUSION The migration of the elements is expressed at each scale of the study. The environmental disturbance is recorded by the mineralogical composition and is expressed mainly by the saponite and berthierine new crystallisation and the kaolinite and smectite dissolution andor weathering. Element transfers are made through an hydrated gel which is probably the mineral precursor and the active support of the elemental and thermal transfers. Clay minerals, compacted clay material (EBS) are extremely reactive under thermal gradient.

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The Chemical Durability of Yttria-Stabilized ZrO, pH and 0, Geothermal Sensors M.F.Manna, D.E.Grandstaff & G.C.Ulmer Temple University, Department of Geology, Philadelphia, Pa I91 22 USA

E .P.Vicenzi Smithsonian Institution, National Museum of Natural History, Washington D.C. 20560 USA

ABSTRACT: In geothermal situations, like the acidic gases emanating from a volcano, the redox potential of the volcanic vapors has been measured with yttrium-stabilized zirconia (YSZ) sensors (Vulcano: Nuccio et al. 1993, and Mount Etna: Sat0 et a1 1973). If a standard glass pH electrode were used in such a hightemperature (>9OoC) geothermal environment, the electrode would quickly lose calibration, making it unreliable. However, YSZ sensors are now being developed that could measure the pH of such fluids. The YSZ sensor is already known to be more stable for use in geothermal environments. However, a major concern is the long-term stability of the YSZ solid solution in acidic environments. The solubility of yttria and zirconia and their solid solutions are less well known then those of the accessory components, such as the oxides and compounds of Ti, Fe, Ca, A1 and Si that are presenting commercial ZrO2 ceramics. We have conducted experiments to determine the amount of the materials that can be leached from three available YSZ ceramics. Leaching likely effect the impedance-T behavior of the YSZ sensors in long-term exposure. YSZ materials were leached in two concentrations of HC1 for 12 weeks with samples being taken every 2 weeks. The leachate solutions were analyzed with ICP-MS for all the elements listed above to determine the mass of leached materials. 1 INTRODUCTION Yttria-stabilized zirconia (YSZ) membranes have been used for many years to measure redox in industrial and geological environments. The most commonly used Y S Z membrane (for example see more than two decades of research papers in the Journal of the Electroanalytical Chemistry Society) is fiom a single manufacturer whose ceramic has been military- and ISO- (International Organization for Standardization) certified. Our collaborative research (Lvov et al. in press) has shown that an accurate (A 0.05 log units) pH sensor can be devised from this YzO3-stabilized (%mole%) cubic ZrO2 membrane. These pH sensors have been used in laboratory testing for hundreds of hours, and in our work at temperatures up to 38OOC. In our tests, these display Nernstian behavior over a studied pH range of 10 -3 molal HCl to lOW4molal NaOH when tested in a flow-through configuration especially designed to minimize thermal gradients and Soret-effects. However, during this research, it has been found that this commercially available 2 3 - 0 2 ceramic (V-I) is not sufficiently durable, particularly in acidic, hightemperature solutions, to allow longer term use in

field applications such as in geothermal wells, nuclear waste disposal sites or general environmental monitoring.

2 THE DURABILITY PROBLEMS In part, the poor long-term durability in V-I is due to dissolution and alteration of interstitial glass. Also of concern is leaching of yttrium from the YSZ solid solution. Figure 1 shows a glass phase (dark gray

Figure 1 is a “mixed” (50% secondary - 50% backscattered) electron image of a fractured edge of the V-I YSZ (hlly yttria stabilized) ceramic. The light grains are the zirconia and the dark gray material is the interstitial Ca-Al-Na-silicate glass.

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forming in triple grain junctions or along grain boundaries near the gent of glass dissolution. The interstitial glass is ionic and when present, increases the bulk conductivity of the sensor; effectively lowering the bulk impedance of the ceramic. As the glass is dissolved or altered to form zeolites, the bulk impedance will increase. Zeolitization of the interstitial glass will introduce a volume change in materials, weakening the triple grain junctions, thus introducing the possibility of mechanical degradation of the ceramic.

material), which forms a virtually continuous 3dimensional network between the ZrOZ grains (light material). Electron microprobe analyses of the interstitial glass (Ulmer et al. 2000) reveal that it is a Ca-Al-Na-silicate that is feldspathic-like in composition but non-stoichiometric. Figure 2 shows etching of this glass in acidic and hydrothermal environments. Note in image A, the glass on the outside edge (bottom of image) has been leached to about 12% of the total thickness by the dilute HC1 (as shown by the roughened zone) and in image B, both sides of the ceramic were exposed to the HC1 and both show a 25% depth leaching with the much stronger acid. Progressive dissolution of the interstitial glass will ultimately cause mechanical and electronic failure of the ceramic membrane.

I Ceramic Material Characteristics:

I

V-I contains the major interstitial feldspathic glass V-11 contains minor interstitial, multi-composition silicates V-111 is made from ultra-pure S-moie% powder

3 THE LEACHABILITY PROBLEMS Another concern is the chemical stability of the ZrOz-Y203 solid solution. Thermochemical data (Roine 1999) suggest that the Y203-Z-02 solid solution is more soluble than pure ZrOz, presumably because incorporation of yttria introduces defects into the ZrO2 structure. We have performed experiments to determine if exposure to high temperature and low pH environments will lead to leaching of Y203 from the solid solution, which would destabilize the cubic phase. Theoretically, the leaching of the Y203 fiom the solid solution would allow the cubic ZrOz-YzO3 to revert back to the tetragonal phase which would lead to a change in electronic behavior as well as reducing the ceramic’s thermal shock resistance, i.e. ability to tolerate polymorphic volume changes resulting from thermal changes. (Stubican et al. 1984) We have studied two commercially available ceramic shapes (designated as V-I and V-11) and a custom-made, isostatically pressed ceramic shape made from 99.99% pure powder (V-111). V-I, as shown in Figure 1-3 contains the interstitial glass described previously. Ceramic V-I1 contains various discrete silicate and oxide grain boundary phases, such as periclase, diopside, and wollastonite (Ulmer et al. 2000). Ceramic V-111 has been custom-made from ultra-pure 8-nioleY0 YSZ powder. We investigated both bulk ‘monolithic chunks’ (sawed-out, tube-shaped pieces) and powders where possible. For powders, the bulk ceramics were crushed to a uniform grain size of 125 to 250 pm; BET analysis showed surface areas for all powders to be from 0.082 to 0.132 m2/g. These materials were then placed into PFA Teflon@bottles, in solutions made with ultra-pure HC1, initially of pH 1 and pH 3. These solutions were maintained at

Figure 2 shows two secondary electron images of V- 1 ceramic. Image A, shows a 75pm leached layer resulting from several hundred hours exposure to dilute HCL, not more acidic then 10-3 molal. Image B, shows a 175pm leached layer from exposure to boiling ION HCL for 10 minutes. (The items in the dark material are bubbles in the epoxy mount and are not part of the ceramic.) Both scale bars are 200pm.

Furthermore, in addition to outright dissolution, the glass in this much-studied ceramic material, also zeolitizes (Fig 3) as it degrades due to exposure to hydrothermal fluids. Figure 3 shows radiating platy or bladed crystals, of an as-yet-unidentified zeolite,

Figure 3 is a mixed electron image (as in figure 1) of the V-I YSZ ceramic with zeolitization of the interstitial glass in the grain triple junctions. Note that the previous interstitial glass i s missing in the triple junctions and as well as in the grain boundaries. Scale bar is 1 pm.

296

90°C and were agitated each day to promote mixing. Solutions were sampled every two weeks. The bottles were refilled at the 8-week point to maintain a reasonable water to solid ratio. Solutions were analyzed by Inductively Coupled Plasma - Mass Spectrometry (ICP-MS) using the method developed by Field et aL (1999). After 12 weeks the experiments were ended and solid materials examined. We determined the leaching characteristics of the YSZ solid solution and interstitial phases (where present).

4 LEACHABILITY RESULTS Concentrations of most dissolved species were fairly constant after about two weeks. These stabilized values are shown in Figures 4 and 5. These figures show concentrations of various dissolved species (indicated by placement of chemical symbols) leached fiom powders and monolithic chunks (indicated by "Solid") fiom three different ceramics. Our ICP-MS data reveal two types of chemical leachability: the 'chemical impurities' (Type 1) Al, Ca, Fe, Mg, Si, Ti, and the ceramic matrix (Type 2) Y and Zr. Figure 4 shows results after two weeks of leaching at 90°C at an initial pH of 1 (conditions that were the most strident examined in our test matrix). Figure 5 shows the results of leaching at 9OoC and an initial pH of 3. As expected, the Type 1 chemical

itself, leaching of glass or impurities could perturb pH measurements in a non- or slowly- flowing system. Zr concentrations are low (0.05-0.3 ppm) for these materials. These concentrations generally agree with those predicted for solubility of Zr02 at 90°C fiom the thermochemical data of Shock et al. (1997). This might indicate that the 8-moleY0YSZ is not significantly more soluble than pure ZrO2, in contrast to data of Roine (1999) or that the dissolved Zr concentration is controlled by precipitation of secondary ZrO2. However, the Zr/Y ratios are highly variable between vendors. The Zr/Y ratios for both powders and monoliths of V-I and V-I11 are nominally stoichometric (-6) given the 8-mole% Y2O3 in the ceramic. However, the Zr/Y ratio in the V-I1 solids is -1/2 and the Zr/Y ratio in V-I1 powdered material is very much lower, approximately 1400. To understand this we have also done quantitative Electron Microprobe (EMP) line scans for both the V-I and the V-I1 solids. These scans show that the Y2O3 concentrations in both the V-I and V-I1 solids vary by only k 0.2 mole% within the zirconia grains. pH 3, 90°C, 2 weeks

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However, the Y concentration is systematically greater in grain boundary phases in V-I1 and systematically lower in the grain boundary phases in V-I. Therefore, the higher leachability of Y and high degree of Zr/Y non-stoichiometry in V-I1 materials appears to result fiom yttrium-enrichment of the interstitial grain boundary regions and may not be the result of being leached fiom the ceramic grains. It is possible that the manufacturing and firing procedures produce this grain-boundary enrichment in the V-I1 ceramic. Rodrigues et al. (1997) found that, during their preparatory firing, Y was extracted

Figure 4 represents the concentration Type (1) and Type (2) elements found in the two-week test solution (9OoC, initial pH 1). The y-axis is concentration in ppm and the x-axis is the material identification. Symbols have been offset, horizontally, where symbol overlap occurs.

impurities, such as the glass in V-I and such as impurities remaining fiom the chemical engineering in V-I1 are easily leached. Despite their small proportion in the solids, the Type 1 elements account for the highest measured leached concentrations. Thus, despite the relative insolubility of the ZrO2 297

into a soda-lime grain boundary glass phase that was purposefully introduced into their material. This resulted in higher yttria leachability with yttrium being leached fiom the glass rather than fiom the ZrO2 grains. However, the feldspathic glass-rich V-I ceramic does not show this effect; this might be due to the difference in glass composition (60% Si02 and no A1203 in Rodrigues' work) and different preparatory firing procedures. At 90°C and an initial pH of 3 (Fig 5), species concentrations are much lower than at pH 1. Data indicate that glass and impurity phases (Type 1 species) are still highly leachable under these conditions. The Zr/Y ratios suggest nonstoichiometry for dissolution of all of the ceramics; however, the V-I1 material is still most highly nonstoichiometric. In general Zr is more soluble in the pH 1, 9OoC environment then in the pH 3, 90°C environment. More Y is leached fiom the ceramics in the pH 3 experiment, except for the V-I11 solid, which is more highly leached at pH 1. After 12 weeks leaching at pH 1 and 9OoC, the V-I1 ceramic solid disaggregated into a course powder. At pH 3 and 90°C the V-I1 ceramic partially disaggregated to yield a crescent shape fiom the original ring-shaped solid. The V-I and V-I11 solids showed no visual signs of degradation. However, SEM images of V-I materials (e.g., Fig 2) do show leaching of glass.

5 OUR PRESENT EFFORTS Given these insights into the quality of the existing commercial YSZ ceramics, we have pursued the customizing of our own YSZ material, V-111, prepared by isostatic pressing of high-purity 8-mole % Y203 stabilized Zr02, starting with 99.99% pure powder commercially prepared by dehydrating nitrate-based gel. These were pressed fiom a fine powder (20.1 m2/g BET surface area); the shapes are then sintered in an electric furnace at 1600°C for 1 hour. This approach has the advantage of minimizing grain boundary phases of Type 1 elements and of insuring uniform yttria distribution. Note that the data in Figure 4 confirm that for the solid chunks, V-I11 is at least an order of magnitude lower in leached levels of Type 1 elements than V-I or V-11. However, the Zr/Y ratio still has a value of U-2 indicating that the Y leachability may still be a problem in long term sensors made in this way.

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6 CONCLUSIONS Our experiments show that while V-I bulk ceramics possess good short-term stability in high temperature, low pH environments, they would not be ideal for long term, down-hole use. With further understanding of the mechanisms that allow the Y and Zr to leach fiom the solid solutions, manufacturing sintering procedures can be changed to improve the stability of the yttria solid solution within the ceramic. As this paper goes to press, additional chemical leachability tests are in progress to assess the longer-term durability of the YSZ pH sensors. Certainly by comparison to glass pH electrodes, with more development the YSZ ceramic sensor can be made more ideal for use in highly corrosive, harsh environments. REFERENCES Field, M.P., J. T. Cullen, & R.M. Sherrell 1999. Direct determination of 10 trace metals in 50 pL samples of coastal seawater using desolvation micronebulization sector field ICP-MS. Journal of Analytical Atomic Spectrometry, 14:1425-1431. Lvov, S.N., G.C. Ulmer, H.L. Bames, D.D. Macdonald, X.Y. Zhou, S.M. Ulyanov, L.G. Benning, D.E. Grandstaff, M.F. Manna, & E. Vicenzi. Electrochemistry and structure of Yttria-stabilized Zirconia sensors for hydrothermal pH Measurements, Chemical Geology, in press. Nuccio, P.M. & M. Valenza 1993. Principles and methods of volcanic surveillance; the case of Vulcano, Italy. Zsotopic and geochemical precursors of earthquakes and volcanic eruptions; proceedings of an advisory group meeting UEATecdoc 726:108-115. Rodrigues, C. M., J.A. Labrincha & F.M.B. Marques 1997. Study of Yttria-Stabilized Zirconia-Glass Composites by Impedance Spectroscopy. Journal of the Electrochemical Society, Vol. 144, pp. 4303-4309. Roine, A. 1999. Outokumpu HSC Chemistry@ for Windows Vs. 4.1. Outokumpu Research Oy Information Service, Finland. Sato, M., J.G. Moore 1973. Oxygen and sulphur fugacities of magnetic gases directly measured in active vents of Mount Etna. Philosophical Transactions of the Royal Society of London, Series A: 274: 137- 146. Shock, E.L., D.C. Sassani, M. Wills & D. Sverjensky 1997. Inorganic species in geologic fluids: Correlations among standard molal thermodynamic properties of aqueous ions and hydroxide complexes. Ceochim. Cosmochim. Acta. 5 1~907-950. Stubican, V.S., G.S. Corman, J.R. Hellmann & G. Senft 1984. Phase relationships in some ZrOz systems. Advances in Ceramics 12: Science and Technology of Zirconia: 96- 107. Ulmer, G.C., M.F. Manna, D.E. Grandstaff, E.P. Vicenzi, H.L. Bames, S.N. Lvov, X. Zhou, & S.M. Ulyanov 2000. Impurity phases and the quest for a robust Zr02-based hydrothermal pH sensor. Applied Mineralogy. Balkema, Rotterdam: 79-82.

Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Gibbs free energies of formation of uranyl silicates at 298.15 K William F.McKenzie & Laurent Richard Department of Earth and Planetary Science, University of California, Berkeley, CA, USA

Sonia Salah Centre de Gkochimie de la Surface, Strasbourg, France

ABSTRACT: New estimates of the Gibbs free energies of formation at 298.15 K of uranyl silicates (uranophane, boltwoodite, sodium boltwoodite, kasolite, sklodowskite, cuprosklodowskite, weeksite, haiweeite, soddyite, synthetic sodium boltwoodite, and synthetic sodium weeksite) using the methodologies of Chermak & Rimstidt (1989, 1990) and Chen (1975) are reported. These results are compared to earlier estimates by Langmuir (1978), Hemingway (1982), Finch (1997) and Chen et al. (1999) as well as estimates based on experimental solubilities measured by Nguyen et al. (1992). 1 INTRODUCTION The thermodynamic properties of minerals containing uranium are of interest,as it is with the application of these properties together with the thermodynamic properties of other solid and aqueous species that one may predict the environmental consequences of introducing uranium into groundwater (e.g., at DOE research, production, and test sites or at the proposed high-level nuclear waste repository at Yucca Mountain, Nevada). These thermodynamic properties can also be used to validate the environmental consequences at natural analogue sites. A compilation and critical review of the chemical thermodynamics of uranium has been published (Wanner & Forest, 1992). This is an excellent source of thermodynamic data for many aqueous species and some solid hydroxy and carbonato compounds. Although there are data for U(1V) silicates [USi04(cr), USiO4(am)], there are none reported for uranyl [U(VI)] silicates in this compilation. Nguyen et al. (1992) have measured solubilities of four uranyl silicates [uranophane, soddyite, sodium boltwoodite (synthetic), and sodium weeksite (synthetic)] and from these solubilities have calculated Gibbs free energies of formation which included a set of assumptions. Langmuir (1978) estimated the Gibbs free energy of formation for uranophane by assuming uranophane is in equilibrium with calcite and schoepite under particular environmental conditions. Hemingway (1982) estimated the thermodynamic properties of uranophane and soddyite as well as those properties for boltwoodite, kasolite, sklodowskite, cuprosklodowskite, weeksite, and haiweeite. Finch (1997) and Chen et

al. (1999) estimated the thermodynamic properties for haiweeite, soddyite, and uranophane. 2 URANYL SILICATE MINERALS Stohl & Smith (1981) reported the crystal chemistry of ten uranyl silicate minerals and classified these minerals on basis of their uranium to silicon ratios: 1: 1 ratios include uranophane, beta-uranophane, boltwoodite, sodium boltwoodite, kasolite, sklodowskite, and cuprosklodowskite; 1:3 ratios include weeksite, and haiweeite; and the sole mineral with a 2: 1 uranium:silicon ratio is soddyite. Chemical formulas for these minerals are in Table 1. 3 LANGMUIR’S (1978) ESTIMATED GIBBS FREE ENERGY OF FORMATION FOR URANOPHANE

Based on equilibrium with schoepite and calcite Langmuir estimated the Gibbs free energy of uranophane from a chemical formula for uranophane then thought correct, Ca(U02)2(Si030H)a (Rosenzweig & Ryan, 1975) and. It was assumed that Pc02 = 10-3.5 atm and [H4Si040] = 10-3.4 moV1. The reaction can be written as: Ca(U02)2(Si030H)2 + C02(g) + 7H20 = CaC03 + 2 U02(0H)2.H20 + 2H4Si04 Langmuir adopted values for AG? for C02(g) [-94.254 kcal/mol] (CODATA, 1976), U02(0H)2.H20 [-390.4 kcal/mol] (estimated by 299

Langmuir from enthalpies of formation calculated from heats of solution measured by Cordfunke 1964, and Hemingway 1977; from the calculated equilibrium constant by Baes & Mesmer, 1976, for the reaction U02(OH)2 + 2 H+ = U022+ + 2 H20 and the free energy of formation for uranium trioxide from Gayer & Leider 1955; and from the observation by Nikitin et al. 1972, and Robins 1966, that U02(OH)2.H20 and U02(OH)2 are in equilibrium at 60OC), H4Si04 [-3 12.58 kcal/mole] (estimated by Langmuir), and H20 [-56.687 kcal/mole] (CODATA 1976). Although not specified, it is likely that Langmuir adopted the value for the AGP of calcite [-269.908 kcal/mol] as that from Robie & Waldbaum (1968). Langmuir’s estimated free energy of formation for uranophane from the above reaction and thermodynamic data was -1,189 kcal/mole. If one uses the chemical formula for uranophane now accepted (Stohl & Smith 1981) and writes the following reaction with uranophane in equilibrium with schoepite and calcite

and makes the same assumptions made previously by Langmuir (with respect to the partial pressure of carbon dioxide and the concentration of aqueous silica) together with the more newly accepted thermodynamic data (Johnson, et al., 1992), the resulting calculated Gibbs free energy of formation for uranophane is -6,168 kJ/mol or -1,474 kcal/mol. 4 HEMINGWAY’S (1982) ESTIMATED GIBBS FREE ENERGIES OF FORMATION FOR URANYL SILICATES Hemingway estimated free energies of formation for uranyl silicates by the method of Chen (1975) and from a procedure suggested by H. C. Helgeson (personal communication to H. P. Eugster) and applied by Eugster & Chou (1973). Hemingway modified the procedure suggested by Helgeson and plotted scaled free energies rather than scaled enthalpies. Without presenting the details of his calculations, Hemingway noted that the free energies for uranyl silicates estimated in this manner were equal to those obtained from summing the free energies for the oxide components. Hemingway further noted that the fact that this method does not provide a good estimate for the free energies of uranyl silicates reflects the structural dependence of the model. Nevertheless, Finch (1997) employed the method of summing oxides in his estimates (Table 1). Chen’s method is based on the postulation that the extrapolation of the calculated Gibbs free ener300

gies of formation of a mineral calculated from the sums of the Gibbs free energies from various sets of compounds with increasing complexity results in a reasonable estimate of the Gibbs free energy for that mineral. Chen ranked reactions, x, from 0 to n depending on their complexity with decreasing free energies. Reactions with values of free energies of approximately equal value were ranked equally. Chen then regressed the free energies as a function of x using the equation

where AGz 298 is the estimated free energy for a particular reaction at reaction order x, a and b are constants, and c is the extrapolated AG;

298

for the

mineral. Chen noted that if b < 0, AGz 298 approaches c as x becomes great in value; and, if b 2(uo2>2(sio4)2'3H20

I

k J rnol-'

AG22g8 -6168a -6211e -6213b -6211' -6213' -6 1 92g -59 14d

Kasolite Pb(U02)(Si04)*H20 Sklodowskite

-631gb -6110' Cuprosklodowskite -5827b CU[(UO~)~(S~O~OH)~]*~H~O -5830' Weeksite -9043b -9039' K2(U02)2(Si205)3'4H20 Haiweeite -9396b -9395' Ca(UO2)2(Si205)35H20 -9329d Soddyite -36Ub -3706' (U02)2( Si04)-2H20 -3671d Sodium Boltwoodite (synthetic) ~] = - 1430.9 kJ mol-'. These results deviate considerably from the data recommended by both Wagman et al. (1982) and Robie & Hemingway (1995). The thermodynamic properties of azurite and malachite have been reinvestigated by performing solubility measurements as a function of temperature. The solubility experiments have been carried out by means of the pH - variation method (Schindler 1963, Gamsjager et al. 1965). Details about the experimental procedures and data analysis are beyond the scope of this contribution and will be reported elsewhere (Preis & Gamsjager, in prep.). The aim of this work is to reveal severe discrepancies between experimental observations in the system Cu2+- H2O - CO2 and theoretical predictions using literature values for the thermodynamic properties of malachite and azurite. Revised thermodynamic quantities for the copper carbonates are presented. These data are applied to calculate predominance diagrams in order to compare our values with the thermodynamic properties given in literature. 2 DETERMINATION OF THERMODYNAMIC DATA FOR COPPER CARBONATES The dissolution reactions for malachite and azurite in aqueous solutions of ionic strength I can be written as

329

1/2 CU~(OH)~CO~(S) + 2 H+(I) + Cu2+(1)+ 112 C02(g) + 312 H2O(I)

of the experimental data yielded for the dissolution of one mole copper at T = 298.15 K:

(1)

AsOIH@(malachite) = (- 27.8 f 3.0) kJ mol-', As,,H@(azurite) = (- 21.9 f 3.0) kJ mol-', As,lS@(malachite) = (28 f 10) J mol-' K-', AS,@(azurite) = (47 f 10) J mol-' K'.

and

+ 2 H+(I) + 113 CU~(OH)~(CO~)~(S) Cu2+(I)+ 213 CO&) + 413 H20(I) ,

(2)

respectively. The pertinent solubility constant valid at a certain temperature and ionic strength of the aqueous medium reads

'K' PSO

= ([CU"

]/ m")(p(C0,) / p"}Y ([H+]/m")'

3 DISCUSSION

(3)

with square brackets, p(C02), 'm and p @ denoting the molalities of the respective ionic species in solution, the partial pressure of carbon dioxide, the standard molality (1 mol kg-') and the standard pressure (1 bar). The exponent y is equal to 1/2 and 2/3 for malachite and azurite, respectively. The stoichiometric solubility constant can be extrapolated to zero ionic strength (infinite dilution) by correcting for the ionic strength dependence of the individual activity coefficients, y, as well as the activity of water, 43201,

where x amounts to 1 and 2/3 for malachite and azurite, respectively. The enthalpy of dissolution at T,. = 298.15 K is obtained from the temperature dependence of the solubility constant (5) Combining the enthalpy of dissolution with the solubility constant results in an expression for the entropy of dissolution

at T,. = 298.15 K. Finally, the standard enthalpy of formation AfB@ and the standard entropy S@ of the respective copper carbonates are obtained from As0lH@,AsolS@,and appropriate auxiliary data for Cu2+,H20(1) and CO& which can be found in the CODATA tables (Cox et al. 1989). The solubility constants of malachite and azurite were determined in perchlorate media from 288.15 to 338.15 K at constant ionic strength I = 1.00 mol kg-' NaC104. Recently, the ionic strength dependence of the solubility constant of zinc carbonate has been reliably modeled up to I = 3.00 mol kg'' NaC104 (Preis et al. 2000) by means of the specific ion-interaction theory (Grenthe et al. 1997). Hence, this electrolyte model was likewise applied to the extrapolation of the solubility constants of malachite and azurite to zero ionic strength. A proper analysis

Using the set of thermodynamic data for azurite, malachite and tenorite (CuO), recommended by Robie & Hemingway (19952), predominance diagrams for the ternary system Cu + - H20 - CO2 can be constructed. The diagrams shown in Figures l a and l b are valid at 298.15 K and 323.15 K, respectively, with log p(C02) plotted versus pH (H+ - activity), at a(Cu2+)= 104. The partial pressure of carbon dioxide for the equilibrium between malachite, tenorite, gas phase and aqueous medium would amount to 3.1 bar at 323.15 K. Thus, CuO should be stable at a CO;! pressure around 0.84 bar (dashed line in Figure lb). However, we were able to transform tenorite into malachite at p(C02) = 0.84 bar and T = 323.15 K, which clearly disproves the prediction using the data tabulated by Robie & Hemingway (1995). Moreover, azurite could be prepared in a highpressure autoclave at p(C02) = 50 bar and T = 298.15 K, following a procedure given by Brauer (1962). According to Figure la, malachite should be the most stable solid phase at these conditions. The formation of azurite would require a CO;! pressure of approximately 8 x 107 bar which is not only unrealistic but again disproved by the conditions of azurite synthesis. As the thermodynamic data for CuO are consistent with the JANAF tables (Chase 1998) and are regarded to be fairly reliable, some of the literature values for the thermodynamic properties of malachite and azurite must be wrong. Using the data compilation of Robie & Hemingway (1995), the enthalpies of dissolution [A,,IH@(malachite) = - 33.6 kJ mol-' and A,,~R@(azurite)= - 34.5 kJ mol-'1 deviate considerably from our results, whereas the entropies of dissolution [As,lS@(malachite) = 30.7 J mol-' K-' and A,,lS@(azurite) = 53.0 J mol" K-'1 are consistent with our solubility experiments within the limits of experimental error. Hence, it can be concluded that in contrast to the values for AfH@of malachite and azurite listed by Robie and Hemingway (1995) the standard entropies, based on low-temperature measurements of heat capacities (Kiseleva et al. 1992), are confirmed by our investigations. The literature data for the standard enthalpies of formation should be replaced by the following recommended values: A ~ H @ [ C U ~ ( O H ) ~=C -~ ~1067.1 ] kJ mol-', and A~B@[CU~(OH>~(CO~);!] = - 1675.1 kJ mol''.

330

Figure 1. Predominace diagrams for the system Cu2+- HzO - COz at a(Cu2') Robie & Hemingway (1995). (a) T = 298.15 K, (b) T = 323.15 K.

The application of the revised thermodynamic properties of malachite and azurite leads to predominance diagrams for the system Cu2+- H20 - C02, shown in Figure 2, which deviate remarkably from those depicted in Figure 1. As the standard Gibbs free energies of formation for the complex Cu(C03)2*-(aq) as well as the solid phase CuCO3 (neutral copper carbonate) are exclusively available at 298.15 K (Schindler et al. 1968, Reiterer et al. 198l), these species are included in our thermodynamic model for the calculation of Figure 2a only. It is worth mentioning that the equilibrium partial pressure for the coexistence of azurite, malachite, the gas phase and the aqueous medium increases slightly with increasing temperature with values around 1 bar. The CO2 pressure for the equilibrium between malachite, tenorite, the gas phase and the aqueous medium increases considerably when the temperature is raised. At 323.15 K and p(C02) = 0.84 bar tenorite is expected to be unstable (see dashed line in Figure 2b) which was proved by our experimental observation where tenorite was transformed into pure malachite at these conditions. Ac-

=

10-4using thermodynamic data selected by

cording to our thermodynamic model, the formation of azurite in aqueous media occurs at equilibrium partial pressures of carbon dioxide around 1 bar. Well-crystallized azurite was prepared at 50 bar and 298.15 K which provides further evidence for the reliability of our revised set of thermodynamic data for Cu2(OH)2C03 and CU~(OH)~(CO&. The standard enthalpies of formation of the copper carbonates based on calorimetric methods (Roth et al. 1941, Richardson & Brown 1974, Kiseleva et al. 1992) are clearly disproved as outlined above. The determination of the standard enthalpy of formation from solubility measurements as a function of temperature avoids all systematic uncertainties which may occur during calorimetric measurements of the enthalpy of dissolution due to the evolution of unknown quantities of carbon dioxide. Moreover, the decomposition experiments on malachite and azurite in a drop calorimeter may be accompanied with systematic errors owing to a poorly defined thermodynamic state of the copper oxide formed rapidly during the decarbonation reaction.

Figure 2. Predominace diagrams for the system Cu2+- HzO - CO2 at a(Cu23= 10-4according to the revised set of therrnodynamic properties of the copper carbonates. (a) T= 298.15 K, (b) T= 323.15 K.

331

4 CONCLUSION

of malachite [CU~CO,(OH)~]. U S . Bur. Mines Rep. Inv. 7851: 1-5. Robie, R.A. & B.S. Hemingway 1995. Thermodyanmic properties of minerals and related substances at 298.15 K and I bar (105 Pascals) pressure and at higher temperatures. Washington: U.S. Geological Survey Bulletin 213 1. Roth, W.A., H. Berendt & G. Wirths 1941. Die Bildungswarme einiger mineralischer und kiinstlicher Carbonate. 2. Elektrochem. 47: 185-190. Schindler, P. 1963. Die Bestimmung der Loslichkeitskonstanten von Metalloxiden und - hydroxiden. Chimia 17: 313330. Schindler, P., M. Reinert & H. Gamsjager 1968. Zur Thermodynamik der Metallcarbonate - 2. Mitt.: Loslichkeitskonstanten und freie Bildungsenthalpien von Cu2(0H),C03 (Malachit) und C U ~ ( O H ) ~ ( C O(Azurit) ~ ) ~ bei 25'. Helv. Chim. Acta 51: 1845-1856. Wagman, D.D., W.H. Evans, V.B. Parker, R.H. Schumm, I. Halow, S.M. Bailey, K.L. Churney & R.L. Nuttall 1982. The NBS tables of chemical thermodynamic properties. J. Phys. Chem. Re$ Data 11, Supplement 2.

Literature data for the thermodynamic properties of malachite, azurite and tenorite were applied to construct predominance diagrams for the ternary system Cu2+- H20 - Col. The prediction of the partial pressure of carbon dioxide for the phase equilibrium between malachite and tenorite in aqueous media was disproved by the experimental observation that tenorite can be transformed into pure malachite at p(C02) = 0.84 bar and T = 323.15 K. According to the data selected by Robie & Hemingway (1995), azurite is unstable in aqueous solutions at 298.15 K and partial pressures of CO2 below 8x10' bar. However, we were able to prepare azurite in an aqueous medium at room temperature and CO2 pressures around 50 bar. The thermodynamic properties of malachite and azurite were reinvestigated by performing solubility measurements as a function of temperature at constant ionic strength. A proper analysis of the experimental data confirmed the reliability of the literature values for the standard entropies. Revised values for the standard enthalpies of formation are recommended: A~H@[CU~(OH)~CO~] = - 1067.1 kJ mol-' and A~B@[CU~(OH)~(CO~)~] = - 1675.1 kJmol-'. REFERENCES Brauer G. 1962. Handbuch der prapurativen anorganischen Chemie, 2"d edition. Stuttgart: Enke. Chase M.W., Jr. 1998. NIST-JANAF Thermochemical Tables, Part II, Cr-Zr, 4" edition. Gaithersburg: American Chemical Society. Cox, J.D., D.D. Wagman & V.A. Medvedev 1989. CODATA Key Valuesfor Thermodynamics.Washington: Hemisphere. Gamsjager, H., H.U. Stuber & P. Schindler 1965. Zur Thermodynamik der Metallcarbonate - 1. Mitt.: Loslichkeitskonstanten und freie Bildungsenthalpie von Cadmiumcarbonat, ein Beitrag zur Thermodynannk des ternlen Systems Cd",,, - H,O,,) - CO,,,). Helv. Chim. Acta 48: 723-729. Grenthe, I., A.V. Plyasunov & K. Spahw 1997. Estimations of medium effects on thermodynamic data. In I. Grenthe & I. Puigdomenech (eds), Modelling in aquatic chemistry: 325426. Paris, OECD NEA. Kiseleva, I.A., L.P. Ogorodova, L.V. Melchakova, M.R. Bisengalieva & N.S. Becturganov 1992. Thermodynamic properties of copper carbonates - malachite Cu2(0H)2C03 and azurite C U ~ ( O H ) ~ ( C OPhys. ~ ) ~ . Chem. Minerals 19: 322-333. Preis, W., E. Konigsberger & H. Gamsjager 2000. Solid-solute phase equilibria in aqueous solution. XII. Solubility and thermal decomposition of smithsonite. J. Solution Chem. 29: 605-618. Preis, W. & H. GamsjSLger, in prep. Solid-solute phase equilibria in aqueous solution. XV. Thermodynamic properties of malachite and azurite - predominance diagrams for the system Cu2+- H 2 0 - CO2.J. Chem. Thermodynamics. Reiterer, F., W. Johannes & H. Gamsjager 1981. Semimicro determination of solubility constants: copper(I1) carbonate and iron(I1) carbonate. Microchim. Acta I: 63-72. Richardson, D.W. & R.R. Brown 1974. Enthalpy of formation

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Water-Rock Interaction 2001, Cidu (ed.), 0 2007 Swets & Zeitlinger, Lisse, ISBN 90 2657 824 2

Thermodynamic calculation of the distribution of organic sulfur compounds in crude oil as a function of temperature, pressure, and H,S fugacity L.Richard & H.C .Helgeson University of California, Berkeley, USA

ABSTRACT: Recent estimates of the thermodynamic properties of organic sulfur compounds of geochemical interest have been used to calculate their relative stabilities as a function of temperature and pressure, and the extent to which they interact with H2S and other sulfur-bearing aqueous species or minerals at the oil-water interface in sedimentary basins. The distribution of organic sulfur compounds in crude oil predicted from such calculations compares favorably with compositional data reported in the literature. pounds with oil maturity, knowledge of the thermodynamic properties of these compounds is requisite.

1 INTRODUCTION Combustion of the sulfur present in fossil fuels contributes to a number of environmental hazards, not the least of which is acid rain. Therefore, understanding the behavior of organic sulfur compounds in petroleum systems is of fundamental importance. The purpose of this communication is to summarize a thermodynamic approach to the problem, which can be used to guide experimental investigations.

3 THERMODYNAMIC PROPERTIES OF ORGANIC SULFUR COMPOUNDS

2 ORGANIC SULFUR COMPOUNDS IN PETROLEUM SYSTEMS The organic sulfur compounds present in immature oils and bitumens are primarily alkylthiolanes, alkylthianes and alkylthiophenes, whereas most of the organic sulfur present in mature crude oils is accounted for by methylated benzothiophenes and dibenzothiophenes (Orr & Sinninghe Damsti: 1990). The sulfur content of crude oils typically ranges between 0.5 and 0.9 weight percent, but can reach 10 percent or more (Tissot & Welte 1984, O n & Sinninghe Damsti: 1990). Since the publication of the classical paper by Gransch & Posthuma (1974), it is generally accepted that the sulfur content of a crude oil is determined primarily by the sulfur content of its source kerogen (On & Sinninghe Damsti 1990). Alternatively, sulfurization of petroleum can occur through reactions between hydrocarbons and abundant H2S produced at relatively high temperatures by thermochemical sulfate reduction (On 1974). In order to quantify such processes as well as to characterize from a thermodynamic point of view the change in the distribution of organic sulfur com-

The standard molal thermodynamic properties and heat capacity power function coefficients of more than 100 organic sulfur compounds of geochemical interest have been calculated by regressing experimental data reported in the literature with carbon number systematics, and the properties and coefficients of many more can be estimated using group additivity algorithms such as those developed by Helgeson et al. (1998) and Richard & Helgeson (1998). These thermodynamic properties for organic sulfur compounds, which include n-alkyl, branched, cyclic, and aromatic thiols, sulfides, and disulfides, as well as thiophenic compounds, thianthrene and carbon disulfide will be published in a forthcoming paper (Richard 2001). These properties can be used in calculations such as those described below to predict the distribution of organic s u l h compounds in crude oil as a function of temperature, pressure, and H2S fugacity. 4 CALCULATION OF THE DISTRIBUTION OF

ORGANIC SULFUR COMPOUNDS IN MATURE CRUDE OILS Following the approach used by Helgeson et al. (1 993) to characterize metastable equilibrium between hydrocarbons, oil field waters, CO2 gas, and authigenic mineral assemblages, a general disproportionation reaction describing metastable equilib-

333

wheref; and ai stand for the fugacity and the activity of the subscripted gas and liquid species, respectively. If in a first approximation we adopt the hypothesis that the activities (ai) of the liquids are essentially equivalent to their mole fractions (xi)in crude oil, Equation 2 can be rewritten under a logarithmic form as

+ 18.5log f H 2 S ( g ) - log K

I/

18.5. (3)

From compositional data reported by Mair (1964), the mole fractions of n-decane and toluene in crude oil are estimated to be 0.023 and 0.012, respectively. These values have been used together with Equation 3 to calculate the mole fractions of the organic sulfur compounds (I-VI) in equilibrium with n-decane and toluene as a function of H2S fugacity at 100°C and 400 bars. The results of these calculations have been plotted in Figure 2. It can be seen in this figure that the mole fractions of the organic sulfur compounds increase with increasing H2S fugacity, which is consistent with the sulfurization Figure 1. Idealized structures of the organic compounds conprocess of crude oils suggested by Orr (1974). It can sidered in the calculations: (I) ethanethiol, (11) 2-thiabutane, also be deduced from Figure 2 that with the excep(111) thiacyclohexane, (IV) truns-2,5-dimethylthiacyclopentane, tion of 2-methylthiophene, the various organic sulfur (V) 2-methylthiophene, (VI) 2,4-dimethylbenzo[b]thiophene, compounds reach appreciable concentrations at fit(VII) toluene, and (VIII) n-decane. gacities of H2S of the order of 10 bars, which are in the range of those expected for reservoirs containing rium between a given organic sulfur compound both oil and sour gas (Orr, 1977). C,H,$, n-decane (C10H22), toluene (C7H8), and hydrogen sulfide (H2S) gas may be written as

-

where (I) and (g) denote the liquid and gas states, respectively. The equilibrium constant K of Reaction 1 has been calculated at 100°C and 400 bars with an updated version of the SUPCRT92 software package (Johnson et al. 1992) for six representative organic sulfur compounds, including ethanethiol, 2-thiabutane, thiacyclohexane, trans-2,5-dimethylthiacyclopentane, 2-methylthiophene, and 2,4-dimethylbenzo[b]thiophene. The structures of these compounds are shown in Figure 1, together with those of n-decane and toluene. The law of mass action for Reaction 1 can be expressed by

Figure 2. Logarithm of the mole fractions of the organic sulfur compounds depicted in Figure 1 in equilibrium with liquid ndecane and toluene as a function of the logarithm of the fugacity of H2S gas at 100°C and 400 bars (see text).

2Sg) .a 5.5n-2.5m+5 c7H8(/)

=

18.5 2n-1.75rn+3.5 aC,HmS(,, * %0%2(/)

(2)

334

Figure 3. Comparison between the concentrations of organic sulfur compounds (expressed as the logarithm of their mole in the API Research Project 48 crude oil fraction Log XOSC,~) from Wasson, Texas (Rall et al. 1972), and those calculated from Equation 3 with LogfH2S = 1.15 at 100°C and 400 bars (see text).

The comparative histogram shown in Figure 3 indicates that for a value of the H2S fugacity of 14 bars, our calculated mole fractions agree reasonably well with those reported by Rall et al. (1972) for the Wasson, Texas crude oil. This result confirms the hypothesis that metastable equilibrium states between organic sulfur compounds, liquid hydrocarbons, and gas H2S are established at petroleum reservoir conditions and corroborates the thermodynamic approach used in our calculations.

-

5 CONCLUDING REMARKS

assemblages: Are they in metastable equilibrium in hydrocarbon reservoirs? Geochimica Cosniochiinica Acta 57( 14): 3295-3339. Helgeson, H.C., Owens, C.E., Knox, A.M. & Richard, L. 1998. Calculation of the standard molal thermodynamic properties of crystalline, liquid, and gas organic molecules at high temperatures and pressures. Geochimica Cosmochimica Acta 62(6): 985-1081. Johnson, J.W., Oelkers, E.H. & Helgeson, H.C. 1992. SUPCRT92: A software package for calculating the standard molal thermodynamic properties of minerals, gases, aqueous species, and reactions from 1 to 5000 bar and 0 to 1000°C. Computers and Geosciences 18(7): 899947. Mair, B.J. 1964. Hydrocarbons isolated from petroleum. Oil and Gas Journal 62(37): 130-134. Orr, W.L. 1974. Changes in sulfur content and isotopic ratios of sulfur during petroleum maturation - Study of Big Horn Basin Paleozoic oils. American Association of Petroleum Geologists Bulletin 58( 11): 2295-23 18. Orr, W.L. 1977. Geologic and geochemical controls on the distribution of natural gas. In R. Campos & J. Goiii (eds), Advances in Organic Geocheinistiy 1975: 57 1-597. Madrid: Empresa Nacional Adaro de Investigaciones Mineras. Orr, W.L. & Sinninghe Damste, J.S. 1990. Geochemistry of sulfur in petroleum systems. In W.L. Orr & C.M. White (eds), Geochemistiy of Sulfur in Fossil Fuels: 2-29. Washington DC: American Chemical Society. Rall, H.T., Thompson, C.J., Coleman, H.J. & Hopkins, R.L. 1972. Sulfur compounds in crude oil. US Bureau of Mines Bulletin 659, 187 p. Richard, L. 2001. Calculation of the standard molal thermodynamic properties at high temperatures and pressures of crystalline, liquid, and gas organic sulfur compounds of geochemical interest. Geochimica et Cosmochimica Acta. Richard, L. & Helgeson, H.C. 1998. Calculation of the thermodynamic properties at elevated temperatures and pressures of saturated and aromatic high molecular weight solid and liquid hydrocarbons in kerogen, bitumen, petroleum, and other organic matter of biogeochemical interest. Geochimica Cosmochimica Acta 62(23/24): 3591-3636. Tissot, B.P. & Welte, D.H. 1984. Petroleuiii Formation and Occurrence. Berlin: Springer-Verlag.

Estimates of the thermodynamic properties of some representative organic sulfur compounds have been used to calculate the speciation of organic sulfur in crude oil as a function of H2S fugacity at 100°C and 400 bars. The calculated mole fractions of the organic sulfur compounds are in good agreement with compositional data reported in the literature. Such theoretical calculations have important implications for more comprehensive studies of the interactions between organic and inorganic sulfur species in diagenetic and hydrothermal systems, as well as for the global geochemical cycle of sulfur.

REFERENCES Gransch, J.A. & Posthuma J. 1974. On the origin of sulfur in crudes. In B. Tissot & F. Bienner (eds), Advances in Organic Geochemistry 1973: 727-739. Paris: Technip. Helgeson, H.C., Knox, A.M., Owens, C.E. & Shock E.L. 1993. Petroleum, oil field waters, and authigenic mineral

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Measurement of quartz dissolution rates with a flow-through type autoclave reactor H .Sugita, 1.Matsunaga & T.Yamaguchi Geotechnology Department, National Institute for Resources and Environment, Japan

H .Tao Hydrospheric Environmental Protection Department, National Institute for Resources and Environment, Japan

ABSTRACT: The kinetics of quartz dissolution was investigated using a flow-through type autoclave reactor. The rate of quartz dissolution by distilled water was determined at various conditions by changing the following parameters: the mass and the particle diameter of quartz, the temperature and the flow rate. The apparent dissolution rate coefficient was independent of the mass of quartz, but the silica concentration in the effluent increased with increasing mass of quartz. The effect of particle diameter on quartz dissolution was described by the total surface area of quartz. The apparent dissolution rate coefficient increased with both increasing temperature and flow rate. dian diameters of the sieved quartz fractions were measured by a light scattering method and found to be 391, 228, 103 and 31.4 pm. Fine-grained quartz separated from the granite core obtained in well HDR-3 at Hijiori, Japan was also used. The median diameter of the separated quartz particles was 76 pm. The former quartz was used in the experiments on the effects of the mass and the particle diameter of quartz. The latter quartz was used to study the effects of temperature and flow rate. These quartz particles were washed with distilled water under an ultrasonic beam and then dried at 70U for about 1 day before being packed in the reactor.

1 INTRODUCTION Dissolution reaction of silica component from rocks to fluid is one of important processes for rock-water interaction in geothermal hot reservoirs and silica is one of major components in geothermal fluid. The silica concentration in geothermal fluid is generally useful as one of geochemical thermometers. On the other hand, in geothermal application plants, silica scale often forms from geothermal fluid and causes the plugging of pipelines and permeable layers and the declines in the performance of heat exchangers and the reducibility of fluid. Electric power plants using hot dry rock systems that are used near river water as heat exchange medium, is also expected to suffer the silica scale problems. Therefore it is very important to understand the kinetics of silica dissolution and precipitation. In this study, quartz dissolution rate to distilled water was investigated on the effects of the mass and the particle diameter of quartz, temperature and flow rate using the flow-through type reactor in a autoclave.

2.2 Apparatus and measurements

2.1 Quartz particles

The reactor set in the autoclave consists of titanium tube (with an internal diameter of 10.2-10.7 mm and an outer diameter of 12.7 mm), titanium filters, and stainless steel connectors. The zirconia particles were packed as dispersers on both sides of a packed bed of quartz. The lengths of the titanium tubes were adjusted in proportion to the mass of packed quartz. Distilled water saturated with argon gas was passed from the bottom toward the top of the reactor. The effluent was continuously sampled in a glass bottle and weighed to determine the actual average flow rate. The silica concentration in the effluent was measured with a spectrophotometer using the molybdate yellow method.

Two kinds of quartz were used in our experiments. One was quartz sand from Bahia, Brazil, supplied by Mitsubishi Material Corporation. The quartz sand (99.9% SO,) was sieved with mesh sizes 350-425, 180-250, 75-106 pm and less than 75 pm. The me-

2.3 Experimental conditions Three series of runs are conducted. The conditions of these experiments were shown in Table 1. In this table, D,,Wqzo,T, and Q are the median particle diame-

2 EXPERIMENTAL

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ter, the initial mass of packed quartz, the temperature and the flow rate, respectively. In all runs, the quartz particles in the reactor were washed at the experimental temperature with distilled water for one day, since it is generally thought that quartz is often covered with a very thin layer of amorphous silica (O’Connor & Greenberg, 1958). Table 1. Experimental conditions. Series Quartz D,X 106 W q z ~ ~ x 1 0 3 T No. kind m kg K ~~

Q x109 m3 s-’ 1.67 1.67 0.333 - 16.7

1 B* 103 3-12 473 6,12 473 B* 31.4-391 2 398-523 H** 76 12 3 * Quartz sand produced in Bahia ** Quartz separated from the granite core in well HDR-3

3 RESULTS AND DISCUSSION 3.1 Effect of mass of packed quartz

In Series 1, the silica concentration in the effluent, C, was almost constant after 30 hours since T and Q were kept at fixed values. The values of C in the range of about 30 to 70 hours were averaged, and their means were treated as steady state values. C increased with increasing W,, though not proportionally. C is plotted against the residence time in the packed bed of quartz, Z, in Figure 1. Z was calculated on the basis of the initial porosity of the packed bed of quartz (45.4%) and the average mass of quartz remaining in the reactor after 30 to 70 hours, Wqz. C increases with increasing Z but the relationship between C and Tseems to be curved rather than linear.

3.2 Analysis of dissolution kinetics Many investigators have studied the kinetics of quartz dissolution and precipitation and suggested some rate equations (O’Connor & Greenberg, 1958; Rimstidt & Bames, 1980; Brady & Walther, 1990; Dove & Crerar, 1990; Tester et al., 1994; Worley et al., 1996; Johnson et al., 1998). The results obtained in this study are modeled by equations 1 and 2 for plug-flow reactor: dC/dt = kAs(Ce- C)

(1)

where C is the silica concentration in the effluent [kg mf-3], k is the overall rate constant, A, is the surface area of quartz per unit liquid volume[m$ mf-’1, Ce is the equilibrium concentration of silica and equivalent to quartz solubility [kg mf-3]. Z is the residence time in the packed bed of quartz [s]. For separating the effects of dissolution and precipitation reactions, the following equations were used (O’Connor & Greenberg, 1958):

Finally, the net rate of change in the silica concentration is represented by equation 5:

where kl is the apparent dissolution rate coefficient [kg m,-2 s-’], and k2 is the apparent precipitation rate coefficient mf3 mSp2s-’]. The equilibrium constant K [kg mf- ] is given by equation 6, which upon substitution into equation 5 gives equations 7 and 8:

5

In our analysis, C e was obtained by using equation 9 (Fournier & Potter, 1982): log S,, = - 1309/T+5.19

(9)

where S,, is the quartz solubility [mg kg-’] and T is the absolute temperature [K]. Table 2 shows the data on the dissolution and precipitation rates of quartz. A , was 6.99-7.07 X 104rn; mfP3in Series 1. Neither kl nor k2 were affected by W,, within experimental error. The mean values of kl and k2 are 0.857X l O P 9 kg rns-2 s-l and 3.23 X lO-’ mf3 m,-2 s-l respectively. The errors for both kl and k2 are less than 5%.

Figure 1. Change in silica concentration in the effltient with residence time in a packed bed of quartz.

338

tion 7 and plotted against the flow rate at each temperature in Figure 4. kl increases with T at any Q. Also, kl increases with Q at any T. This is considered to reflect the effect of the liquid velocity in the packed bed rather than that of the flow rate. The diffusion layer develops on the surface of the quartz particles when the liquid velocity is low. On the other hand, the effect of the diffusion layer becomes smaller when the liquid velocity becomes higher, and finally the apparent dissolution rate coefficient should reach a constant value.

Table 2.Data on rates of dissolution and precipitation of quartz. wq,x103 cx103 ~ 4 ~ x 1 0klx109 ~ ~ _ _ _ _ s-' kg ms-2s-1 kg kg mf-3 0.849 0.224 3.01 31.7 0.880 0.235 5.86 60.5 0.819 0.218 8.77 80.2 0.882 109 0.233 12.0 Average

0.857

kzx109 ~ mf3ms-*s-' 3.20 3.32 3.09 3.33 3.23

3.3 Effect of particle diameter of quartz The results of Series 1 and 2 are shown in Figure 2. In this figure, C is plotted against the total surface area of the quartz in the reactor, St. Circles in this figure show the results of Series 1and other symbols show the results of Series 2. The results of Series 1 and 2 fall on the same line. This means that the effect of the particle diameter of quartz on the dissolution may be described by S,. This quantity is included in the factor& in the rate equations.

Figure 3. Silica concentration in effluent vs. flow rate at each temperature. * F&P: quartz solubility obtained from equation 9

Figure 2. Silica concentration in effluent vs. total surface area of packed quartz.

3.4 Effect of flow rate and temperature The results of Series 3 are shown in Figure 3. In this figure, C is plotted against Q at each T. At any T, C decreases with increasing Q, i.e. increasing Q is equivalent to shortening Z . C increases with increasing T at any Q. C at Q = 0 may be equal to the saturation concentration of silica, the quartz solubility at the given temperature. The quartz sohbilities obtained using equation 9 are plotted as open circles in Figure 3. These plots may be consistent with the plots for Series 3. Therefore these values of the quartz solubility are treated as the saturated silica concentration at each temperature. The quartz dissolution kinetics was analyzed for Series 3 in a similar way as for Series 1 and 2. The values of k , for Series 3 were calculated using equa-

Figure 4. Apparent dissolution rate coefficient of quartz vs. flow rate at each temperature.

3S Apparent activation energy Figure 5 shows the plots of kl' against the reciprocal of temperature at each liquid velocity v. kl' is the apparent dissolution rate coefficient, obtained by converting kl into mole units. These plots are fitted 339

by the following Arrhenius equation lines using a least square method:

where A is the frequency factor, A E is the apparent activation energy of quartz dissolution [J mol-'I, R is the gas constant (8.314J.mol-'-K-'). A E obtained from equation 10 is plotted against v in Figure 6. A E becomes lower with increasing v when v is under 0.2 X 1OP3 m s-'. When v is exceeds this value, A E seems to be constant, 42-43 kJ mol-'. This tendency was also seen in the apparent activation energy of quartz precipitation in our experiments. The activation energy of quartz precipitation was 17-18 kJ rnol-' when v is over 0.2 X 10-3m s-'.

4 CONCLUSIONS In order to analyze the behavior of silica dissolution on water-rock interaction, the dissolution rate of quartz by distilled water was examined using a flowtype reactor. The following results were obtained. 1)The apparent dissolution rate coefficient of quartz is not affected by the mass of packed quartz, but the silica concentration in the effluent increases with increasing mass of packed quartz. 2) It was confirmed that the kinetics of quartz dissolution follows the equations below: dCldt = kp4,(1- CIC,) or dCl& = kl A, -k2C A,

3) The effect of the particle diameter of quartz on the apparent dissolution rate coefficient can be represented by the total surface area of packed quartz. 4) The silica concentration in the effluent increases with increasing temperature and liquid velocity. 5 ) The apparent dissolution rate coefficient of quartz becomes higher as the temperature and the liquid velocity become higher. 6) The apparent activation energies of quartz dissolution and precipitation both become lower with increasing liquid velocity when the liquid velocity is under 0.2X l O P 3 m s-'. When the liquid velocity is over the above value, the activation energies of quartz dissolution and precipitation seem to be constant, 42-43 and 17-18 kJ mol-', respectively. REFFERENCES

Figure 5. k , ' vs. the reciprocal of temperature at each liquid velocity.

Brady, P.V. & Walther, J.V. 1990. Kinetics of quartz dissolution at low temperatures. Chemical Geology 82: 253-264. Dove, P.M. & Crerar, D.A. 1990. Kinetics of quartz dissolution in electrolyte solutions using a hydrothermal mixed flow reactor. Geochimica et Cosmochimica Acta 54: 955-969. Fournier, R.O. & Potter, R.W.Jr. 1982. An equation correlating the solubility of quartz in water from 25 to 900°C at pressure up to 10000 bars. Geochim.Cosmochim. Acta 46: 1969-1973. Johnson, J. W., K. G. Knauss, W. E. Glassley, L. D. DeLoach & Tompson, A. F. B. 1998. Reactive transport modeling of Plug-flow Reactor Experiments : Quartz and Tuff Dissolution at 240'c. Journal ofHydrology 209: 81-111. O'Connor, T.L. & Greenberg, S.A. 1958. The kinetics for the solution of silica in aqueous solutions. Journal of physical chemistry, 62, 1195-1198 Rimstidt, J. D. & Barnes, H.L. 1980. The kinetics of silicawater reactions. Geochim.Cosmochim. Acta 44: 1683-1699. Tester, J.W., W.G. Worley, B. A. Robinson, C.O. Grigsby, & Feerer, J.L. 1994. Correlating quartz dissolution kinetics in pure water from 25 to 625'c. Geochim.Cosmochim. Acta 58: 2407-2420. Worley, W. G., J.W. Tester & Grigsby, C. 0. 1996. Quartz dissolution kinetics from 100-20Ooc as a function of pH and ionic strength. AIChE Journal 42: 3442-3457.

Figure 6. Apparent activation energy of quartz dissolution vs. liquid velocity in a packed bed of quartz

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Experimental study of rocWwater/C02 interaction at temperatures of 100 - 350°C YSuto, L.Liu & T.Hashida Fracture Research Institute, Tohoku University, Japan

N.Tsuchiya & N.Yamasaki Department of Geoscience and Technology, Tohoku University, Japan

ABSTRACT: In order to reduce the amount of CO2 emitted to the atmosphere, injection of CO2 into underground is considered one of the useful approaches. Although all reservoirs under investigation to date for CO2 disposal have temperatures below 100°C and most are in sedimentary rocks, existing data of CO2 solubility in water suggests that hotter and deeper rock masses may also be candidates for underground CO2 injection due to enhanced CO2 solubility under high pressures even at high temperatures..We conducted some experimental studies of rocldwaterlC02 interaction on granite and sandstone at temperatures of 100-350°C using a batch type autoclave. It was shown that the presence of CO2 promoted dissolution of the rocks, especially alkali metals and alkaline-earth metals, and deposition of some kinds of aluminosilicate, compared with those in experiments without C02. 1 INTRODUCTION Recently, underground CO2 injection has been proposed to reduce the amount of CO2 emitted to the atmosphere, because underground injection is now proven to be feasible. For example, in Europe the Joule I1 project (Holloway 1996) was undertaken to assess the potential of underground disposal, by injection into a subsurface aquifer, of CO2 emitted from a fossil fuel power plant. The Joule I1 project assessed the amount of CO2 emitted from the plant, CO2 storage ability of the reservoir, stability and safety of the reservoir, undertook reservoir modeling and geochemical experiments, and concluded that “underground disposal is a perfectly feasible method of disposing very large quantities of carbon dioxide”. Since October 1996, at Sleipner Vest in the North Sea, the SACS project has been undenvay, led by Statoil, Norwegian oil and natural gas company, whereby CO2 gas in natural gas is separated and injected into a sandstone formation (Torp 2000). So far, reservoirs under investigation for CO2 disposal have temperatures below 100°C and most are in sedimentary rocks. Solubility of CO2 in water decreases as temperature increases, up to 100°C according to existing data of CO2 solubility in water (Wiebe & Gaddy 1939, 1940, Takenouchi & Kennedy 1964, Scharlin et al. 1995). The data shows that there is a range of temperatures and pressures at which the solubility of CO2 in water is high, even above 100°C. Our simple calculations, using existing data of CO2 solubility in water suggest that hot-

ter and deeper rock masses may prove a suitable candidate for underground CO2 injection, due to enhanced CO2 solubility under high pressures, even at high temperatures. At present, mainly sedimentary formations have been investigated, but magmatic rocks, even with formation temperatures above 100°C may be considered, and provide valid reservoir conditions for CO2 storage and disposal. Consequently, we might expect a range of potential CO2 storage and disposal sites. In subsurface reservoirs there are likely to be many pores and cracks, which may be filled with pore water, in which reservoir rocks and underground water are in (or near) equilibrium. When CO2 is injected into the reservoir, there is an equilibrium shift between the rock, water and C02, whereby dissolution of the host rock, formation of secondary minerals and changes of amounts of dissolved CO2 will occur. It is these changes that affect the ability of the reservoir to capture C02. Consequently, it is essential to understand the interactions between the rock-forming minerals, fluid and CO2 under simulated reservoir conditions, to evaluate the long-term CO2 storage capability of the reservoir. However, there are few investigations that have considered rock/water/C02 interaction processes above 100°C compared to more general rocldwater interaction studies, and it is necessary to promote more work in this important area of research. The aim of this study is to establish a fundamental understanding of rocWwaterlC02 interaction processes under a wide temperature range, from 100°C to

341

near critical temperatures, using granite to represent magmatic rocks and sandstone as representative of sedimentary rocks.

2 EXPERIMENTAL In order to investigate the effect of CO2 on waterlrock interaction processes, some experiments were undertaken on granite samples and sandstone samples in both H20 and H20/C02 systems. Two types of rock were used in this reaction experiment. The first rock-type used was Iidate granite (from Iidate, Fukushima, Japan), which is composed of quartz, plagioclase, K-feldspar and other subordinate minerals. Table 1 gives composition of Iidate granite. The other rock material was Kimachi sandstone (from Shishido, Shimane, Japan), which consists of altered volcanic rock fragments (mainly plagioclase), granitic rock fragments and a fine matrix of clinoptilolite Composition of Kimachi sandstone is given in Table 2. Table 1. Composition of Iidate granite. wt% Modal mineral Oxide Si02 73.99 Quartz A1203 13.40 K-feldspar Fe203 2.05 Plagioclase MgO 0.36 Biotite CaO 1.80 Others Na20 3.58 K20 3.78 Others 0.26 Total 99.22 Total

vol% 37.1 21.8 34.0 6.3 0.6

perimental temperatures were 100, 200, 300 and 35OoC, and experimental pressures were established by saturated vapor pressure of mixed fluid at each temperature condition. After each reaction, solid material, residual solution and gas were separated and analyzed. The reacted rock specimen was dried and weighed, and surfaces textures examined by an SEM equipped with EDX. The residual solution was measured for pH and major cation (Si, Al, Fe, Mg, Ca, Na and K) concentration determined by ICP emission spectrometry.

Figure 1. Schematic of a batch type autoclave used in this study.

3 RESULTS AND DISCUSSION 99.8

In order to investigate the effects of CO2 on the dissolution and deposition behaviors of granite and sandstone, we compared the data between from the rocuwater system and from the rocWwaterlC02 system. The data compared in this study were major element concentration, total major element concentration, weight loss of rock specimen, and surfaces textures of rock specimen. The major element concentration was determined by ICP emission spectrometry analysis of the residual solution. The total major element concentration was expressed as sum of elements’ oxide concentration calculated from the measured concentration of each element. The weight loss of rock specimen was calculated from weight before and after each reaction.

Table 2. Composition of Kimachi sandstone. Modal mineral vol% wt% Oxide Altered clastics 89 6 1.43 Si02 A1263 16.14 (plagioclase main) Fe203 6.83 Granitic rock fragments I1 MgO 2.91 (Plagioclase) (6) CaO 5.98 (K-feldspar) (3) Na20 3.19 (Quartz) (2) K20 1.49 Others (Bi, Mt, Py, ...)* 1 Others 1.09 Total 99.91 Total 101 *Bi, Mt and Py mean biotite, magnetite and pyroxene respectively.

The starting materials comprised a rock sample that was dried and weighed, distilled water (16ml) and dry ice (3.688) as the CO2 supply. Filling ratio of water to the inner volume of the autoclave was 42%. The pre-determined amounts of starting materials were placed in a valve-type batch autoclave as shown in Figure 1. Using an induction heating system, the autoclave was heated up to a predetermined temperature and then kept at the temperature for a week, under rocking conditions. Ex-

342

3.1 Granite Figures 2 and 3 give concentration of major elements, total major element concentration and weight loss of granite specimen after each granite/water reaction and granite/water/C02 reaction respectively. Figure 2 indicates high concentrations of alkali metals and alkaline-earth metals, especially Ca, at lower temperatures ranges, in the presence of COZ.

Concentrations of these elements decreased in the residual fluid, for reaction experiments undertaken at more elevated temperature conditions. Mg and Cabearing minerals seemed to preferentially dissolve at lOO"C, whilst Na and K were added to the reaction solution at 200°C. In contrast, the residual solution from the waterKO2 system experiments was deficient in Al. SEM observation showed that different types of deposit andor secondary minerals were formed in the waterKO2 systems compared with the water systems above 200°C. They were especially observed on surfaces of dissolved plagioclase. It seemed that the presence of CO2 much produced a variety of deposits, andor secondary minerals on surface of the granite. An infinite number of small, flake-like crystallites of an unidentified aluminosilicate, occur on the surface of dissolved plagioclase.

dition due to addition of C02. The solution in acid condition increases dissolution rate of the granite. A rapid progress of dissolution of the granite leads the solution to be supersaturated. Supersaturated solution results in deposition of secondary minerals. It is shown in Figure 3a that the total major element concentration in the residual fluid for the water/C02 system is higher than that for the water system. This indicates that the total dissolution of the granite is enhanced due to the presence of C02. In contrast, Figure 3b shows that the sample weight loss for the water/C02 system is smaller than that for the water system, in spite of the enhanced dissolution in the waterKO2 system. The trend of Figures 3a and 3b and mass balance consideration imply that some CO2 in the solution may be fixed on to the rock sample as the secondary minerals deposited in the waterKO2 system, possibly as carbonate minerals. Indeed, our numerical computations of phase equilibrium calculation have suggested that some carbonate minerals were supersaturated in the water/C02 system.

Figure 3. Total major element concentration (a). Weight loss of granite specimen (b).

3.2 Sandstone Figsures 4 and 5 show concentration of major elements, total major element concentration, and weight loss of sandstone specimen, after each sandstone/water reaction and sandstone/water/C02 reaction respectively. Analysis of residual solutions (Fig. 4) indicated very high Na concentration in both systems, compared with other elements. According to SEM observation, Na in the solution for the water system probably comes from the dissolution of clinoptilolite. For the waterKO2 system, Na may mainly derive from clinoptilolite and plagioclase in rock fragment. The addition of CO2 to the water system significantly increased the rate of dissolution of alkali metals and alkaline-earth metals, especially Na and Ca, from the primary minerals. The concentration of these elements in the residual water decreased as reactiodexperimental temperature increased. Increased Na and Ca concentrations in the

Figure 2. Concentration of major element in the residual solution after each reaction on granite.

Based on the analysis of residual solution, and visual observation of the granite, it is suggested that dissolution of granite and deposition of secondary minerals is enhanced by the presence of excess C02. This behavior may be occurred as follows. In initial of reaction, solution of water/CO2 system is in acid con343

solution due to the presence of C02, are considered to derive from the alteration of plagioclase andor other minerals containing Ca and Na included in rock fragments, in the waterKO2 system. In contrast, the concentration of A1 was repressed in the waterKO2 system. Compared with the granite case, trend of temperature dependency of each element concentration for the sandstone was similar to that for the granite. For the sandstone, it was difficult to decide secondary minerals through SEM observation because of its altered surface before reaction. Main secondary minerals in both systems were some flake-like deposits on rock fragments and lining deposit on boundary between rock fragment and matrix. The presence of CO2 seems to facilitate the deposition of various secondary minerals and promote the flakelike deposition on the surface of altered plagioclase included in rock fragments, including an unidentified aluminosilicate. In contrast, the lining deposit formed at 300 and 350°C, much developed in the water system than in the water/C02 system.

fluid (Fig. 5a), and the determination of sample weight change (Fig. 5b) imply that the secondary minerals deposited in the waterKO2 system have the potential to take-up some C02.

Figure 5. Total major element concentration (a). Weight loss of sandstone specimens (b).

4 CONCLUSIONS

Dissolution and deposition experiment was conducted to better understand rocWwaterlC02 interaction processes on granite and sandstone. The result showed that, in the waterKO2 system, alteration of plagioclase andor other minerals containing Ca and Na was preferentially accelerated. It was suggested that secondary mineral phases have the potential of up-taking dissolved C02. ACKNOWLEDGEMENT A part of this work was supported by “Research for the Future” Program (JSPS-RFTF 97P0090 l), The Japan Society for the Promotion of Science. REFERENCES Holloway, S. (ed.) 1996. The Underground Disposal of Carbon Dioxide, Final Report of Joule I1 Project No. CT92-003. Scharlin, P., Young, C.L., Clever, H.L. & R. Crovetto 1995. Solubility of Carbon Dioxide in Pure Water. In: P. Scharlin (ed.), Solubility Data Series, Vol.62 “Carbon Dioxide in Water and Aqueous Solutions ”: 1-66. IUPAC, Oxford University Press. Takenouchi, S. & G. Kennedy 1964. The binary system H20CO2 at high temperatures and pressures. Am. J. Sci. 262: 1055-1074. T o p , T.A 2000. “SACS”-Saline Aquifer CO2 Storage - Final Technical Report. Wiebe, R. & V.L. Gaddy 1939. The solubility in water Of carbon dioxide at 50,75 and 1OOOC, at pressures to 700 atmospheres. J. Am. Chem. Soc.6 1: 3 15-3 18. Wiebe, R. & V.L. Gaddy 1940. The solubility of carbon dioxide in water at various temperatures from 12 to 40°C and at pressures to 500 atmospheres. Critical phenomena. J. Am. Chem. SOC.62: 815-8 17.

Figure 4. Concentration of major element in the residual solution after each reaction on the sandstone.

As in the case of the granite, the assessment of the total major element concentration in the residual 344

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

The source of sodium in groundwater, Pannonian B asin, Hungary I. Varsinyi Department of Mineralogy, Geochemistry and Petrology, Universiry of Szeged, Hungary

L.O .KOVBCS Hungarian Geological Survey, Budapest, Hungary

ABSTRACT: One of the main sources of sodium in the subsurface waters of the study region is feldspar weathering. Standard formation Gibbs energy and standard enthalpy of albite weathering into montmodonite were calculated based on the method developed by Tardy & Garrels (1974). Ion activity product of the reaction was also calculated in the subsurfke waters of the 100-2500 m depth interval. Temperature dependence of the equilibrium constant and that of the IAP values indicate that the driving force of the reaction is effective up to 70-80 "C.

1 INTRODUCTION

In the central part of the Pannonian Basin there are two main types of groundwater: Ca and Mg bicarbonate, and Na bicarbonate types. Ca and Mg bicarbonate waters are characteristic of the coarser-grained Pleistocene sediments in the recharge area. Na bicarbonate water type occurs in the Pleistocene discharge areas of her-grained materials, in the Pliocene and Pontian aquifers, and at the Pannonian/Pontian boundary (Upper Miocene). The amount of dissolved material is low in the water flow system within the Pleistocene layers in the western part of the study area. All Ca and Mg bicarbonate type water samples and several Na bicarbonate water samples belong to this flow system. In these waters the concentrations of Ca, Mg and Na are controlled by ion exchange (Varshyi & 0.Kovks 1997). Na bicarbonate type waters are characteristic of the Pleistocene layers located in the eastern part of the study area where water flows upward from the Pliocene layers, and of the Pliocene and Pontian layers. The dissolved solid content is high in these waters. The aim of the present work is to identltjr processes controlling the sodium concentration in the Na bicarbonate type waters in the 100 to 2500 m depth interval.

the study area since the Miocene. 1000-3000 m thick marine sediments are overlain by 1000-3000 m of lake sediments. In the upper part of the Pliocene, lacustrine and fluvial sediments are interbedded. At about the end of Pliocene the basin was uplifted. Fluvial sedimentation started in the inner part of the basin about 2.4 million years ago, at about the beginning of the Pleistocene. Local scale tectonic events together with climatic factors led to a cyclic fluvial sedimentation during the Pleistocene ( R h a i 1985). The thickness and dip of the studied layers is portrayed in a WSW-ENE cross-section (Fig. 2).

Figure 1. Location of the study area. 2 DESCRIPTION OF THE STUDY AREA The study area is located in the south-eastern part of Hungary (Fig. 1). It represents the central part of the Pannonian Basin filled up with Neogene sediments. Lithological and paleo-geomorphologicalconsiderations support that continuous sedimentation occurred over 345

In the fluvial sediments the most common minerals are quartz, feldspar, calcite, dolomite, illite-smectite mixed layers, muscovite and chlorite (Viczh 1982, Varsinyi & 0.Kovhcs 1994). The mineralogical composition of the deeper lacustrine sediments is similar to that of the

components within the silicate structures difEers fiom the standard Gibbs energy assigned to the same components as separate phases. AGP values of the oxide and hydroxide components in the silicate structure were used to calculate AGP value for montmodonite. In the present work, the Naos(AL5Mg05) (S&O1o)(OH)2montmorillonite formula written in form of oxides and hydroxides as 0.25Na20 0.75Al203 * 0.5Mg(OH)2 4sio2 0.5H~0, allows us to estimate both AGP and AHp values. The standard Gibbs energy and standard enthalpy of formation for different silicate minerals of the simplest structure (chrysotile, talc, sepiolite, kaolinite, pyrophyllite, paragonite) were calculated fiom the standard Gibbs energy and standard enthalpy of their dissolution reactions (AG? and AH?). Data in thermodynamic data bases available for the aqueous species at 298.15 OK (Wagman et al. 1982, Nordstrom et al. 1990, Bloom & Weaver 1982), and equations of dissolution reactions, AG? and AH: values for chrysotile, talc, sepiolite, kaolinite and Al(0H)i given in WATEQF (Plummer et al. 1984), for pyrophyllite and paragonite given in SOLMINEQ.88 (Kharaka et al. 1988) were used to calculate AGfo and A€€;of these minerals: 0

Figure 2. WSW-ENE cross-section. Q=Quarternary, Pl=Pliocene, M3Po=Pontian stage of Upper Miocene, M3Pa=Pannonian stage of Upper Miocene, MZ=Middle Miocene, dashed arrow=Pleistocene flow system. fluvial layers; they consist of quartz, mica, montmorillonite (mixed ate-smectite layers), calcite, dolomite, Na- and K-feldspars, chlorite, and a very small amount of kaolinite (Varsimyi 1975, Vie* 1982).

3 RESULTS AND DISCUSSION In the study area both albite and montmorillonite are available in the sediment. It has been supposed that albite weathering into montmorillonite is the source of sodium except in waters of the Pleistocene flow system where ion exchange processes occur (Fig. 2). The albite +montmorillonite reaction is the following: (albite) 3NaAlShOs + Mg2'+4H20 -+

chrysotile Mg3Si205(0H)4+ 5H20 = 3Mg2" + 2&Si04 + 60HAGp = 295.3 kJmol-' AHp = 115.41 kJmol-' talc = 3Mg2" M~S&Ol~OH)2+loH,o+10H20 AG; = 355.1 1 kJmol-' AHp = 188.55 kJmo1-'

+ 4&Si04+ 60H-

--+ ~ N ~ o . ~ A ~ ~ . ~ M+~2Na' ~.~ + H4SiO4 S ~ ~ O ~ Osepiolite (OH)~ (montmorillonite)

Standard Gibbs fiee energy (AG?), standard enthalpy

(AH?),and equiliirium constant (K) of the above reaction were calculated at 298.15 OK. In a smectite-water system the composition of the solution is controlled by the dissolutiodprecipitation reactions of gibbsite or amorphous aluminium hydroxide rather than by those of smectite. That is the reason why fiee energy of formation for smectite could not be experimentally determined. Several empirical methods were developed for estimating the standard fiee energy of formation (AGP ) for layer silicates (Tardy & Garrels 1974, Mattigod & Sposito 1978). These methods are based on the solubility measurements of the silicate minerals with the simplest structure which attain equilibrium in a mineral-water system. In this work the method developed by Tardy & Garrels (1974) was used. This method is based on the assumption that layer silicates can be represented by oxide and hydroxide components. The standard Gibbs energy of formation of these oxide and hydroxide

346

Mg&306(0H)4+6H20 = 2Mg2" AGp = 228.6 kJmol-' A H ~ =111.01 kJmol-'

+ 3&sio4+ 40H-

kaolinite A12Si20AOH)4+7H20 = 2A1(OH)4 + 2&sio4+ 2H+ AG? = 2 10.42 kJmo1-' AHp = 205.64 kJmol-' AI% + 4 0 B = AI(0H)i AG; = -188.13 kJmol" AHp = -46.7 kJmol-' AGfoN(* = -1305.73 Wmol-' AHFHOw = -1 507.64 kJmol-' pyrophy llite A12S&Ol~OH~+6H" +4H20=2A1%+ 4&sio4 10gKm = -0.10 10gKm = 1.42 AGp = 0.57 kJmol-' A H =~-94.5 Wrnol-'

transformation is -10.86 kJ/mol (log K=1.90), and AHP is -169.1 kJ/mol. Equilibrium constant and ion activity product (IAP), given as

paragonite = Na' + 3A13' + 3KSi04 NaA13Si3010(OH~+10H' 1 0 6 2 9 8 = 14.39 l0gK2n = 18.42 AGP = -82.14 Wmol-' AJ4P = -250.81 kJmol-' The calculated AGfo and AHfo values of chrysotile, talc, sepiolite, kaolinite, pyrophyllite and paragonite are summarized in Table 1, and those of oxides and hydroxides in Table 2.

log IAP = 2-log ["I

+ log [HbSiO4]- [log Mg '+

1,

resulting fiom albite weathering into montmorillonite, were plotted versus bottom hole temperature (Fig. 3).

Table 1. AGfoand AHfo values of minerals. Mineral

~ ~ (kJmol-') f o ~ ~ (kJmo1-I) f o

chrysotile

-4037.0

-428 1.92

talc

-5529.11

-5940.51

sepiolite

-4273.36

-462 1.97

kaolinite

-3779.90

-4134.71

pyrophyllite

-5265.61

-5673.18

paragonite

-5573.36

-5984.13

Table 2. AGfoand AHfo values of oxides and hydroxides. component

~ ~ f o (k.Jmol-') sil ~ ~ f o (Hmol') sil

wow2

-848.3

-874.44

MgO

-622.56

-617.65

SO:!

-858.92

-957.69

A1203

-1597.8

- 1465.51

H20

-232.13

-376.91

NazO

-735.54

-1071.77

Figure 3. Temperature dependence of IAP and K on albite weathering into montmorillonite. With increasing temperature the driving force of the reaction is decreasing.

Based on the AGfo and AHfo values for the minerals, AGP and AHfo for the oxide and hydroxide components in the layer silicates structure were calculated by solving sets of simultaneous equations. The fiee energy and enthalpy of montmorillonite formation fiom the AGfo components and A€€;of the oxide and hydroxide are -5358.16 kJ/mol and -5823.37 kJ/mol, respectively. At 298.15 "K the AGP and AHP for the albite -+ montmorillonite transformation were calculated fkom the standard Gibbs energy and standard enthalpy of the albite formation determined by Wagman et al. (1982), and fiom the standard Gabs energy and standard enthalpy of montmorillinite formation calculated above. The calculated AG; for the albite -+ montmorillonite 347

Temperature of the waters belonging to the Pleistocene flow system is about 20-30 "C. In these waters, IAP is independent of the temperature. The increase in LAP is caused by the increase in the sodium concentration due to ion exchange. In the other samples, logIAP and logK of the albite weathering into montmorillonite reaction show a strong temperature dependence. Below 70 "C the very high undersaturation means a strong driving force for the reaction. With increasing temperature, undersaturation and the driving force are decreasing, and between 70 and 80 "C the albitemontmorillonite-water system approaches the equilibrium. Above 80 "C there is no more driving force for the albite weathering into montmorillonite, and the Na concentration remains constant. In the water samples fiom the Pontian layers, however, the Na concentration increases above 80 "C. These samples originate fiom the PaPo boundary. In these PaPo waters the concentrations of Sr and Ba are increasing together with that of Na (Fig. 4), while in the waters fiom the overlying layers there is no correlation

Figure 4. Barium vs. sodium. When albite weathering controls the Na concentration there is no correlation between Na and Ba. In the PdPo samples, the concentration of Na is increasing with that of Ba, indicating that the source of Na is partly different from that in the overlying layers. between them. These relationships suggest that at the PaPo boundary Na originates from albite weathering and from other, unidentified weathering or dissolution processes, as well. 4 CONCLUSIONS There are three main sources of sodium in the subsurface waters in the studied region of the Pannonian Basin. In the Pleistocene flow system the main source is ion exchange. In the Pleistocene layers discharging the Pliocene, and in the Pliocene and Pontian up to 80 "C, albite weathering into montmorillonite provides Na. Above 80 "C, dissolution reactions, other than albite -j montmorillonite, occur producing not only Na but Sr and Ba, as well. ACKNOWLEDGEMENTS The work was financed by the Hungarian Research Fund (OTKA). Project number is T 02624 1.

REFERENCES Bloom, P.R. & R.W. Weaver 1982. Effect of the removal of reactive surface material on the solubility of synthetic gibbsites. Clays Clay Miner. 30: 28 1-286.

348

Kharaka, J.K., Gunter, W.D., Aggarwal, P.K., Perkins, E.H. & J.D. Debraal 1988. SOLMINEQ.88, a computer program for geochemical modeling of water-rock interactions. USGS Water Resources InvestigationsReport 88-4227. Mattigod, S.P. & G.I. Sposito. 1978. Improved method for estimating the standard free energies for formation (AGP 298.15) of smectites. Geochim. Cosmochim. Acta, 42: 17531762. Nordstrom, D.K., Plummer, L.N., Langmuir, D. et al. 1990. Revised chemical equilibrium data for major water-mineral reactions and their limitations. In: Melchior, D.D. & Bassett, R.L. (eds): Chemical modeling in aqueous systems I t 857892. American Chemistry Society (Symposium Series 416). Plummer, L.N., Jones, B.F. & A.H. Truesdell 1984. WATEQF a Fortran IV version of WATEQ, a computer program for calculating chemical equilibrium of natural waters. USGS Water Resources Investigations Report 76- 13. Ronai, A. 1985. The Quarternary of the Great Hungarian Plain. Geologica Hungarica, 2 1. Institutum Geologicum Hungaricum, Budapest. Tardy, Y. & R.M. Garrels 1974. A method of estimating the Gibbs energies of formation of layer silicates. Geochem. Cosmochem. Acta, 38: 1101-1 116. Varsknyi, I. 1975. Clay minerals of the Southern Great Hungarian Plain. Acta Miner, Petr. Szeged, XXIVl: 5 1-60. Varsknyi, I. & L.O.KOV~CS 1994. Combination of statistical methods with modelling mineral-water interaction: a study of groundwater in the Great Hungarian Plain. Applied Geochemistry, 9: 419-430. Varsanyi, I. & L.O.KOV~CS 1997. Chemical evolution of groundwater in the River Danube deposits in the southern part of the Pannonian Basin (Hungary). Applied Geochemistry, 12: 625-637. Viczian, I. 1982. An expanding mixed-layer clay mineral in Upper Pannonian to Pleistocene fine-grained clastic rocks of the borehole Pusztaottlaka IIP (SE Hungary). Ann. Rep. Geol. Ins. Hun. 1980: 449-457. Wagman, D.D., Evans, W.H. & V.B. Parker 1982. The NBS tables of chemical thermodynamic properties: selected values for inorganic and C1 and C2 organic substances in SI units. J. Phys. Chem. ReJ: Data,ll (Suppl. 2): 1-392.

Water-RockInteraction 2001,Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Silica solubility geothermometers for hydrothermal systems M .P.Verma Geotermia,Instituto de Investigaciones Electricas,Apdo. 1-475,Cuernavaca 62001, Mor., Mexico

ABSTRACT: From the existing solubility data, the regression equations for quartz and amorphous silica along the liquid-vapor saturation for pure water are derived for the temperature range 0-374°C. Amorphous silica and quartz are the extreme silica phases under the conditions of hydrothermal systems and their solubilities are widely different. Therefore, the application of silica geothermometers provides a wide difference in the estimated temperatures for a specific silica concentration. Another limitation is the correction of total discharge silica concentration from a well for the vapor fraction in the geothermal reservoir and the definition of the silica phase in equilibrium with reservoir fluid. An iteration process is used to calculate the deep reservoir temperature and vapor fraction, but a basic assumption is that the reservoir fluid was in equilibrium with either amorphous silica or quartz and existence of water-vapor saturation condition. In both cases, the error in the estimated temperatures is close to k3O"C.

1 INTRODUCTION White et al. (1956) initiated the development of silica geothermometry through comparing the field evidences on the silica content in geothermal fluids with the experimental temperature dependent silica solubility data. Recently, many workers (Rimstidt 1997, Gunnarsson & Arnorsson 2000, Verma 2000a) compiled the silica solubility data in order to obtain the regression equations along the water-vapor saturation. Silica is found in many stable phases in natural and engineered earth systems including quartz, chalcedony, tridymite, moganite, cristobalite, coesite, stishovite, lechatelierite (silica glass), opal and amorphous silica. The dissolution-precipitation equilibration of such multi-phase minerals depends upon the solution-mineral contact time, and it requires an understanding of mineral solubility kinetics. Quartz is the most stable phase and has the lowest solubility, whereas amorphous silica is the least stable phase and has the highest solubility. Thus quartz and amorphous silica must represent two extreme cases of silica dissolution-precipitation equilibria in hydrothermal systems. The solubility of others silica phases will be in between the two extreme solubilities. It has been emphasized that the resident time for geothermal reservoir fluid is high enough to reach in equilibrium with quartz (Fournier & Rowe 1966). Recently, in case of CP-M-19A well, Verma ( 1997) observed that the calculated concentration of

silica in the deep reservoir fluid was substantially higher than the quartz solubility at the reservoir temperature. In case of natural manifestations there are many more limitations like dilution with cold water, loss or gain of steam, re-equilibration, etc., which will not be discussed here. In this article all the quartz solubility data along the liquid-vapor saturation curve and compressed liquid for pure water will be analyzed. The causes of decrease in the quartz solubility values in the existing data along the liquid-vapor saturation above 300°C will be discussed. On this basis, the regression equations for quartz and amorphous silica will be derived for whole range of liquid-vapor saturation temperature (0-374°C). Similarly, the limitations to use these equations as geothermometers for geothermal fluids will be presented. 2 EXPERIMENTAL SILICA SOLUBILITY All the quartz solubility data from literature (Rimstidt 1997, Verma 2000a and references cited therein) are divided in two groups (Fig. 1): a) along the liquid-vapor saturation curve and b) in the compressed liquid region. The quartz solubility at room temperature (25°C and 1 bar) is reported in the range 6- 16ppm. But Rimstidt (1 997) conducted the solubility determination experiments for a long period and got the value of ll.Ok1.1 ppm. Attainment of equilibrium between water and quartz at room

349

Figure 1. Experimental quartz solubility data along the liquidvapor saturation curve and in the compressed liquid region.

Figure 2. A regression relation for quartz and amorphous solubility data along liquid-vapor saturation curve (modified after Verma, 2000b)

temperature requires doing experiments for geological time period without supersaturating the solution at any instant, to prevent equilibrium with other silica phases. Rimstidt (1 997) critically analyzed the solubility data and fitted the regression expression up to 300°C. He took repeated datasets by same author from literature. Therefore, the solubility data are refitted here removing the duplicated values (Fig. 2). The refitted expression is more or less the same proposed by Rimstidt; but it is more realistic, because a biased statistical evaluation of the dataset due to repetition of data points is reduced. The errors in the coefficients of the following regression equation are only the statistical errors f l CT (standard deviation). The regression expression for quartz is: log SiO, (ppm) = -

1175.7(f31.7) T(K)

low the critical point and as temperature gets lower the difference in the specific volumes increases. At 25°C the difference is very high; therefore any amount of vapor fraction does not affect the silica solubility. But considerable mass of water can be present even in the small fraction vapor in experiments performed above 300°C when this mixture of vapor and liquid is quenched; the vapor condenses and dilutes the liquid water. That may be a reason of decrease in quartz solubility above 300°C in the Figure 1. Verma (2001) presented that some problems might also be associated with the experimental design for extracting a small amount of solution from the reaction solution while the vessel is maintained at the specified temperature and pressure. For example, the evaporation of the solution due extraction in order to fill the empty part of the vessel and trapping of some condensed vapor in the extraction pipeline are important reasons, which may affect the concentration of dissolved silica. The experimental details were not reported in the literature in order to perform such corrections. Therefore, there is need to repeat the silica solubility data above 300°C along the liquid-vapor saturation curve. The values of pressure and temperature for all the solubility experiments in the compressed liquid region together with the theoretical PT curves for the two extreme cases (i.e. when the vessel (bomb) is completely filled with water and when there is just enough water to make the total specific volume equal to the critical volume of water at 25°C) are plotted in Figure 3 (Verma 2000a). All the experimental pressure and temperature data lie between the theoretical curves. Therefore the pressure and temperature during all these solubility studies were

+ 4.88(4 0.08) (1)

Similarly the regression expression for amorphous silica is 724.68(f 8 1 .O) logSiOz(ppm)= + 4.50(k 0.13) (2) T(K) The silica solubility determination has been performed using one of three methods: (a) weight loss of quartz in a known amount of water, (b) chemical analysis of dissolved silica remaining in solution after rapid quenching and opening of the reaction vessel and (c) chemical analysis of dissolved silica in a small amount of solution extracted from the reaction solution while the vessel is maintained at the specified temperature and pressure. Verma (2000a) demonstrated that the early measurements of silica solubility were carried out mostly using the first two methods and were affected by the fraction of vapor present in the reaction vessel. The specific volume of water vapor and liquid are close to the same just be350

Figure 3. The pressure and temperature values for all the experimental determinations of quartz solubility data in the compressed liquid region together with theoretical curves for the two extreme cases for existence of liquid water in the reaction vessel. The liquid-vapor saturation curve is extended with dashed curve after the critical point (see Verma, 2000a).

probably controlled by the different amount of water in the reaction vessel. Similarly, it can be observed that all the data points even in the supercritical region (e.g. the region for temperature and pressure above the critical point) fall in the compressed liquid region. There is not even a single point in the vapor region. A question comes in mind why there is no experimental quartz solubility data in the superheated vapor region, when the solubility data have been measured in vapor phase along the saturation curve (Fournier & Potter 1982 and others). Similarly, if quartz is soluble in vapor phase, it should be in molecular form. The atmosphere is a mixture of oxygen and nitrogen, both are in molecular form. Gases can be mixed in any proportion. Why does quartz has specific solubility values in vapor? There are many geothermal systems around the world like Lederrello, Geysers and Los Humeros, which produce dry steam at high temperatures (higher than 300°C). The condensed vapor does not have dissolved silica. Even we do not find more volatile species like Na, C1, etc in the vapor phase of these geothermal systems and no one has measured the solubility of these species in vapor phase. Therefore, the silica solubility data in vapor phase are incorrect. Secondly, if it is true that there is decrease in the quartz solubility after certain maximum along the saturation curve. There will be two temperature values for a specific value of silica content in water. For example, there will be approximately temperature 280 and 370°C for 600 ppm of silica and 260 and 372°C for 500 ppm (see curve from Fournier & Potter (1982) in Figure 1). Thus there could be higher temperature even for low concentration of

Figure 4. The calculated concentration of silica in the liquid phase in the reservoir and in the total discharge concentration (after Verma, 2000b).

silica. Similarly, it is difficult to predict the right temperature for a given silica concentration in the reservoir fluid. Under these circumstances a combined evaluation of quartz solubility data along the saturation curve and in the compressed liquid region will be helpful in understanding the behavior of quartz solubility data after 3OO0C (Fig. 1). Ragnarsdottir & Walther (1983) presented a pressure dependence study on the quartz solubility. Rimstidt (1 997) used the result for the study to adjust the solubility (log m) values along the saturation curve at 300°C and 8.581 MPa to the values at 300°C and 0.1 MPa. This is done by subtracting only 0.02 units. It means that the effect of pressure from 8.581 MPa to 0.1 MPa on the quartz solubility is quite less. Therefore it is more reliable to extrapolate the linear tendency for the data up to 300°C in the absence of the correct silica solubility data along the liquid-vapor saturation curve above 300°C. 3 CRITIQUE ON SILICA GEOTHERMOMETRY

The silica geothermometers have been applied extensively to estimate geothermal reservoir temperature from the silica concentration of the fluid obtained from natural manifestations and drilled wells. Unfortunately, the predicted temperatures generally show a wide dispersion even when applying a single geothermometer to all the wells in a geothermal field. Many reasons have been proposed to justifL the discrepancies, including gain or loss of steam phase in the reservoir, mixing of different types of fluids, re-equilibration during ascension to the sur-

351

face, precipitation-dissolution, etc. Enormous works have been done on improving the geothermometer equations and their applications. But there is a fundamental question to be answered on the correction of silica total-discharge concentration for vapor fraction in the geothermal reservoir fluid in order to use silica content in geothermal fluids as a chemical geothermometer. Similarly, it is also needed to justify that the silica in the quartz phase is in the equilibrium with the geothermal fluid. Verma (1997) presented a two-phase flow approach to calculate the fluid thermodynamic parameters including chemical speciation, pressure, and temperature in a geothermal reservoir from the parameters measured in the geothermal fluid (vapor and liquid) at the wellhead separator. Using this approach the geothermal reservoir fluid parameters were calculated in the well M-19A at Cerro Prieto. The concentration of silica is 666 ppm and the reservoir temperature is 248°C. The data point is shown in Figure 4. It can be observed that the value is higher than the experimental quartz solubility, but is lower than the amorphous silica solubility. Thus the fluid is supersaturated with respect to quartz, or the solubility of silica is controlled by another phase or there is lost of vapor in the reservoir.

REFERENCES Fournier, R.O. & R.W. I1 Potter 1982. A revised and expanded silica (quartz) geothermometer. Geotherm. Resourc. Counc. Bull., 11: 3-12. Fournier, R.O. & J.J. Rowe 1966. Estimation of underground temperatures from silica content of water from hot springs and wet-steam wells. Amer. J. Sci., 264, 685-697. Gunnarsson, I. & S. Amorsson 2000. Amorphous silica solubility and the thermodynamic properties of H4Si02 in the rang 0’ to 350’C at P,,,. Geochim. Cosrnochim. Acta, 64, 2295-2307. Ragnarsdottir K.V. & J.W. Walther 1983. Pressure sensitive “silica geothermometer” determined from quartz solubility experiments at 250°C. Geochim. Cosmochim. Acta, 47: 941-946. Rimstidt, J.D. 1997. Quartz solubility at low temperatures. Geochim. Cosmochim. Acta, 61: 2553-2558. Verma, M.P. 1997. Thermodynamic classification of vapor and liquid dominated reservoir and fluid geochemical parameter calculations. Geofisica hternacional, 36: 181-189. Verma, M.P. 2000a. Chemical thermodynamics of silica: a critique on its geothermometer. Geothermics, 29: 323-246. Verma, M.P. 2000b. Limitations in applying silica geothermometers for geothermal reservoir evaluation. Verma, M.P. 200 1. Silica (quartz) solubility regression equation along the water vapor saturation curve: a chemical geothermometer for hydrothermal systems. J. Volcan. Geotherm. Res. In revision. White, D.E., Brannock, W.W. & KJ. Murata 1956. Silica in hot-spring waters. Geochim. Cosmochim. Acta, 10: 27-39.

4 CONCLUSIONS New quartz and amorphous silica solubility regression expressions are derived by a critical scrutiny of all the existing experimental data along the watervapor saturation and in the compressed liquid region. The unknown value of reservoir vapor fraction is a fundamental limitation in using the quartz solubility expression as a chemical geothermometer on the total discharge composition of a well. In case of the CP-M-19A well at Cerro Prieto it has been calculated that the reservoir temperature is 248°C and a vapor fraction 0.224 by weight with a two-flow approach. There could be a wide range of silica concentrations in geothermal fluids depending on the silica phase in equilibrium at a specified temperature. Additionally, the uncertainty in equilibrium temperatures is approximately +30”C. The reservoir fluid is supersaturated with respect to quartz, but sub-saturated with respect to amorphous silica according to the two-phase flow calculations. ACKNOWLEDGEMENTS The author appreciates the constructive comments and suggestions from Dr. Luigi Marini and an anonymous reviewer, which contributed considerably in improving the manuscript.

352

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Leaching kmetics of a quartz-chlorite schist and consequent changes in the rock structure T.Wells, P.Binning, G.Willgoose & A.Mews Department of Civil, Surveying and Environmental Engineering, University of Newcastle, Australia

ABSTRACT: Recent research has shown that the rate limiting process in erosion is often the production of fine material from parent rock. In this study, the weathering of a quartz-chlorite schist from the Ranger Uranium mine in Northern Australia has been examined. The leachate composition, rock surface area, as well as pore and particle size distributions were examined over time. Leachate compositions were found to follow a power law trend after an initial surge. The initial particle size of the rock did not affect leachate composition because particle surface areas per unit weight were similar for all particle sizes. Tests on the rock matrix after 14 days leaching showed up to a 20% drop in surface area and a change in the pore structure.

1 INTRODUCTION Significant advances have recently been made in the long term modelling of the erosion of man-made landforms such as the waste rock dump generated by the Ranger Uranium mine (Northern Territory, Australia; Hancock et al. 2000, Willgoose & Riley 1998). Such modelling is important in assessing the environmental impact of mines well after they have ceased operation. In the longer term, (1000s of years), it is the breakdown of larger rock specimens to finer, more transportable material that becomes the rate limiting step in the erosion process. Thus to extend the predictive capabilities of landform evolution models it is important that the breakdown processes specific to that region are understood and can be quantified. As a first step in addressing this lack of knowledge a study has been conducted on the dissolution kinetics of a quartz-chlorite schist, a rock type common to the Ranger Uranium mine waste rock site. In different experiments, several size fractions of ground rock were contacted with de-ionized water and a simulated monsoonal rainwater. The concentration of several cations in the leachate solution was determined periodically. This approach is similar to that taken by many previous studies that have examined the time dependent composition of the leachates derived from numerous rock and mineral types, (e.g. Busenberg & Clemency (1976), Wollast (1967) and Pickering (1962)), including quartzchlorite schists, (Murakami et al. (1996) and Herbillon & Murakami (1975)). In order to quantify the rock breakdown process

this study also examined particle size distributions and the internal rock structure after a set leaching period. This focus on the changes in rock structure is less common in the literature, one example being the study of Sweevers et al. (1998). 2 EXPERIMENTAL PROCEDURE All powdered rock material was generated from a single sample of schist retrieved from the Ranger Uranium mine waste rock dump. All external surfaces were cut from the sample to ensure no extraneous impurities were present and the remaining rock was then ground in a small hammer mill. Samples of the powder were taken, ground to a talc consistency and the chemical composition determined via x-ray fluorescence. X-ray diffraction was employed to provide a semi-quantitative mineralogical assessment. The results of both examinations are listed in Table 1. In order to minimise sampling error the remaining crushed material was then dry sieved into the following size fractions: ( 60%) occurs in the upper l m (Fig 2), with smaller losses at greater depths. The pattern of chloride temporal variation in the 4 m samples I'lags'' slightly behind the 0.5 m samples (Fig 1). Visual examination of these data, as well as statistical analysis by cross-correlation, suggests that the deeper lysimeter "lags" by about 1 to 2 months, suggesting that this is the residence time of water in the vadose zone. This is consistent with residence times of 13 to 300 days calculated from the hydraulic conductivity and porosity. Thus, conductivity calculations and temporal variations of C1 and other species in the vadose zone are consistent with residence times from ca. 30 to 60 days. Concentrations of major species and pH for one representative sampling date (January 7, 1997) are shown in Figure 2. Precipitation (Pptn) and throughfall (Tf) samples are plotted schematically above 0, the land surface. Samples from 0.25 to 4 m depth are from the vadose zone. Groundwater (Gw) and a sample from an effluent stream (Strm) are plotted at 4.5 and 5 m. Concentrations in groundwater and the effluent stream are generally similar to those in the deep vadose zone. Variation of pH in samples is shown in Figure 2. The precipitation pH is 4.3. The throughfall pH is

higher. Organic constituents, cyclic salts, and mineral grains deposited on plant surfaces tend to neutralize acid rain. The pH in the near surface (0.25 m lysimeter) sample is 5.0; values of pH increase with depth to near 6.0. Concentrations of most dissolved species appear to increase with depth within the vadose zone (Fig 2). However, after compensating for evapotranspiration, Na, Mg, and SO4 are approximately constant whereas K, Ca, and Si02 do increase with depth. Na is almost constant because vadose zone minerals contain relatively little Na. NO3 and SO4 are greater in precipitation and throughfall than in the vadose zone. NO3 decreases at the top of the vadose zone due to biological uptake and denitrification. The evolution of the vadose waters in the K20A1203-Fel03-Si02-H20 system is shown in Figure 3. In this plot the size of the symbol corresponds with depth; small symbols are from near the surface, larger samples are from deeper in the vadose zone. Shallow vadose waters are in the kaolinite field, and move toward or into the smectite and glauconite stability fields at greater depths. There is no obvious temporal variation in solution path. 4 REACTION PATH AND REACTION RATE MODELING

364

NETPATH (Plummer et al. 1991) was used to determine amounts of mineral dissolution/ precipitation consistent with changes in composition of water samples (Fig 2) and phase stability diagrams (Fig 3) as a function of depth within the vadose zone. Results for one representative sampling period, January 23, 1997, are presented in

Figure 3. Phase relations in the K20-Al203-Si02H20 system. Symbol size indicates depth within the vadose zone. Small symbols are near the surface, large symbols indicate greater depths. Figure 4. The results are consistent with the data, but not unique. Previous studies of glauconite weathering (e.g. Wolff 1967, Courbe et al. 1981) have shown that glauconite may weather to form goethite, kaolinite, and mixed-layer or smectite clays. All of these are present in the HTF vadose zone. Courbe et al. (1981) have shown that weathering products are extensively depleted in potassium, often retaining less than 25% of the initial amount, with groundwaters being enriched in dissolved K and silica (Wolfe 1967). The secondary clays could not be completely characterized in this study. Following Courbe et al. (1981), we assume that most (75%) potassium has been removed during glauconite weathering. Precipitation/dissolution amounts for four minerals, glauconite, smectite, kaolinite, and silica are plotted as a function of depth (Fig 4). Reaction rates have been normalized to constant (1 meter) layer thickness. Positive numbers indicate dissolution and negative numbers precipitation. The greatest amount of glauconite reaction occurs, not at the top of the vadose zone, but between about 0.5 to 1.5 m below the surface. Sverdrup & Warfvinge (1995) have previously found that maximum rates of weathering at Gardsjon, Sweden occurred somewhat below the surface. The slightly higher porosity and number of macropores at the top of the vadose zone may result in faster water movement and less time for mineral reactions. Further, most rapidly reacting and finest grained minerals may have been removed from the upper vadose zone. The glauconite reaction rate decreases at depth and is essentially zero at the water table. Kaolinite precipitates near the surface, consistent with phase relations in Figure 3. Smectite and goethite (not

Figure 4. NETPATH model consistent with mineral distribution and phase chemistry for data from January 23, 1997. Normalized amounts (mmol/kg H20-m) of mineral dissolution (positive) and precipitation (negative) are given as a function of depth. pictured) precipitate throughout the vadose zone. Mass-balance calculations suggest that more silica must be removed from solution than can be readily removed by precipitation of clay minerals. This excess is indicated by "silica" precipitation in Figure 4. Although the vadose water is somewhat supersaturated with quartz (Fig 3), quartz does not readily precipitate at low concentrations; silica may be taken up by plants or adsorbed on clays (Pochatila et al. 2000). The rate of glauconite weathering can be calculated from the amount of mineral dissolution from NETPATH coupled with the residence time and vadose zone properties. The total amount of glauconite reacting, for depths from 0.5 to 4 m, ranged from about 0.3 to 0.5 mmol/kg H20, with higher values during the summer and fall months. The average rate of dissolution is given by:

R=

D(moZ / kg)* PPT(kg / c d ) * (I - E T ) /gm)*d(cm)*p(~~z/cmi)*F

T(S)*SA(M?

where D is the total amount of glauconite weathering (0.4 2 0.1 mmol/kg HzO), PPT is the rainfall during the residence interval (based on 0.145 kg/cm2 annual precipitation), ET is evapotranspiration loss (0.8), T is the residence time (30 to 60 days), SA is the geometric surface area (0.024 m2/g), d is the interval thickness (350 cm), p is the bulk density (1.5 g ~ m - ~and ) , F is the fraction of glauconite in the vadose zone (0.5). Substituting these values results in a calculated weathering rate of 4.6 2 0.6 x lO-I4 mol m-2 s-' 365

(0.046 pmol m-2 s") for glauconite in contact with vadose waters at pH 6.0, based on a formula containing Olo(0H)z. This calculated rate is a maximum based on the geometric surface area of the pelletized grains. The surface area of the individual glauconite clays is much greater (perhaps as much as 1 m2 g-I). Thus, the actual rate of glauconite dissolution may be between 4.6 x lO-I4 to as little as 1.5 x lO-I5mol m-2 s-I, if the total mineral surface area is used in calculations. The rate of glauconite weathering has not been measured in the laboratory. The rate of biotite dissolution at pH 6.0 in the laboratory is about 2 x 10-l2rnol m-2s-' (see Kalinowski & Schweda 1996), about 30 times faster than the rate in this study. Although not strictly comparable, this tends to confirm the observation that rates measured in the laboratory are greater that those in natural environments by one to two orders of magnitude. The calculated rate is similar to or less than rates from some previous field studies. For example, Paces (1983) calculated the weathering rate of Nafeldspar as being ca. l O - I 4 rnol m-*s-I. Velbel (1985) calculated a total weathering rate in forested watersheds of 1.2 x lO-I3 mol m-2 s-I. SwobodaColberg & Drever (1993) found a rate of silica release of 7.8 x 10-14rnol m-2 s-' at the Bear Brook site, Maine. The rate of silica release in this study is 8 x 10-15rnol m-2s-'. Differences between dissolution rates calculated in this and in previous studies may be related to mineralogy and water contact, in addition to problems of surface area and defect densities (e.g. Paces 1983, Swoboda-Colberg & Drever 1993). Clay minerals should react more slowly than do primary igneous and metamorphic minerals. Most previous studies were conducted in catchments, and weathering rates reflect both vadose and groundwater processes. In this study only vadose processes were investigated. Further, slow rates of water movement, and isolation of some pores and surfaces, may also allow equilibrium to be approached, or high concentrations of reaction inhibitors to slow dissolution rates. The net chemical denudation rate calculated from these data is about 0.7 cm/1000 yrs. This rate is much less than, for example, the ca. 4 cd1000 yr rate for metamorphic rock from the Southern Blue Ridge (Velbel 1985).

Based on temporal variations in chemical composition and calculations, the residence time of water in the sandy loam vadose zone was from 1 to 2 months. Glauconite weathering produces goethite, kaolinite, and smectitic clays. The most rapid rate of weathering was somewhat below the surface. The average weathering rate of glauconite in the Hornerstown Formation is 5 4.6 x lO-I4mol m-2 s-'. The chemical denudation rate is about 0.7 cm/1000 yrs. REFERENCES Courbe, C., B. Velde, & A. Meunier 1981. Weathering of glauconites; reversal of the glauconitization process in a soil profile in western France. Clay Minerals, 16: 23 1-243. Kalinowski, B.E. & P. Schweda 1996. Kinetics of muscovite, phlogopite, and biotite dissolution and alteration at pH 1-4, room temperature. Geochim. Cosmochim. Acta 60: 367-385. Pates, T., 1983. Rate constants of dissolution derived from the measurements of mass balance in hydrological catchments. Geochim. Cosmochim. Acta 47: 1855-1863. Pochatila J., B.F. Jones, & J.S. Herman 2000. Geochemical investigation of weathering reactions in South Fork Brokenback Run, Shenandoah National Park. EOS 8 1: S259. Plummer N.L., P.C. Prestemon, & D.L. Parkhurst 1991. An interactive code (NETPATH) for modeling net geochemical reactions along a flow path: U.S. Geological Survey Open-File Report 91-4078. Sverdrup H. & P. Warfvinge 1995. Calculating field weathering rates with a mechanistic geochemical model PROFILE, Applied Geochemistry 8: 273283. Swoboda-Colberg, N.G., & J.I. Drever 1993. Mineral dissolution rates in plot-scale field and laboratory experiments. Chem. Geol. 105: 5 1-69. Velbel, M.A. 1985. Geochemical mass balances and weathering rates in forested watersheds of the southern Blue Ridge. Am. J. Sci. 285: 904-930. Wolff, R.G. 1967. X-ray and chemical study of weathering glauconite. Am. Mineral. 52: 11291138.

5 CONCLUSIONS The composition of pore fluids in a glauconite-rich vadose zone in the Hornerstown Formation, central New Jersey, was monitored over a one-year period.

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Mineralogical evolution of bituminous marl adjacent to an alkaline water conducting feature at the Maqarin analogue site A .Cassagnabkre, J.C .Parneix & S .Sarnrnartino ERM, Poitiers, France

L.Y.Griffault Andra, Chatenay-Malabry,Fralzce

U .Maeder University of Bern, Bern, Switzerland

T.Milodowski B.C.S., Nottingham, United Kingdom

ABSTRACT: The Maqarin site, in Jordan, has been investigated over the last ten years as a natural analogue for the effects of an alkaline plume progressing through a fractured rock. Classical analytical methods have been applied to obtain the mineralogical and chemical evolution of the rock when leached by hyperalkaline fluids. Drilling of cored boreholes in the A6 Adit wall rock was attempted in order to collect a long core section crosscut by a single feature. A 30 cm long sample was obtained and is being investigated using classical analytical methods (petrography, SEM, XRD, porosity, B.E.T. measurements, chemistry). The overall distribution of the elements together with the textural modifications in the adjacent rock will be interpreted taking into account the potential rock matrix diffusion processes.

1 INTRODUCTION The Maqarin site in Jordan site has been investigated over the last ten years as a natural analogue for the effects of an alkaline plume progressing through a fractured rock (Alexander & Smellie 1998). The Maqarin site is an active system characterised by many superimposed fractures. Even if a sequence of mineralogical alteration could be established some difficulties were raised for the interpretation of the geochemical processes occurring in the adjacent rock. As well the extension of the perturbation in the wall rock of those fracture was difficult to estimate because of superimposition of fracture. The aim of further investigations was then to obtain the mineralogical and chemical evolution of the marl adjacent to a single water-conducting feature at the Maqarin site. This study reports on a petrographical and chemical study performed on a core section of about 30 cm long crosscut by a single fracture, using classical analytical methods (petrography, SEM, XRD, porosity, B.E.T. measurements, chemistry, UTH dating) (ERM 1997: Cassagnabere et al. 2000). The objective is to obtain a description of the distribution of the elements in the adjacent rock together with the mineralogical and textural modification. Interpretation will take into account the potential rock matrix diffusion processes.

groundwaters result from leaching of a cemented zone created during a high temperature and lowpressure metamorphism. The reconnaissance mission conducted in 1999 at Maqarin (Smellie 2000) allowed to perform core drilling in the A6 Adit at the Maqarin site. Coring succeed and a sample 30 cm long crosscut by a single vein was obtained. This core section was first cut along its length in two parts preserving the infilling products and the adjacent wall rock (Fig. 1). Five thin sections were performed along the profile including the fracture. Ten samples were picked in order to perform X-ray diffraction and other analyses along the profile.

Figure 1. Aspect of the core section with fracture on the left part. The maximum length is 27 cm.

The vein, 2 mm thick, is outlined by two distinct zones. The first one is about four millimeters thick and the second is fifteen millimeters thick, distinct with its light colour on the picture (Fig. I).

3 STUDY AND MEASUREMENTS

2 SAMPLING

3.1 Optical microscopy

The Maqarin site is a natural occurrence of alkaline groundwater penetrating bituminous marl. Those

Under optical microscope the rock is a biomicrite in which numerous foraminifera tests can be observed.

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The fracture 2 mm thick is infilled by CSH (Calcium Silicate Hydrates) displayed as flaky minerals and dark material having the aspect of organic matter. In the first adjacent zone, the biomicrite appears darker and some microcracks can be observed with an orientation sub-perpendicular to the fracture. In contrast, the second zone is lighter crosscut by slanting cracks filled with material similar to the one observed in the fracture. Optical microscopy didn’t allow identifying any mineralogical changes in the biomicrite beyond the second alteration halo. However an increasing amount of dark minerals corresponding to the organic matter was observed.

3.2 Scanning electron microscopy A S.E.M. JEOL 6400 equipped with an E.D.S analysis system was used in order to observe carbon coated thin sections. Elementary analyses were performed for the main chemical elements constitutive of the minerals present in the rock. Knowing the chemical composition of the different minerals identified by X-ray diffraction, it is possible to combine the different chemical maps in order to get a map of the mineralogy. This is performed with software developed by ERM (in progress). Considering the small size of some minerals, the probability to get pixels representing a mixture of minerals is to be taken into account. So possible binary mixtures are integrated in the database used to obtain the mineralogical map. At this point, fifteen minerals (or binary mixtures) can be integrated in that database. This technique allows getting the spatial distribution of the mineralogical phases constituting the rock (even nanometric ones) and their evolution along the studied profile (Fig. 2). Such a map allows to clearly distinguishing the different mineral components of each zone. Ettringite and tobermorite are present in the vein. Precipitation of ettringite extends up to 2 mm in the first alteration zone. Beyond, its occurrence is mainly as infilling of the microcraks. No ettringite is observed beyond 4 cm from the vein (Fig. 3).

3.3 X-ray diffractometry X-ray diffraction confirms the occurrence of CSH in the fracture: ettringite, 11 8, - tobermorite and 14 8, - tobermorite. In the wall rock, calcite is present in all the samples as the main mineral. However, its relative abundance seems to be higher in the first alteration wall rock. Another carbonate, siderite, can be detected in this zone. 11 8, - tobermorite was also identified in small quantities as well as probably cesanite (Ca2Na3(SO4)3(OH)). Other minerals occurring in too small mounts could not be detected using x-ray diffraction.

Figure 2 . SEM cartographic interpretation of the mineralogy in the vein and the nearest wall rock.

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no sufficient amount of material. Trace elements including rare earth were analysed with ICP-MS. Concerning major elements the following remarks can be made: - CaO content does not vary and is in agreement with the results of calcimetry ; - Si02 is lower (about 2 %) in the first zone near the fracture and remains constant farther in the wall rock. Concerning trace elements: there is an enrichment in V, Cr, Zn, MO near the vein (Fig. 4) in the two first samples and a depletion in Ni, Cu. Sr and Ba (Fig. 5). Ba is the only one notably more concentrated in the vein than in the nearest 3 cm wall rock.

Figure 4. Evolution of trace elements whose concentration increases near the vein according to the distance to the vein (cm).

3.6 Specific surface by NZ adsorption (B.E. T. method) The B.E.T. method was used in order to determine the specific surface of the samples with nitrogen adsorption (Brunauer et al., 1938). Measurements were performed on a MICROMERITICS (Asap 20 10) device. The results are shown in Table l. There is an important variability of the external surface of interaction (between 8.8 and 12.9 m2/g) along a zone of about 40 mm from the fracture. The values are lower near the fracture and higher between 8 mm and 35 mm far from the vein. Farther the specific surface is more or less stable near 10 m2/g. Figure 3. SEM backscattered electron photograph of the fracture, the first zone and the second zone.

3.4 Calcimetry

Values obtained with Bernard calcimeter are relatively high and homogeneous between 63.8 % and 68.3 %' o of calcite.

3.5 Chemical analyses The core section was cut accordingly to the different alteration halos distinguished by the SEM observations. Chemical analyses were performed as a function of available quantity of material. Major elements were obtained with XRF for samples A5 and A10, ICP for others, except A1 for which there was

Figure 5. Evolution of the concentration (ppm, except for Sr p p d l 0 ) Ni and Cu decreasing near the vein, and Ba and Sr having the higher concentration of trace elements in the vein.

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3.7 Hgporosimetry The porosity of the rock was measured by forced injection of mercury using a MICROMERITICS Autopore I11 9410 V1.02 (LEPTAB, University of La Rochelle) (Washburn 1921). The results are given in the Table 1. As for B.E.T., there is a great variability of values near the vein until about 40 mm from the vein. It is interesting to note that there is opposition of trends between the B.E.T. curve and the porosity one (Fig. 6).

Table 1. Specific surface (BET method) and Hg porosity values as a function of the distance to the fracture (D). Sample D(mm) 2 0-4 3 4-10 4 10-18 5 18-24 6 24-38 7 38-48 8 48-58 9 58-78 10 78-112

B.E.T. (m2/g) 8.810.1 1 1.510.1 9.810.1 12.910. 1 11.9AO.1 10.3AO.l 10.310.1 10.210. 1 10.510.1

Hg porosimetry (?A) 44.OAl 35.110.5 37.110.5 33.0A0.4 3 1.410.4 32.810.4 32.7A0.3 32.810.3 32.510.2

The nearest wall rock is characterised by the dominance of ettringite as pervasive precipitation. The second wall rock, lighter is marked by the presence of microcraks filled with ettringite. An enrichment in Zn, CryV and MO near the vein in the two first samples while there is a depletion in Ni, Cu, Sr and Ba. Barium is the only element notably more concentrated in the vein than in the wall rock. Concerning major elements, silica seems to decrease near the vein, indicating the dissolution of silicates and the precipitation of CSH in the vein. Calcium remains constant but as calcite concentration decreases near the vein, this element is probably integrated in other phases like ettringite and tobermorite. The presence of microcracks induces a higher porosity near the veins but the reactive surface area decreases as shown by BET values. This can be due to the dissolution of the major fluid pathways with a precipitation in the lower porosity. Overall results indicate that the alteration affects the rock on about 3 cm around the 2 mm thick vein, considering that all parameters begin to change between 3 to 4 cm away from the vein. The macroscopic evidence of alteration marked by changes of colours is about 2 cm. REFERENCES Alexander, W.R. & J.A.T. Smellie 1998. Maqarin natural analogue project: ANDRA, CEA, NAGRA, NIREX and SKB synthesis report on phase I, I1 and 111. Report 98-08. Cassagnabere A., Bouchet, A. & S. Sammartino 2000. Etude d’echantillons du site d’analogues naturels de Maqarin (Jordanie), ANDRA report D RP OERM 00 023. Brunauer, S., Emmet, P.H. & E.Teller 1938. Adsorption of gases in multimolecular layers, J. Am. Chem. Soc. 60: 1553. E.R.M. 1997. Techniques et protocoles analytiques utilises par la socikte ERM. Internal Report ERM 97 022 FR 060 (available near ERM, Poitiers). Smellie, J.A.T. 1998. Reconnaissance mission report, Maqarin natural analogue project: phase IV, SKB Report R-00-34. Washburn, E.W. 1921. Note on a method of determining the distribution of pore-size in a porous material. Proc. Nut. Acud. Sci.:7: 115-116.

Figure 6. BET and porosity curves according the distance to the vein (cm) ; proportion of ettringite by SEM mapping on the nearest wall-rock.

4 CONCLUSION Results obtained on a core section crosscut by a single vein are in agreement with previous results, i.e. occurrence of ettringite and tobermorite in the fracture. However, the different analytical methods applied allowed to better understand the water rock interaction in the adjacent wall rock of the fracture. The biomicrite mainly constituted by calcite, a few silicates, pyrite and organic matter displays the following alteration features: The hyperalkaline fluid lead to the deposition of ettringite, 14A-tobermorite and again 11 A-tobermorite in the vein. 370

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Oxidation of an argillaceous formation: mineralogical and geochemical evolution D Charpentier, M.Cathelineau & R.Mosser-Ruck UMR 7566 G2R-CREGU, BP23, F-54501 Vandoeuvre 1;s Nancy Cedex

G.Bruno* IPSNIDPREISERGD, BP6, F-92265 Fontenay aux Roses Cedex “Present address :IPSNfDESfSESID,BP6, F-92265 Fontenay-Aux-Roses

ABSTRACT: The evolution of the water-rock system at the surface of rocks in presence of atmospheric oxygen has been studied on the example of an indurated argillite sampled in the Tournemire underground experimental site. Penetration of water and atmospheric oxygen modifies redox conditions and subsequently the water-rock equilibrium at the surface of argillite. Although no macroscopic alteration was suspected careful microscopic observations show that systematic mineralogical changes occurred at the surface of the exposed rocks: pyrite shows marks of oxidation and a typical association of oxidation products (gypsum, celestite, jarosite, goethite) has been identified. Experimental and numerical modeling has been carried out to reproduce the oxidized rock-water system, and the role of newly formed assemblage on the water chemistry. 1 INTRODUCTION The evolution of the water-rock system at the surface of rocks in presence of atmospheric oxygen after digging of an underground work has been studied on samples of an indurated argillite from Tournemire (France). At this locality, a railway tunnel, built about 100 years ago through the Tournemire Massif, and new galleries are used as a laboratory for the generic scientific and geotechnical investigations in underground conditions. The Tournemire site is located in southern Aveyron (France) in subhorizontal Jurassic sedimentary rocks. These sedimentary rocks comprise a 250 m thick indurated Toarcian and Domerian shale unit between two aquifers (Cabrera 1991, Barbreau & Boisson 1993, Boisson et al. 1996). The shale from Tournemire is characterized by very low hydraulic conductivities and apparent diffusion coefficients (Boisson et al. 1996). Within the clay-rich matrix of the shale, pore-water contents are very low (around 3.5%, Bonin 1998, De Windt et al. 1998). Water circulation is low and essentially driven by diffusion (De Windt et al. 1998). Modifications of the physico-chemical conditions in the excavated zones yield to changes in rock properties. The presence of oxygen causes a disturbance of the redox conditions and subsequent Eh-pH changes in the pore fluid due to water-airminerals interaction Dehydration affects rocks over a significant thickness from the tunnel (or gallery) walls. In the case of pyrite-rich argillites, processes at the pyrite-fluid interface may yield significant

mineralogical changes at the surface of exposed rocks (Blowes & Jambor 1990, Cathelineau et al. 1995, Roussel et al. 1999, among others). However, the presence of oxidation products, and the degree of alteration was still debated at Tournemire. This motivated a detailed study of the rock surface using light and electron microscopes, TEM, electron microprobe, and XRD. Both experimental study of water-rock interaction and numerical modeling of the water-rock reaction in oxidizing conditions have been carried out in order to estimate the mineralogical and chemical changes during oxidation. 2 EVOLUTION OF THE MINERALOGY BY OXIDATION The mineralogical composition of studied samples is similar to the one obtained by previous authors (Bonin 1998, De Windt et al. 1998). The argillite contains about 55% clay minerals (illite and mixed layered I/S, kaolinite, biotite, K-micas, chlorite), 20% quartz, 15% carbonates (bioclasts and diagenetic calcite, dolomite and siderite), Kfeldspars, pyrite. However, the study of rock surfaces from the tunnel and gallery walls reveals significant changes in the mineralogy and textures. At the surface of argillite blocks, pyrite shows marks of oxidation such as dissolution pits mainly observed on surfaces of cubic pyrite. A typical association of oxidation products has also been identified: 371

Procedure 2: Samples 2 and 3 come from milled oxidized surfaces. Leaching was carried out under atmospheric conditions. Procedure 3: Sample 4 was obtained after drilling of the argillite and leaching was carried out under atmospheric conditions. Leaching of sample 1 provides the composition of water interacting with a rock not affected by oxidation. In contrast, leaching of samples 2 and 3, on one hand, and sample 4, on the other hand, provide water compositions interacting with rocks affected by long duration of oxidation (a 100 years for tunnel walls) and short exposition to air (a few hours of the experiment) respectively. The concentrations of cations in the solutions were measured by ICP-AES, and anions were analyzed by ion chromatography. Modeling was carried out using the geochemical code EQ3/6 (Wolery & Daveler 1992) with the following assumptions: temperature of 25°C and closed system behavior with constant oxygen fugacity (logf,,= - 0.68). The initial fluid is taken as pure water with pH = 5.8. The assumed mineralogy is the one of the oxidized argillite, composed of quartz, calcite, dolomite, pyrite, smectite, illite, kaolinite, chlorite and K-feldspar and the two main oxidation products (gypsum and jarosite).

- Gypsum (CaSO,. 2H,O) is detected in great amount. The amount of gypsum, the morphology and size of the crystals vary with the location of the samples. Close to the tunnel, the argillite is covered by macroscopic crystals and by small size needleshape crystals. - Celestite (SrSO,) is observed as small euhedral crystals in samples located near the tunnel and on fracture surfaces. - Jarosite (KFe,(SO,),,(OH),) is mostly observed on the surfaces of altered pyrite-rich zones such as diagenetic pyrite-calcite fractures. - Brown color patches enriched in iron hydroxides around pyrite crystals are indicative of iron diffusion in the nearby environment of oxidized pyrites. Wellcrystallized Fe-hydroxides, identified as goethite, have been detected onto the surfaces of samples close to the tunnel. Oxidation does not involve significant mineralogical modifications of the clay particles. The same clay phases are present both in preserved and oxidized argillite. However, microprobe and TEM analyses of the fine clay particles ( c 2 pm) show a weak chemical evolution after oxidation, in particular an increase in the silicium content and a decrease in the interlayer charge of the (2:l) clays. These transformations involve modifications of the mixed layer (US) nature, which could be explained by a preferential dissolution of illite layers.

3.2 Results

3 CHANGES IN PORE-WATER CHEMISTRY IN PRESENCE OF OXIDATION PRODUCTS 3.1 Principles and conditions The modification of the interstitial fluid in the oxidized argillite has been investigated through water-rock experiments. The liquid/rock ratio (L/R) was enhanced, and water-rock reaction studied under varying L/R ratios. Several parameters were tested during the experiments. Rock grains (2 mm c size c 3 mm) were put, together with the leaching solutions, in 100 ml vessels at variable L/R ratios from 2 to 50. The mixtures were shaken from 1 min to 10 days. 0.45 pm millipore were used to filtrate the solutions. Leaching experiments were carried out on rocks preserved from oxidation, and also on samples displaying the oxidation assemblage, in order to determine the influence of oxidation products on the chemistry of the fluids. We followed three experimental procedures. Procedure 1: Sample 1, obtained after drilling of the argillite, was sealed in evacuated AI-coated plastic sheets at the sampling site and has been opened, crushed and leached within a glove box, under Ar atmosphere.

The evolution of the element concentrations is strongly dependent on the leaching conditions (time, liquidholid ratio, redox condition). Calcium and sulfate are the most sensitive elements to leaching conditions and, in what follows, they are chosen to illustrate the parameters controlling the chemical characteristics of the analyzed solutions.

Figure 1. Evolution of Ca and SO, concentrations in leaching fluids versus WR ratio (case of the sample 4, time of reaction = 10 days).

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Effect of the LIR ratio The calcium and sulfate concentrations in the solutions increase with decreasing liquid/solid ratio (Fig.1). Nevertheless, the difference in behavior between sulfate and calcium (depicted by the difference in slope in Figure 1) indicates that dilution is not the only parameter controlling the evolution of concentrations for increasing L/R. Other potential controlling parameters may be variations in accessibility to dissolved minerals, precipitation of new minerals or exchange processes. Effect of water-rock interaction duration: case of sample 4 For short experiments, the sulfate concentrations increase whereas the calcium concentrations stay almost constant in aqueous solutions (Fig.2). For experiments spanning more than 12 hours, the Ca concentration increases and shifts to the Ca = SO, line. Effect of the redox conditions The fluid of experiments carried out on oxidized surfaces (samples 2 and 3) shows high concentrations, close to saturation with respect to gypsum (Fig.2). When the experiments were carried out under Ar-atmosphere on sample preserved from oxidation (sample l), the sulfate and especially calcium concentrations are significantly lower. Under atmospheric conditions on a drilled sample (sample 4), the leaching solutions have intermediate concentrations, involving that oxidation during a few hours to a day was enough to yield significant

modifications of element concentrations. The Ca and SO, concentrations in solution are not on the Ca = SO, line. Part of the calcium generated by gypsum dissolution can be fixed by another mineral. Alternatively, dissolution of gypsum might not be the only source of sulfates, as other sulfates are present. Ca and SO, concentrations in solution tend to the stability field of gypsum when the duration of the experiments is increased. The gypsum present in altered samples is partially or totally dissolved after 10 days of leaching. At this stage, it becomes the main phase contributing to the changes in water chemistry which, in turn, becomes progressively enriched in Ca, (K, Mg) and SO,. From Ca and SO, concentrations in solution, it is possible to calculate the amount of gypsum formed by oxidation of pyrite and then to estimate the rate of pyrite alteration. This calculation leads to 3 to 10%. The evolution of Ca and SO, contents in solution obtained from numerical modeling is reported in Figure 2 (dotted line). The modeled line is parallel to the trend of experimental Ca and SO, concentrations and shows a good agreement between experimental and numerical modeling of oxidation processes.

4 CONCLUSION Argillites from a tunnel excavated 100 years ago show clear evidences of pyrite oxidation, with subsequent formation of gypsum and other phases (celestite, jarosite, iron hydroxides). Leaching

Figure 2. Ca and SO, concentrations of leaching fluids and modeled evolution of the fluid composition.

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experiments on argillite affected by various degrees of oxidation (from short duration experiments to “natural” long duration water-rock interaction) show that: i) significant oxidation of pyrite occurs after a few hours of exposition to the air, and the Ca/SO, balance in leaching experiment is obtained after a few hours, ii) the amount of formed gypsum is related to the duration of the air-rock interactions. These processes are of critical importance for the consideration of mineral equilibrium in the excavated disturbed zones and especially for the prediction of the future behavior of the excavated disturbed zones during re-hydration.

ACKNOWLEDGMENTS This work has been carried out through a research contract with IPSN, and has been supported partly by GdR FORPRO (Action 99-1).

REFERENCES Barbreau, A. & J.Y. Boisson 1993. CaractCrisation d’une formation argileuse/synth&se des principaux rksultats obtenus partir du tunnel de Tournemire de janvier 1992 a juin 1993. Rapport d’avancement n”1 du contrat CCE-CEA n”FI2W Cl91-0115: EUR 15736FR. Blowes, D.W. & J.L. Jambor 1990. The pore-water geochemistry and the mineralogy of the vadose zone of sulfide tailings, Waite Amulet, Quebec, Canada. Applied Geochem. 5: 327-346. Boisson, J.Y., Cabrera, J. & L. De Windt 1996. Investigating faults and fractures in argillaceous Toarcian formation at the IPSN Tournemire research site. Workshop on Fluid jlow through faults and fractures in argillaceous formation ”,Bern. Bonin, B. 1998. Deep geological disposal in argillaceous formations: studies at the Tournemire test site. Contam. Hydrol. 35: 315-330. Cabrera, J. 1991. Etude structurale dans le milieu argileux: LEMI du tunnel de Tournemire. Rapport n”3 du contrat CEA- CABRERA . Cathelineau, M., Guerci, A., Ahamdach, N., Cuney, M., Mustin, C. & G. Milville 1995. Smectite-goethite-Fe, U, Ca and Ba sulfate assemblage resulting from bio-oxidation derived acid drainage in mine tailings: an analytical, experimental and numerical approach. Proc. of the third biennal SGA meeting: 647-649. De Windt, L., Cabrera, J. & J.Y. Boisson 1998. Hydrochemistry in an indurated argillaceous formation (Tournemire tunnel site, France). In Arehart, G.B. & J.H. Hulston (eds.), Proc. W.R.I.-9: 145-148. Rotterdam: Balkema. Roussel, C., Bril, H. & A. Fernandez 1999. Evolution of sulphides-rich mine taillings and immobilization of As and Fe. C. R. Acad. Sci. 329, 11: 787-794. Wolery, T.J. & S.A. Daveler 1992. EQ6, a computer program for reaction path modeling of aqueous geochemical systems: user’s guide and documentation. Lawrence Livermore National Laboratory, University of California. 337p. “

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Microscopic processes at the interface between metal sulphides and water Giovanni De Giudici Dipartimento di Scienze della Terra-Universitci di Cagliari, Italy

Pierpaolo Zuddas Lab. de Gebchimie des Eau-CNRS 7047-Institut de Physique du Globe de Paris and Univ. Paris 7 France

ABSTRACT : In this work, the reaction of galena dissolution in oxygen saturated and acidic solutions is investigated by coupling Atomic Force Microscopy (AFM) with in situ solution chemistry measurements. In the first hours of microtopography investigation in the AFM liquid cell, square etch pits grow on the (001) investigated surface. After this short stage, the direct crystallographic control upon surface features is lost, and the surface is characterised by microroughness. Interacting solutions were strongly undersaturated in both galena and anglesite. Then, we calculated dissolution rates from lead concentration in solution and found a continuous decrease in kinetic rates by one order of magnitude. Surprisingly, we also found that the thermodynamic driving force governing dissolution does not depend upon galena nor anglesite. Using thermodynamic and X P S literature data, we predicted that protrusions forming microroughness are made of native sulphur. Finally, we interpreted that dissolution of protrusions is the rate limiting step of the overall reaction, and protrusions constitute a microscopic geochemical reservoir that rules the processes of galena dissolution.

1 INTRODUCTION

2 EXPERWLENTAL METHODS

Galena is one of the main minerals controlling the release of lead and other toxic metals to solutions circulating at the Earth surface, where the conditions are often acidic and oxidising. The overall process of galena dissolution involves both acid-base and redox reactions, and is further complicated by the high number of electrons involved in the irreversible oxidation of sulphide to sulphate. Previous works have shown that galena surface reactivity depends upon surface stoichiometry and presence of trace impurities (Tossel & Vaughan 1987, and references therein), and solution pH (Hsieh & Huang 1989). More recently, Scanning Tunnelling Microscopy has shown that surface dissolution in electrochemical oxidising conditions proceeds via retreat of steps oriented along [ 1101 crystallographic directions (Higgins & Hammers, 1996). In this work, we investigated the reaction of galena surface dissolution in oxygen-saturated solutions by in situ Atomic Force Microscopy (AFM) and solution chemistry measurements, both carried out together on the same sample. The questions that we would like to answer are the following : 1) do surface features remain constant during the interaction, and 2) is the dissolution rate constant during the two days long period of investigation?

A specimen of natural galena was carefully cleaved in the laboratory and then imaged by in situ AFM (Molecular Imaging) equipped with Digital Instrument software. The oxygen saturated solution circulating in the liquid cell of the AFM was continuously renewed at a constant flow rate. The ratio between solution volume and mineral surface area was assumed to be constant and the initial solutions did not contain lead nor sulphur. Thus, for a given experiment, dissolution rate (mol m-*sec-'), R, was measured as: R = (,b2' x flowrute)l Area (1) The two initial solutions (pH 1 and 3) were prepared by adding HCl to distilled water under room atmosphere . Solutions flowing out from liquid cell were periodically sampled and analysed for dissolved lead (GFAAS) and pH value. 3 RESULTS 3.1 Solution chemistry

In the experimental conditions, the stable phase of sulphur in solution is sulphate, and the overall

375

dissolution of galena proceeds in oxidising solutions according to the following irreversible reaction :

galena or anglesite. By contrast, we interpreted that thermodynamic driving force is ruled by microscopic phases, other than galena, forming at the interacting surface.

However, while Reaction 2 is the overall process, the oxidation of sulphur to sulphate implies the loss of eight electrons and should then proceed via intermediate steps. In previous works, the process of galena surface interaction with acidic and oxygen saturated solution has been summarised by Hsieh & Huang (1989) and Fornasiero et al. (1994) as a twostage process, composed of a fast adsorption of hydrogen ions at the surface followed by adsorption of oxygen that produces the oxidation of sulphur at the surface or the interface. In our AFM experiments, initial solutions did not contain lead nor sulphur, and the concentrations of these elements in solution depend exclusively upon the reactivity of galena surface. Lead concentration in solution was detected by GFAAS, while sulphur concentration was our below detection limit. Dissolution rates were calculated (Eqn. 1) and plotted as a function of time. Figure 1 shows that dissolution rates decrease by more than one order of magnitude over the 45 hour interval of investigation. The observed change in kinetic regime at constant pHs, suggests a change in the reaction mechanism.

3.2 Microtopography measurements

Figure 1. Log of dissolution rates (nanomol m-* sec-') as hnction oftime (hours) for the conditions of pH=l and 3.

Although sulphur was not directly determined, we can reasonably assume that sulphate is 90% of total dissolved sulphur and sulphide is 10%. In such a way we calculated solution saturation state with respect to both galena and anglesite finding strong undersaturation (these two values are below 2x 104 and 2x 10'3, respectively). Surprisingly, we also found that dissolution rates decrease with solution saturation states and, thus, should not depend upon thermodynamic driving force ruled by dissolution of 376

After careful cleavage of our galena specimen and the 5 min necessary to optimise the instrument parameter, the fresh (001) surface was imaged by AFM Contact Mode (see De Giudici & Zuddas 2001). In the first hours of AFM microtopography measurement during dissolution at pH=l we observed nucleation and growth of square etch pits (Figs 2a and 2b). After this short stage, surface features evolve by widespread formation of protrusions 1-3 nm high and homogeneously distributed, while nucleation and growth of new etch pits is inhibited, and large rough terraces delimited by macrosteps are formed by dissolution (Fig.s 2c, 2d and 2e). At the conditions of pH=3, we observed the same surface behaviour (Fig.s 2f and 2g). However protrusions showed rounded shape.

4 DISCUSSION Because of their appearance on AFM images, protrusions are made of phases that dissolve slower than galena. Since the residence time of interacting solution in the liquid cell is shorter (about 30 sec) than the time needed for deposition of small particles, we argued that protrusions are not produced by precipitation from bulk solution but rather by interfacial reactions. On the other hand, if the dissolution reaction proceeds via the production of HS- and its migration through the interface towards the bulk solution, surface features should be controlled by the primary crystallography of galena. However, our AFM data clearly indicate that crystallographic control is lost after some hours of interaction, and surface features are governed by protrusions. We interpreted that the protrusion formation is due to interfacial reactions, and inhibits direct crystallographic control upon surface features by coating the surface. We propose that the reaction of protrusion dissolution represents the step limiting the rate of the overall galena dissolution, and the evolution of the rate regime is a transitory stage due to the continuous increase in the thickness of the altered layer that coats the surface. From thermodynamic and X P S data in the literature, we predicted that nanometric protrusions are composed of native sulphur, an intermediate and slowly dissolving phase. In turn, the sulphur reacts with oxygen to dissolve and migrates towards bulk

solution in a more oxidised state (eg. as SO:-, s20,2-, etc.).

377

REFERENCES De Giudici G. & Zuddas P. In situ investigation of galena dissolution in oxygen saturated solution : evolution of surface features and kinetic rate. Geochim. Cosmochim. Acta, in press Fornasiero D., Fengsheng L., Ralston J. and Smart R. (1994 ) Oxidation of galena surfaces. I. X-Ray photoelectron spectroscopic and dissolution kinetic studies. J. Colloid Inter- Sci. 164, 333-344 Higgins R.S. & Hamers R.J. (1996) Chemical dissolution of the galena (001) surface observed using electrochemical scanning tunnelling microscopy. Geochim. Cosmochim. Acta 60, 3067-3073. Hsieh Y.H. & Huang C.P. (1989) The dissolution of PbS in dilute aqueous solutions. J. Colloid and Inter$ Science(% 1, 537-549. Tossel J.A. & Vaughan D.J. (1987) Electronic structure and the chemical reactivity of the surface of galena. Tke Canad. Min. 25, 381-392. Figure 2. Galena surface interacting with oxygen saturated solution at pH 1 (from a to e) and pH 3 (f and g). The tip was scanned over the (001) surface only for the time necessary to collect the images. Galena surface features after 28 min (Fig. 2a) and 38 min (Fig. 2b) of interaction are characterised by square etch pits that grow with an horizontal velocity of 15 nm/min and vertical velocity of 1.5 d m i n (estimates by measuring width and dept of surface features using DI image analysis software). Fig. 2c shows the surface after 17 hours of interaction. Fig. 2d is a zoom made on Fig. 2c, and clearly shows the presence of protrusions 1-3 nm high that cover all the surface. Fig. 2e shows the surface after 40 hours of interaction, etch pits have been redissolved and surface is even characterised by microroughness. At the condition pH=3, surface evolution is very similar, but protrusions have a rounded shape (Fig. 2f and 2g, after 10 min and 40 hours respectively).

5 CONCLUSIONS

During galena dissolution in oxygen saturated solutions, the crystallographic control upon surface features is lost after some hours of interaction. Thermodynamic driving force is not ruled by galena or anglesite, but rather by nanometric phases forming at the interacting surface, most likely sulphur. This study confirms that AF'M is a useful method to understand the microscopic processes ruling mineral-water interaction also in non-steady-state systems like galena dissolution in oxygen saturated solutions. In future work, the experimental apparatus used in this investigation can be used to assess rate laws at both laboratory and field scales for metal sulphides and other geochemically important minerals.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Dissolution of calcite in CaC03-C02-H20 systems in porous media P.A .Diaz & V.Alvarado PDVSA Intevep, Los Teques, Edo. Miranda Venezuela

M .I. Rodriguez U.C.V. Facultad de Ciencias, Caracas, Venezuela

ABSTRACT: The study of many geological processes requires the understanding of the dissolution and pre~ O under dynamical conditions. These processes take place cipitation of calcite in C ~ C O J - C O ~ - Hsystems during the deposition of carbonate scales in sandstone reservoirs, which turns out to be an operational problem in the oil industry. This paper presents the results from some calcite dissolution experiments performed in columns packed with calcite in a quartz matrix that simulates sandstone composition. Carbonated water was pumped and then recirculated through the column until calcium saturation in the system was reached. The experiments were designed to investigate the effect of the fluid flow rate, grain size of the sand, CO2 partial pressure and temperature of the system, on the calcite dissolution rate. A correlation between the studied parameters effects is observed. 1 INTRODUCTION The dissolution of calcite in CaCO,-CO2-H20 systems has been widely studied due to its importance in natural processes involving seepage water and groundwater, such as weathering and karstification (see for instance Dreybrodt 1980, Buhmann & Dreybrodt 1985a, b, Domenico & Schwartz, 1998). This system has also paramount importance in the oil industry, due to its relevance to the formation damage originated by the precipitation of carbonate minerals. This phenomenon, known as carbonate scale deposition, deteriorates the permeability of the rock formation and consequently affects the hydrocarbon recovery. The deposition mechanism of the carbonate scale may involve the dissolution of natural carbonate minerals present in some sandstone reservoirs, by their interaction with flowing-through aqueous solutions (i.e. production or injection water). Subsequent precipitation occurs when the resulting carbonate solution is transported towards regions with lower pressure and temperature, due mainly to changes in CO2 fugacity. In addition to the thermodynamic parameters, the transport phenomena may play an important role in the dissolution of calcite in these systems. The use of models developed to describe the mechanisms of heterogeneous reactions provides a conceptual guide to understand the coupling between transport phenomena and chemical kinetics. The adsorption-layer model has been used to explain

the dissolution and precipitation of carbonate minerals (Plummer et al. 1979, Dreybrodt 1980). This model assumes a thin adsorption layer separating the mineral surface from the hydrodynamic limiting layer. Chemical species located in the adsorption layer are linked to the mineral surface, thus they are less movable than those species from the hydrodynamic layer. Reaction mechanisms are proposed considering species migration. from and towards the mineral surface. Palciauskas & Domenico (1976) studied mass transfer, solution chemistry and chemical equilibrium in porous carbonate rocks. They solved the transport equation for calcium in the pore fluid of a porous structure, introducing a fluid-solid reaction as a source term. Buhmann and Dreybrodt (1985a, b) proposed a parallel-planes model to describe the dissolution process between two grains of calcite; they consider the chemical kinetics at the fluid-solid interface and the hydrodynamic of a solution flowing through the grains. These models state how diffusion, mechanical dispersion or convection may control the dissolution and precipitation of calcite under different fluid flow regimes and porous media geometry, and how their effect may be coupled with chemical kinetics. The application of these concepts is displayed in the work of Baumann et al. (1985) related to the study of calcite dissolution kinetics in columns packed with pure calcite, under various flow conditions. Their work is relevant to carbonate rock systems; 379

further investigations have to be made for sandstone systems. This work presents an experimental study of calcite dissolution kinetics in columns packed with this mineral in a quartz matrix that simulate sandstone composition. The results are interpreted under the light of the models mentioned above.

Table 1. Properties of the columns packed with calcite and quartz. Column Grain size

2 DISSOLUTION EXPERIMENTS

* In order to study the calcite dissolution, a physical model consisting on columns packed with calcite in a quartz matrix was used to simulate the reservoir rock. Some calcite dissolution experiments were performed injecting and recirculating carbonated water through the columns. Parameters such as grain size of the packing material, CO2 partial pressure (PC02) and temperature of the system, and the flow rate of the calcareous solution through the column, were varied in the experiments. 2.1 Calcite-quartz columns Slurries prepared with a calcite and quartz mixture (1:9 in weight) and water were used to fill columns (7 cm long, 0.9 cm internal diameter). To reduce grain segregation, the same range of grain size for calcite and quartz was used in each column (Table 1). During the packing process, grain settling was aided by means of an ultrasound bath. The resultant columns were characterized in terms of their permeability (K), porosity ($) and the surface area of the calcite present in the column (Table 1). Total pore volume, and henceforth porositiy, was determined as the difference between the column volume and the volume occupied by the calcite and quartz grains. Permeability was calculated using Darcy's Law: q=K-

Calcite

%WxW prn 1 38-63 9.04 2 125-200 10.45 3 200-300 10.80 4 38-63 10.10 5 200-250 10.12 calcite surface area

4

K

0.51 0.44 0.43 0.25 0.33

Surface area'

darcies 0.18 >66.6 >66.6 0.044 2.96

m2 109 78 11 109 10

2.2 Experimental conditions Figure 1 shows a sketch of the experimental equipment used in the dissolution experiments. The parameters varied in each experiment are presented in Table 2. PC02 was provided by bubbling c 0 2 - N ~ mixtures into distilled water located in the reservoir (160-ml). Once the resulting solution reached a constant pH value, it was recirculated through the calcite/quartz packed column. Gas supply to the solution was maintained to the end of each run.Calcium concentration and pH measurements were registered during the experiments using a Ca-specific and a calomel electrode, respectively. Table 2. Parameter levels varied during the experiments. Parameter level 1 2 qi (10" cm3/s) 7.2 31.2 10 22 7;: ("C) PC02 (10-3atm> 3.8-5.3 0.08-0.2

3 4 39.2 78.4 25 30 1.0-1.1 --

5 7.8 35

--

AP PL

where q = fluid flow rate (cm3/s); AP = pressure drop (atm) between the column inlet and outlet; p = fluid viscosity (centipoise); L = column length; K = permeability (darcies). A more effective packing process, as result of longer ultrasound treatment, leads to columns with lower porosity and permeability (see rows 4 and 1 in Table 1). Pictures of calcite grains obtained by scanning electron microscopy were digital processed to determine the average perimetedarea ratio of the grain images, P,,,/A. The surface area of the calcite present in the column, S, was estimated using this ratio, as:

Figure 1. Schematic view of the experimental equipment for calcite dissolution. 1 = calcite/quartz column; 2 = solution reservoir; 3 = magnetic stirrer; 4 = peristaltic pump; 5 = pH and calcium specific electrodes; 6 = COz or C02/N2mixture; 7 = thermal bath.

The experiments were stopped after the solution had reached the apparent e uilibrium calcium concentration [Ca2'],. The [Ca9+] profile for each experiment was adjusted with the following function: [Ca2'] = [CaZ'],(r-Be-"')

(3) where B and m are parameters from each reaction (Table 3). This function has the form of the solution

where Y = calcite volume; W = calcite weight; p = calcite density.

380

for the one-dimension reactive-transport model proposed by Baumann et a1.(1985)for porous calcite. PC02 was calculated for each experiment from the equilibrium expression involving [CaZ+], and [&Ie (Stumm & Morgan 19Sl), the error involved in this calculation is estimated to be about 10%.

On the other hand, the thickness of the equivalent adsorption layer of the calcite grains 6, was estimated as the ratio D,la, where D, is the molecular diffusion coefficient. This definition corresponds to that given by Baumann et al. (1985) to half the distance between calcite grains, as a limit for the adsorption layer thickness. In the present work 6 may coincide with this definition when a calcite grain faces another calcite grain, otherwise, when quartz grains surround a calcite grain, 6 corresponds to the grain separation distance. Therefore, 6 values calculated in the present work estimate the effective thickness of the adsorption layer. In order to estimate the calcite dissolution extent in each run, the following model is considered. Each experiment may be thought as run in a very long column resulting from the combination of the loops. The time t0.98 and the equivalent linear distance xO.98 required to reach 98%[Ca2’], are calculated and used to compare the extension of each reaction.

Table 3. Equilibrium calcium concentration and pH, calculated PC02, and calcium profile function parameters for each experiment. Experiment [Ca”], pH, PCO, I o - mol/I ~ 10”atm

[Ca”],’

B

~ O - ~ ~ O U I

m 10%’

1 Plql’* 4.1 5.96 4.2 4.17 0.96 2.8 5.3 4.12 1.01 12.4 5.89 1 Plq2 4.1 5.96 3.8 3.73 0.96 30.6 1 Piq4 3.7 6.99 0.08 2.09 1.05 1.5 1 P2ql 2.0 6.90 0.10 2.14 1.10 3.8 1 P2q2 2.0 1 P2q4 1.4 6.91 0.12 1.43 1.04 6.9 5.79 5.3 5.53 1.19 14.8 2Piq3 5.5 5.90 4.2 5.48 1.34 16.6 2Plq4 5.5 6.89 0.19 1.22 0.89 4.0 2P2q3 1.2 3.1 9.64 1.00 6.1 5.81 3 Piq2 9.4 3.4 9.40 1.02 8.1 9.4 5.79 3 Piq4 4.42 6.22 1.1 1.03 3.6 3 P3q2 4.2 1.0 2.17 1.04 6.43 1.7 3P3q4 1.8 5.75 7.4 5.04 1.00 11.0 4 Tlq, 5.0 5.86 6.5 4.46 1.04 11.9 4T2q5 4.5 5.89 4.8 4.11 1.03 7.9 4T3q5 4.1 5.88 6.2 3.33 1.08 9.6 4T4q5 3.3 5.3 5.64 1.03 7.0 5.6 5.80 5 Tlqs 4.1 5.08 1.07 9.5 4.9 5.69 5 T2q5 5.88 10.8 4.68 0.96 5.1 5T3q5 4.5 3.70 1.08 10.1 3.7 5.72 11.6 5T4q5 * equilibrium calcite concentration from profile adjustment **lPIql run in column 1, PCOz and q in level 1 as in Table1

’

The rate at which CaZf is produced by chemical reaction at the solid-fluid interface r, (mol l-’s*’),may be expressed in terms of the Cak flux from and towards the calcite surface, F, as:

(4) where V, = total pore volume; a = constant defined by the diffusion coefficient and the separation between the calcite grains (Buhmann & Dreybrodt 1985a, b). A reaction constant k is defined in terms of S, V, and a,thus, equation 4 may be rewritten as: r, = k{[Ca’+],- [Ca”]}

Table 4. Calculated and corrected reaction constants, calcite surface, time and equivalent distance to reach 0.98[Ca2+] for each column experiment. S ~1 6* to.98 ~ 0 . 9 8 Exp. k‘ k 1 0 - S-I ~ IO-~S-’ cm2 10-~cms-’ cm 104s m 1 Plqi 2.7 1.5 98 0.44 0.32 14.5 27 1 Plq2 12.4 6.7 105 1.89 0.07 3.2 26 1 Piad 30.4 16.4 109 4.46 0.032 1.3 26 1 Pi& 1.5 0.8 74 0.32 0.43 26.1 49 84 0.70 0.20 10.6 86 1 P2q2 3.7 2.0 94 1.17 0.12 5.6 116 1 P2q4 6.9 3.7 0.047 2.7 32 9.2 78 2.99 2 Plq3 14.6 56 0.4 70 3.79 0.037 2.3 2Piq4 16.6 116 0.084 9.8 2.5 38 1.68 2 P2q3 4.0 156 3.9 11 8.87 0.016 6.4 3 Plqz 6.1 47 5.2 10 12.96 0.011 4.8 3 Plq4 8.1 11 5.24 0.027 10.9 264 3 P3q2 3.6 2.3 0.050 28.0 270 8 2.80 3 P3q4 1.4 0.9 3.6 15 4.5 106 0.60 0.13 4Tlq5 10.9 3.4 14 4.8 98 0.64 0.20 4 T2q5 11.6 4.9 21 89 0.43 0.33 4T3q5 7.9 3.2 4.2 17 97 0.52 0.31 4T4q5 9.4 3.9 18 8 4.16 0.019 5.6 5 Tlqs 6.9 2.2 9 5.66 0.023 4.2 13 9.4 3.0 5 T2q5 24 1.6 9 3.04 0.046 7.7 5 T3q5 5.1 9 6.03 0.027 3.9 12 5 T4q5 10.1 3.2

* D,= 0.7 x 10‘5(0.56+ 0.058T)cm2/s,T (“C) Buhmann & Dreybrodt (1985a)

(5)

2.3 Fluid flow rate and particle size effects

The d[Caz+]ldt in the system may be related to r,

The reaction constant k increases while the thickness of the adsorption layer 6 decreases as q is higher for experiments performed in columns with the same particle size, PC02 and T, (lPlql, lPlqzand 1Plq4,Table 4). The removal rate of the species from the calcite grain surface determined the dissolution rate. The species are removed quickly from the adsorption layer as q increases, therefore this layer has less chance to grow and the chemical reaction is also accelerated. At higher q, diffusion controls the proc-

as: d[Ca2’] V, dt V p - rs

where Vr = total volume system. Rearranging 5 and 6:

(7) where k’ = kVdV,(Table 4).

381

ess. In the same set of experiment, the saturation distance ~ 0 . 9 8is the same for the three experiments, although it is reached at different times, which is determined by q. However, in experiments run in column 2, with a larger particle size, the q effect was not clearly observed, as other variables such as particle size and PC02 were also affecting the process. The extent of the grain compaction achieved during the column packing process affects the porosity and permeability, consequently it affects 6; in run lPtql 6 was about three times thicker than in run 4T2q5. As pore size in column 4 was smaller, the 10cal fluid velocities were higher, and the species in the pore fluid were more disturbed by the fluid flow, thus the adsorption layer was unable to grow thicker. Comparing the results from experiments run in columns with different grain size and similar q, PC02 and T (1Plq4, 2P&, 3Plq4, Table 4), it was found that the dissolution reaction constant is larger for smaller particle size columns. The calcite surface area in contact with the solution is larger, providing more reaction sites, therefore the reaction rate increases. On the other hand the [Ca2']e is higher in columns with larger particle size (Table 3). In this case, the pore fluid volume is also larger, thus the bulk fluid volume is less affected by the chemical kinetics taking place in the fluid-solid interface which allows to dissolve more calcite.

2.4 PC02 and T effects The experiments run at higher PC02 reached a higher [Ca2'], and 6. Under the experimental conditions used in this work, the CO2 content limited the extent of the calcite dissolution; longer reaction times were needed to saturate the solution in experiments run at low PC02 (i.e. 3p3q2,3p3q4). As the H2CO3 was the only acid present in the system, the final pH was the result of the reaction equilibrium involving the C02; experiments run at lower PC02 exhibited a higher pH, this may be explained by looking at the equilibrium equations. Lowerin the temperature system results in a higher [CaEt ]e. At lower temperatures, the CO2 fugacity diminishes, thus the resulting carbonate solution is more reactive and dissolves more calcite grains. 3 CONCLUSIONS The calcite dissolution experiments performed under the parameters depicted in this work exhibited a calcium concentration profile adjusted by a function of the form: [Ca"] = [Ca2']e(l-Be-'''),which coincides with the solution for the one-dimension reactivetransport model for porous calcite proposed by Baumann et aL(1985).

382

The adsorption layer model was used to interpret the effects of the experimental parameters on the calcite dissolution process. The dynamics in the adsorption layer determine the dissolution rate; changes in the concentration of the chemical species affect the thickness of this layer, so does the geometry of the porous media. The temperature and PC02 effects are correlated. The latter affects the adsorption layer thickness and limits the extension of the calcite dissolution process. The effect of the fluid flow rate on the dissolution process may be explained in terms of the removal rate of the species present in the adsorption layer. However, this effect depends on the porous media grain size, permeability, and the PC02. The dissolution experiments allow a first approach towards the understanding of what may be occurring in a CaC03-C02-H20 system located inside the reservoir rock. Further work has to be done to study systems closer to real conditions; for instance, the injection and formation water composition may be reproduced in some experiments to investigate its impact on the calcite dissolution. ACKNOWLEDGEMENTS The authors thank PDVSA-Intevep for permission to publish this work. REFERENCES Baumann, J., Buhmann D., Dreydrodt, W. & Schulz H. D. 1985. Calcite dissolution kinetics in porous media. Chemical Geology 53: 219-228. Buhmann, D. & Dreydrodt, W. 1985a. The kinetics of calcite dissolution and precipitation in geologically relevant situations of karst areas 1. Open system. Chemical Geology, 48: 189-211. Buhmann, D. & Dreydrodt, W. 1985b. The kinetics of calcite dissolution and precipitation in geologically relevant situations of karst areas 2. Closed system. Chemical Geology 53: 109-124. Domenico, P. & Schwartz, F. 1998. Physical and Chemical Hidrogeology. 2nd ed. Wiley & Sons, N y . Dreybrodt, W. 1980. Deposition of calcite from thin films of natural calcareous solutions and the growth of speleothems. Chemical Geology 29: 89-105. Palciauskas, V. & Domenico, P. 1976. Solution chemistry, mass transfer, and the approach to chemical equilibrium in porous carbonate rocks and sediments. Geological Society of America Bulletin 87: 207-214. Plummer, L. N., Parkhurst, D. L. & Wigley, T. M. L. 1979. Critical review of the kinetics of calcite dissolution and precipitation. In Chemical Modeling in Aqueous System; Amer. Chem. Symp. Ser. 93: 537-573. SW. & Morgan, J. 1981. Aquatic Chemistiy. 2nd ed. Wiley & Sons, NY.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Composition of charnockite weathering products in three climatic zones W.1.S .Fernando, & R K i t a g a w a Dept. of Earth and Planetary System Science, Hiroshima University,Higashi Hiroshima 739-8526, Japan

B .P.Roser Dept. of Geoscience, Shimane University,Matsue 690-8504, Japan

Y.Hayasaka, & Y.Takahashi Dept. of Earth and Planetary System Science, Hiroshima University,Higashi Hiroshima 739-8526, Japan

ABSTRACT: Weathering products developed on Sri Lankan charnockitic basement rocks in three differing climatic zones (dry, intermediate and wet) have been examined to test the effect of precipitation rate on chemical weathering of a common parent material. Clay fraction mineralogy in the dry zone profile shows typical biotite/vermiculite- vermiculite- smectited h$lloysite- kaolin$ transitions. Smectite is absent from the intermediate and wet zones, which show 10A halloysite- 7A halloysite- kaolinite and vermiculite4 halloysite- kaolinite- gibbsite sequences, respectively. A-CN-K diagrams show a clearly differing weathering trend in the dry zone (C.I.A. < 70) compared with the wet and intermediate zones (C.I.A. 75 to 98). In all three zones, parental REE patterns are essentially inherited intact, even though REE concentrations are moderately depleted in the wet zone products. Although the results overall suggest that precipitation is the primary control on weathering, the XRD mineralogy clearly shows that reaction pathways differ in each climatic zone. mm/yr), intermediate (1250-3000 mm/yr) and wet ( > 3000 mm/yr) (Fig. 1).Rain falls on over 50% of days in the wet zone (up to 70-75% in the wettest areas). In the dry zone rain days are < 25%. Mean annual temperatures in the study areas are 22.5-27.5"C7 but diurnal range is higher than the annual variation. Climatically, the wet zone is classified as humid tropical, and the dry zone is seasonally dry tropical (Panabokke 1996).

1 INTRODUCTION The chemical weathering of parent rocks is affected by a complex suite of factors. Many studies of tropical weathering processes (e.g. Curtis 1990, Johnsson et al. 1993) have reported a good correlation between precipitation and clay composition. Other work (e.g. Nesbitt 1979, Nesbitt et al. 1980) has examined mobility of major elements and REE during weathering. However, it remains unclear whether bedrocks weather directly to different products in different rainfall regimes, or if chemical weathering proceeds at different rates but along the same pathway in differing rainfall regimes, as suggested by Carroll & Hathaway (1963). No weathering-related studies of Sri Lankan rocks have yet been reported. We here examine the weathering products of charnockitic rocks (orthopyroxene-bearing granitoids) in three different climatic zones. The aims of our study are to clarify the relationship of weathering products with precipitation and to investigate major element and REE characteristics of weathering profiles in each zone.

3 SAMPLESITES

Weathering products and soils have accumulated on charnockitic basement in all three climatic zones. Weathering products and parents were collected from sites in each zone (Fig. 1). Samples were collected stratigraphically from the dry and intermediate zones, and along lines away from basement and corestones in the wet zone. At all sites fresh charnockitic basement or corestones display only thin (10-20 cm) leached margins, and transitions to overlying weathering products are abrupt. The ages of the profiles are not constrained. At the dry zone site, a thin (2 m) soil horizon rests directly on charnockite; retention of the texture of a pegmatitic vein from the basement through the soil clearly shows the weathering products are in place. The intermediate zone locality profile is thicker (8.5 m) and parent basement is not exposed. However, retention in outcrop of gently inclined foliation, perpendicular joint sets, and

2 CLIMATE Due to a combination of monsoonal and orographic effects, annual rainfall in Sri Lanka varies from 25005000t mm in the southwest, while the northwest and southeast receive < 1250 mm (Fig. 1).The country is divisible into three main climatic zones: dry ( < 1250

383

oriented corestones show this profile is also in place. Field relations at the wet zone site are less certain. Kaolinitic products fill depressions and joints in the basement, but large irregular corestones are also present. There is no clear evidence for transportation, yet the most altered material is adjacent to parent rock, suggesting that groundwater circulation has occurred along fractures in the basement (Calvert et al. 1980).

parent minerals, and feldspar grains adjacent to opaques are stained, presumably as a result of Fe leaching from the latter. 5 RESULTS AND DISCUSSION 5.1 X-ruy diffruction (XRD)dutu Conventional XRD analysis of air-dried < 2 pm fractions followed by heat and chemical treatment was carried out to identify the weathering products present. XRD patterns of representative samples from the three climatic zones are shown in Fig. 2.

Figure 2. XRD patterns of representative samples subjected to various treatments. 1. 500°C; 2. 350°C; 3.150"C; 4. Formamide solvated; 5. Mg-saturated, glycerol solvated; 6. Glycerol solvated; 7. Mgsaturated; 8. Ethylene glycol solvated; 9. Air-dried.

Figure 1. Sample locations, mean annual rainfall, and distribution of the dry, wet and intermediate zones.

4 PARENT CHARNOCKITE

All parent charnockites are massive or only weakly foliated at outcrop scale, and exhibit medium-grained gneissic or weakly foliated textures in thin-section. Plagioclase, quartz, K-feldspar, pyroxene, biotite and hornblende are the main constituents, and apatite, zircon, magnetite and ilmenite are accessories. The mafic minerals are homogeneously distributed. At the initial stage of weathering, pyroxene and hornblende grains show highly weathered features. Adjacent plagioclase and K-feldspar are cloudy, but no weathering products are obvious. Our microscope studies show pyroxene is most prone to weathering in all climatic zones. Weathering is initiated along grain boundaries, micro-fractures and cleavages of all 384

Diffraction patterns of air-dried samples from the wet zone (Fig. 2) show that vermiculite, mica, kaolinite, and halloysite were present. Chemical and heat treatment indicat5d that chlorite and smegtite were absent. Peaks at 4.84A (18.33" 20) and 4.18A (21.25" 20) were not affected by chemical treatment, but were collapsed by heating at 500"C, showing that the nonphyllosilicates gibbsite and geothite were also present. Intermediate zonc samples (Fig. 2) cont@n peaks for kaolinite, and 7A (12.5" 20) and 10A (8.85" 20) halloysite. No other weathering products were detected. Dry zone patteqs (Fig. 2) contained a prominent peak at 14.28A (6.19" 20), which is expected for smectite, vermiculite and chlorite minerals. Chemical and heat treatment shoyed that chlorite was not present, and that a peak at 12A (7.16" 26) in the oair-dried state represeated interstratified biotite (10A) and vermiculite (14A). Kaolinite and halloysite were also identified in air-dried patterns. 5.2 Proportions of Clay minerals Relative proportions of clay minerals in individual samples were estimated from (001) peak areas. Halloysitic and kaolinitic phases are unevenly

distributed at the sites examined. Although the proportions are only semi-quantitative, several correlations are evident. Most significant is a marked variation in clay mineralogy with respect to climatic zone. Smectite group minerals, which are structurally complex three layer clays rich in soluble cations, are developed only in the dry zone, and constitute 50% of the clay fraction at the base of the section; vermiculiteibiotite forms the remainder. Contents of both decrease upward to near zero over 1.5 m as kaolinite bFcomes abundant (84% in the uppermost sample); 7A halloysite is present in variable amounts, and may form up to 34%. Most of the intermediate zone profile consists of relatively constant proportionso of the 1:l layer miner$ kaolinite (-30%), 7A halloysitc (30%)0 and 10A halloysite (40%). In the upper meter 10A halloysite falls to < 5% while both others increase. In the dry and intermediate zones, the upper parts of the profiles tend to be enriched in cation-poor phases and exhibit increasing kaolinite/halloysite ratios. Wet zone weathering products are extensively leached and cation;depleted phases such as kaolinite (20-67%) and 7A halloysite (17-60%) dominate; vermiculite (5-50%) and vermiculite/biotite (1-10%) are significant in some samples. Distribution of the species is irregular, with highest contents of kaolinite generally adjacent to basement or corestones, suggestive of groundwater leaching. Although precise estimation of relative percentages of gibbsite and goethite is @practical, the high intensity of the peak near 4.82A (18.4'20) indicates that gibbsite is relatively abundant in the < 2 pm fraction in this zone.

alkali and alkaline earth elements are strongly depleted relative to parent in the intermediate and wet zones. Depletion is not marked in dry zone samples. Al,O,-(CaO*+NazO)-K,O (A-CN-K) plots (Fig. 3) show differing weathering trends in each zone. The pH of wet zone soils is relatively low ( < 4). Abundant acid supply caused by the high rainfall creates an environment in which plagioclase and mafic minerals (pyroxene, hornblende, biotite) weather faster than Kfeldspar (Brady & Walther 1989; White & Brantley 1995), producing cation-depleted weathering products (kaolin and gibbsite). This initially produces a trend toward the A-K edge, followed by a trend toward the A apex, as primary phases are degraded and weathering products are produced (Fig. 3). A-CN-K trend in the intermediate zone may also be similar, but the pattern is not clear due to the uniformly high C.1.A ratios there. In contrast, dry zone products lie near the A-CN edge and trend toward the A apex parallel to that join. Because soils in this low rainfall regime have higher pH ( > 6) values, plagioclase and K-feldspars weather at a slower rate and weathering products are mainly smectite (rich in Ca and Na). However, slightly weathered samples in the dry zone are depleted in K,O relative to the parent, possibly

5.3 Chemical charges with weatheririg Whole rock samples were analyzed for major elements by XRF, and REE abundances in selected samples were determined by ICP-MS. Analyses of eight charnockites show the dry and intermediate zone parents are silica-poor (56.6-65.6 wt%) compared to those at the wet zone site (69.7-72.9 wt%). Intensity of weathering can be measured by the Chemical Index of Alteration (C.I.A.) of Nesbitt and Young (1984). Primary minerals generally have CIA ratios of 50 or less, whereas secondary products have higher ratios (smectite 70-85; kaolinite, halloysitc, gibbsite 100). C.1.A ratios near 100 in the wet zone (kaolinite and gibbsite-rich); highest values occur adjacent to basement or corestones. Ratios are variable but uniformly high (89-98) throughout the h&opsite and kaolinite-rich inteiTnecliate zone profile. In contrast, maximum C.1.A in the dry zone is < 70, due to the abundance of mainly smectitic clays and presence of remnant feldspar. Ratios increase quite smoothly from 50 at the base of the section to 70 at the top. Whole rock major element data show all

Figure 3. Molar A-CN-K diagrams (Nesbitt & Young 1984) for parents and products in each zone.

due to rapid degradation of biotite to vermiculite, as suggested by presence of interstratified vermiculite and biotite in the lower part of the profile. REE analyses of six charnockites (Fig. 4A) show contrasting patterns between the parent in the wet zone compared to the dry and intermediate zones. In

385

keeping with the contrasts in major elements, the wet zone parent REE patterns are strongly kdctionated with marked LREE enrichment and significant negative Eu anomalies. In the other two zones patterns are flatter and Eu anomalies are almost absent. Whole-rock REE analyses of 11 weathering products were made. Average REE abundances normalised against respective average parents (Fig. 4B) show that the REE are slightly enriched overall in the dry zone products, but that no fractionation occurs. Average abundances in the intermediate zone are similar to parent or are only slightly depleted. Significant negative Ce anomalies occur in all four samples analysed, suggesting some Ce has been lost. However, the intermediate zone parents also exhibit slight negative Ce anomalies, so protolith heterogeneity may also be a factor. The wet zone average plotted excludes one anomalous sample which has a positive Eu anomaly. This sample aside, the wet zone average is strongly depleted (40-50%) relative to parent (Fig. 4B). Apart from slight relative depletion in the HREE, little fractionation occurs.

Figure 4. REE by zone, (A) Chondritenormalized patterns for parental charnockites; (B) average abundances in products in each zone compared to average parents.

In all three zones REE patterns are thus essentially inherited intact. The decreasing average abundances between zones most likely simply reflect increased leaching related to precipitation.

6 CONCLUSIONS

Weathering products developed on charnockitic basement rocks in Sri Lanka have different compositions in each climatic zone. Although the products in the wet and intermediate zone have chemically similar parents, clay mineral assemblages and CIA ratios differ markedly. Intermediate zone products have mineralogy and CIA ratios more similar to the wet zone, despite a more felsic parent in the latter. This suggests that annual precipitation is here the major control on weathering, and thus the transition sequences of clay minerals are dependent on climate. REFERENCES Brady, P. V. & J. V. Walther 1989. Controls on silicate dissolution rates in neutral and basic pH solutions at 25°C. Geochim. Cosmochim. Acta 53: 2823-2830. Calvert, C. S., S. W. Buol & S. B. Weed 1980. Mineralogical characteristics and transformations of a vertical rocksaprolite-soil sequence i n the North Carolina Piedmont: I. Profile morphology, chemical composition, and mineralogy. Soil Sci. Soc. Amer. 44: 1096-1103. Carroll, D. & J. C. Hathaway 1963. Mineralogy of selected soils from Guam: U.S.G.S. Prof Pap. 403-F, 51 p. Curtis, C. D. 1990. Aspects of climatic influence on the clay mineralogy and geochemistry of soils, paleosols and clastic sedimentary rocks. .Jour. Geol. Soc. Lordon 147: 35 1-357. Johnsson, M. J., S. D. Ellen & M. A. McKittrick 1993. Intensity and duration of chemical weathering: An example from soil clays of the southeastern Koolau Mountain, Oahu, Hawaii. Geol. Soc. Amer. Spec. Paper 284: 147-169. Nesbitt, H. W. 1979. Mobility and fractionation of rare earth elements during weathering of a granodiorite. Nature 279: 206-210. Nesbitt, H. W., G. Markovics & R. C. Price 1980. Chemical processes affecting alkalis and alkaline earths during continental weathering. Geochim. Cosmochim. Acta 44: 1659-1666. Nesbitt, H. W. & G. M. Young 1984. Prediction of some weathering trends of plutonic and volcanic rocks based upon thennodynamic and kinetic considerations. Geochim. Cosmochim. Acta 48: 1523-1534. Panabokke, C. R. 1996. Soils md Agro-Ecological Environments of Sri Lanka. Nat. Res., Energy Sci. Authority of Sri Lanka Pub. 220pp. White, A. F. & S. L. Brantley (eds) 1995. Chemical weathering rates of silicate minerals. Mineral. Soc. Amer. Reviews in Mineralogy 31: 583pp.

Wafer-Rock Interaction 2007, Cidu (ed.), 02001 Swefs & Zeitlinger, Lisse, lSBN 90 2651 824 2

Basilica da Estrela stone decay: the role of rain-water C .A .M.Figueiredo, A .A.Mauricio & L.Aires-Baffos Laboratory of Mineralogy and Petrology, IST, LAMPIST, AV.Rovisco Pais 1049-001,Lisboa, Portugal

ABSTRACT: Chemical analyses of rain and seepage waters were performed on samples collected outside and inside of Basilica da Estrela, Lisbon (Portugal). Located in a moderately air polluted area about 15 km far from the sea, it is the most relevant 18' century monument in Lisbon built with limestones. Phase diagrams, scatter and multivariate statistical analysis of possible hydrogeochemical significance have been attempted. Using a computational hydrogeochemical model (HIDSPEC), saturation indexes (S.I.) of some minerals commonly involved in the weathering process of carbonate stone monuments were calculated. Relative to rainwater, seepage water S.I. values increase for most of the minerals, but saturation and supersaturation is reached only with respect to calcite. Further interpretation of results allowed the assessment of water typology and the origin of chemical species analysed as well the discussion on the possibility of occurrence of potentially stone damaging minerals other then calcite (gypsum, for instance) in the church. 1 INTRODUCTION

2 METHODOLOGY

Periodic chemical analyses were carried out on both rain and seepage waters in order to establish the relationship between water composition and stone decay processes observed inside of Basilica da Estrela. Basilica da Estrela, the most relevant 18' century monument in Lisbon, was started in 1779 and finished eleven years later. It is located in a moderately air polluted area about 15 Km far from the sea. It was built with Jurassic and Cretaceous calcareous limestones exploited at Lisbon region. The interior is all covered with beige, greyish-blue, rose and ochre limestones. These are essentially pure and calcitic limestones with more than 95% of calcium carbonate and less than 3% of silica. The yellowish variety is slightly dolomitic and clayey. With effective porosity less than 1% and permeability ranging from 1.34 x 10-' (mD) to 4.96 x 10-I (mD), these limestones have very little porosity and are practically impermeable materials. Physical weathering forms (granular disintegration, flakes, scales and spalling) prevail inside of Basilica da Estrela. Chemical weathering forms (salt efflorescences and calcitic concretion), largely dominated by calcite re-precipitation forming large white zones are, however, also present. Soluble salts, including gypsum, commonly involved in the weathering process of carbonate stone monuments were, in contrast, practically non existent.

Chemical analyses were performed on rain and seepage waters collected during three years on a weekly basis, outside (terrace) and inside (high choir) of Basilica da Estrela. Cl-, NOi, SO?-, HC03-, CO?-, Na', K', Ca2', Mg2+,W ' w e r e the main solutes determined. Silica, given as ,5302, pH, electrical conductivity (0)and temperature (T) were also measured. Scatter diagrams, multivariate statistical analysis of possible hydrogeochemical significance were attempted. Saturation indexes (S.I.) calculation with respect to some minerals commonly involved in the weathering process of carbonate stone monuments was also made using the HIDSPEC (Carvalho & Aimeida 1989) computational program. This is a hydrogeochemical model based on physical and chemical analyses of waters. It estimates the activity of 68 aqueous species and the S.I. of 55 minerals. To study calcite and gypsum deposition conditions, a phase diagram was also computed. 3 DATA ANALYSIS, INTERPRETATION AND DISCUSSION 3.1 Multivariate data analysis results Principal component analysis (PCA) approach was used to help data interpretation. Factor loadings 2D diagrams show the correlation among the original variables themselves and also between theses and the 387

factor axes (Figs 1-3). On the other hand, projecting the factor scores (samples co-ordinates) onto the first two principal axes, some significant insight into the inter-samples relationships in the data set could also be obtained (Fig. 4). Ca2', SO?-, Mg2", Na' and C1- ions play a significant role in the characterisation of rainwater (Fig. 1).

alteration process involving SO?-. All these variables do not provide, however, enough discrimination of seepage waters to provide sub-classifications other than richedpoor samples in the content of these variables. This could reflect a significant uniformity contribution of ion sources and stone alteration processes.

Figure 1. Plot of loadings on the first two factors axes of rainwater analysis data.

Figure 2. Plot of loadings on the first two factors axes of seepage water analysis data.

Ca2' and SO?' are strongly and positively correlated with each other, as well are Mg2', Na' and Cl-. These variables forming two uncorrelated clusters reflecting the operation of different processes, sources or sinks for these ions. Na' and CI- seein to be essentially derived from marine s ray while the relationship between Ca2' and SO4!indicates the action of other sources (air pollution influence and/or gypsum particles from the atmosphere). Rainwater composition can be explained mainly by a factor composed of marine elements and also by a less important second factor composed of terrestrial elements. Rainwater composition reflects therefore the combination of the sea and anthropogenic derived elements. The other variables, temperature (T), HC03-, pH, K' and N03-, involving either the third, fourth and fifth principal axes are thus less important to explain the overall variance between rainwater composition. K', Na', Cl-, SO?-, HC03-, N03- and pH are variables playing a significant role in characterising seepage waters composition (Fig. 2). They explain most of the variation observed in the chemical composition of the samples, while the other variables including Ca2' as well, are not. This surprising secondary role of Ca2' is possibly associated with stalactite formation observed inside of Basilica da Estrela. Only K', Na', Cl-, HC03- and NOS- form one cluster that is, in general, positively correlated with SO?'. This seems to suggest the same source or process involving the strongly and positively correlated variables forming the cluster as well as a not very different source and

The other variables: Ca2', COj2-, Mg2', T and CT involving also the third or fourth principal component seems thus to be less important to explain the overall variability in the seepage water samples. A clear distinction between rain and seepage waters becomes well established from the PCA performed on the whole data set composed simultaneously of rain and seepage water analyses. Rain (A) ) waters are in opposite position in terms of their physical and chemical properties (Figs 3 7 4 ).

Figure 3. Plot of loadings on the first two factors axes of rain and seepage water analysis data.

388

almost completely described by only the following variables: Mg2', K', Na', Cl-, HC03-, pH and CT. It can be concluded that the water-rock interaction and environmentally-induced processes promote essentially the enrichement of seepage waters in K', Na', C1- and HC03-. 3.2 Water geochemical results Rain and seepage water analyses were plotted in the Piper diagram reproduced in Figure 5. Rain (0)waters belong to SO4 - C1 - Ca and Cl - SO4 - Na type, with pH values ranging from 5 to 7. Conductivity values are around 90 pS/cm and total mineralisation lies between 40 and 100 mg/l. Seepage waters ( belong to HC03- - Na type, showing pH values between 7 and 12. Compared to rain waters, seepage waters are much more mineralised as shown by conductivity (about 700 pS/cm) and total mineralisation (between 200 and 3000 mg/l) values. Estimated S.I. values indicate that all rainwater samples are undersaturated with respect to many minerals. For the most important minerals of interest to our study, rainwater shows decreasing S.I. values for: gypsum > calcite (Fig. 6). Concerning seepage waters, S.I. values increase for most of the minerals relative to rainwater, but saturation and supersaturation is reached only with respect to calcite (Fig. 6). Given that rain and seepage waters were collected in an urban environment, it is important to consider the possibility of the transformation of calcite into gy-

Figure 4. Factor scores plot of rain and seepage water analysis data onto the first two factor axes.

Seepage waters are essentially characterised by higher values of K', Na', Cl', pH, CT and lower values of Mg2' than the mean values calculated for all the rainwater and seepage samples. These samples show also the highest HC03' concentrations. Rainwater sam les have as their main characteristic values of MgP' higher and values of K', Na', Cl-, HC03-, pH and CT lower than the mean values calculated for all the samples. Subgroups of rain and seepage waters samples richer in sulphate content than the others could also be clearly established in this plan, on the negative side of the second factor axis. The differences between these two types of waters can be

Figure 5. Piper diagram of rain (0)and seepage (4) water analyses.

389

is consistent with the large white stain zones of calcite precipitation and small stalactites and stalagmites observed inside the church. Furthermore, gypsum precipitation from rain and seepage water evaporation can only be expected in smaller relative quantity. This forecast can be ascertained in hture works for the outside of the church. 4 CONCLUSIONS

Figure 6 . Calcite SI and Gypsum SI for rain and seepage waters.

psum according to the following chemical equation: Ca CO3 (S)+S02'+HzO(~,+2H'tCaS04.2H20(,)+ CO2 (g)

Using standard free energy Of formation data (MagalhHes et al. 1997) and the methodology presented the between and gypsum Obtained from the equation ( l ) 7 at 298.15 K is given by: 2 pH + log (acoz) = 14.20 + log (aso.?-)

1. Ca2', SO:-, Mg2', Na' and CI- ions play a significant role in characterising rainwaters. Rain waters belonging to SO4 - C1 - Ca and Cl - SO4 - Na type, reflects therefore the combination of sea and anthropogenic derived elements. 2. Seepage waters belong to HC03 - Na type. K', Na', Cl-, SO:-, HC03-, N03' and pH are variables playing a significant role in their characterisation and explain most of the variance observed, 3. The positive and, in general, strong correlation observed between Cl-, K', Na', HC03-, Nos- and SO:, reflects a significant uniformity contribution of ion sources and water-rock interaction processes, promoting essentially the enrichement of seepage waters in K+, Na+, cl-and HC03-, 4, The unlikely Occurrence of gypsum should not be surprising from this type of seepage waters, reinforced by gypsum-calcite equilibrium diagram interpretation

(2) ACKNOWLEDGEMENTS

where (ax) = activity of chemical species x.

This study was partially financed by PRAXIS/P/ ECW 13 0 12/1998 and PRAXIS/P/CTE/ 1 1003/ 1998. REFEENCES Carvalho, M.R. & Almeida, C. 1989. HIDSPEC, um programa de especiado e calculo de equilibrios agudrocha. Geocitncius, Rev. Univ. Aveiro, 4(2): 1-22. Magalhiies, M.C.F., Ares-Barros, L. & Aives, L.M. 1997. Thermodynamics of carbonates and sulphates. Applications to stone decay studies - the case of "Mosteiro dos Jeronimos, Lisboa". Geocitncius, Rev. Univ. Aveiro, 11(12):139-147.

Figure 7. Rain and seepage water composition plotted onto calcite and gypsum stability fields.

The stability field diagram for the equilibrium between gypsum and calcite is shown in the Figure 7, where data concerning rain and seepage water samples chemical composition are also plotted. The plot 390

Water-RockInteraction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

A new model of rock weathering: design and validation on a small granitic catchment L.FranGoisl, A.Probst2, Y.Goddkris1, J.Schott2, D.Rasse1, D.Viville3, 0.Pokrovsky2, &

B .DuprC2 'Lahoratoire de Physique Atmosphkrique et Planktaire, Universitk de Lidge, Lidge, Belgium 2Laboratoire des Me'canismes Transfert en Gkologie, Universite'Paul Sabatier, CNRS UMR 5563, Toulouse, France Scentre d 'Etudes et de Recherches Eco-Gkoographiques,CEREGKJLPICNRS,3 rue del 'Argonne, 67083 Strasbourg Cedex, France

ABSTRACT: A new process-oriented model of rock weathering and soil chemistry is presented. As a first validation test, this model is applied to a small granitic catchment, the Strengbach watershed located in the Vosges mountains, France. A long record of meteorological, hydrological and chemical data is available for this site. The vegetation is composed of spruce over two-thirds of the watershed and a mixed beech and fir forest over the remaining third. Daily soil CO:! fluxes and water runoff are calculated with the ASPECTS model (Rasse et al. 2000). The weathering model can calculate the budget of the major ions in the soil solution in a fully transient mode from rain composition data and dissolution kinetics for a wide set of primary minerals. In the experiments performed over the Strengbach watershed, the model is run to steady state. Preliminary results of this experiment are compared to some available data. 1 INTRODUCTION Chemical weathering of continental silicates is a key factor regulating fluctuations of atmospheric CO2 concentrations and associated climatic changes at geological timescales (>106 years). At shorter timescales, chemical weathering of the bedrock is responsible for soil genesis, thus contributing to the nutrition of the land biosphere. Finally, chemical weathering of the bedrock drives the chemistry of streams and rivers. Since river water is ultimately delivered to the world ocean, weathering is the original provider of ocean alkalinity and nutrients, which drives ocean chemical and biological evolutions. Despite these fundamental roles of weathering in the global Earth's system, today there is no process-oriented model describing weathering at large spatial scales. However, to understand rock weathering impacts on the chemistry of large rivers and the transfer of chemical elements from continents to oceans, it is necessary to develop process-oriented models of weathering which can be scaled up from small monolithologic catchments to large river basins or even larger scales. Here, we combine a model of the water and carbon cycles in forest vegetation and soils (ASPECTS) with a model of rock weathering and soil chemistry. At this stage this system is not coupled in the sense that information is transferred only one-way from the water/carbon module to the weathering module. ASPECTS calculates soil water drainage and soil CO:! production. The drainage output constrains the

39I

temporal evolution of the water flow in the weathering model, while soil CO2 production is used to estimate soil PCO2. As a validation test, we apply this chain of models to a stream watershed.

2 THE ASPECTS MODEL ASPECTS (Atmosphere-Soil-Plant Exchanges of Carbon in Temperate Sylvae, Rasse et al. 2000) is a fully coupled scheme of the water and carbon cycles in the vegetation and soils of temperate forest ecosystems. All fluxes and pools are calculated with a time step of 30 minutes. Photosynthesis and transpiration are calculated separately for shaded and sunlit leaves (De Pury & Farquhar 1997). The stomata1 conductances of CO2 and H20 are related to the net assimilation of the leaves (Leuning 1995). Photosynthetic assimilates transit through a carbohydrate pool before being allocated to other plant reservoirs (leaf, starch, branches, stem, coarse roots and fine roots). This allocation varies with stress factors (water, nutrient, temperature, etc) influencing the development of forest vegetation. The soil is divided into 5 layers. Temperature and water amount is calculated in each soil layer, as well as coarse and fine root biomass, litter and soil carbon. The soil CO2 production used to estimate soil PC02 in the weathering model is the combined respiration flux of roots and soil microorganisms integrated over the soil profile. The drainage used in the weathering model is the drainage at the bottom of the deepest soil layer.

nevertheless important rainfall occurs in spring whereas the driest season is in autumn. Usually, snowfall season lasts four months per year from December to April. The total runoff reaches 853 mm for the 1986-1995 period (Probst et al. 1995), which corresponds to a mean annual discharge of 21.7 Us. The bedrock is a base poor leucogranite (The Brtzouard granite) aged of 3 15 k 7 Ma. This granite is coarse grained and has undergone hydrothermal alteration, which is particularly obvious on the south-facing slope of the catchment (El Gh'Mari 1995). At the upper margin of the catchment, a banded gneiss lies in contact with the granite. The soils are rather deep (80 cm average), sandy and stony and lie on a saprolite, which can reach 10 m depth in places. These soils belong to the brown acidic to ochreous podzolic soils serie. A small saturated area with permanent moisture conditions, which only represents 2% of the total catchment area (Probst et al. 1990), takes up the valley bottom near the outlet. This area can contribute significantly to stream water output particularly during storm events (Idir et al. 1999). The forest cover (95% of the total area) is complete and homogenous. Norway spruce (Picea abies Karst.) dominates while mixed silver fir (Abies alba Mill.) and beech (Fagus sylvatica L.) only represent one third of the area.

3 THE WEATHERING MODEL The weathering model explicitly calculates the budgets of 7 different chemical elements present in the soil solution or the exchange complex: Ca, Na, K, Mg, Si, S (SO?-) and C1. At every time step (1 day), the concentration of 30 ionic species involving these elements, as well as AI and C, are calculated from 29 different chemical equilibrium reactions and 4 Gaines-Thomas equations governing cation exchange. The concentration of AI3+ is derived by selecting the lowest concentration inferred from equilibrium conditions with kaolinite or gibbsite. Inorganic C species (H2CO3, HC03', C032-) are assumed in equilibrium with soil PC02. Weathering is assumed to occur below the root zone. A elemental budget is however performed in the root zone which has the consequence of concentrating the rain solution reaching the soil. Below the root zone, the solution is chemically altered through the dissolution of primary minerals and the possible precipitation of secondary minerals. Currently, a set of 16 primary primary minerals is taken into account in the model, allowing its applicability to soil underlaid with various bedrocks. The dissolution rate R, of each primary mineral is described as:

4.2 Site equipmerit and sampling where the rate rH(resp. roH) at low (resf. high) pH is a power function of the activity of H (resp. OH-), rH20 is the rate at neutral pH and rL is the rate representing organic ligand promoted dissolution. Whenever relevant, the inhibition by AI andor alkaline ions is taken into account in these rate laws. The factor ( I - Q k ) is a chemical affinity factor which describes the reduction of the dissolution rate when the equilibrium with the primary mineral is approached.

The Strengbach catchment has been investigated since 1986 (see Probst et al. 1990 for further details) and has been previously monitored to study the effects of acid rain on a forested ecosystem and particularly on the hydrochemistry of surface waters as well as on weathering (e.g., Probst et al. 1995). This site was progressively fitted out and many geochemical, mineralogical and biological studies have been performed (Fichter et al. 1998, Idir et al. 1999, Amiotte-Suchet et al. 1999, Probst et al. 2000). As a routine, bulk open field precipitation is regularly collected (every two weeks) at four sites in polypropylene funnel collectors exposed all times. During snow season buckets are used. Throughfall is sampled using 2 m-long open gutters and soil solutions are collected at different depths using zerotension lysimeter plates, both in a beech stand and in an old spruce stand. Four springs emerging 4 m down in the granite at the upper part of the basin are the main contributors to the stream. They flow into a general collector (CR), which is partly harnessed for drinking water supplies (for 2% of the total runoff, Probst et al. 1992). Stream water is controlled by an H-Flume notch weir and water level is monitored both by ultrasonic and mechanical limnigraphs. Stream water and spring water are collected weekly and

4. APPLICATION TO THE STRENGBACH WATERSHED 4.1 Site description The Strengbach forested catchment (80 ha area) is located on the eastern part of the Vosges Mountains (North East of France), 58 km SW from Strasbourg. The elevation ranges from 883 m at the outlet to 1146 m at the catchment divide. The slopes are rather steep. The climate. is temperate oceanic nountainous and westerly wind dominates. The monthly average of daily mean temperature ranges fiom -2°C to 14°C (Probst et al. 1990). The mean annual rainfall is about 1400 mm (Probst & Viville 1997) and rainfall is spread all over the year, 392

stream water is also sampled more frequently during flood events by automatic samplers. Samples are stored in polyethylene bottles and filtered in the laboratory (0.45 pm Millipore membrane). All waters are analysed in the laboratory as follows: pH, conductivity and alkalinity electrometrically (the latter by Gran tit r ation); sodium, potassium, calcium, magnesium by atomic absorption spectrometry; ammonium and silica by colorimetry; alurninum by ICP-AES; nitrate, chloride and sulphate by ion chromatography.

the data. This is presumably due to incorrect initialization conditions, since Oct 94 is actually preceded by Sep 95 in the simulations. Similarly, the

4.3 Design of the model experiments Model experiments were performed with meteorological inputs corresponding to the 1-year period October 1994 to September 1995. The runs of both models were conducted over many years, by repeating this meteorological dataset as many times as needed. In ASPECTS, the model was run over the whole lifetime of the trees, starting from a young population and allowing for some regular tree cuttings to reach approximately the measured value of the stand leaf area index today. The weathering model was run to achieve a steady state (with seasonal changes, but no change from one year to the next). Both models were run at two hypothetical sites corresponding to the old spruce and beech stands. The primary minerals included in the model weathering zone are the same for both sites, but their abundances slightly differ (Tablel). Small amounts of sericite and apatite are also present, but are not neglected in the current model runs. The weigth percent abundance is transformed into volume percentage. For the sake of simplicity, the primary mineral areas are assumed proportional to these volume percentages. The total primary mineral area is a model parameter. Tablel. Primary minerals over old spruce and beech stands in the weathering - model (wt %). __l-l__l-

Quartz Orthose Albite (An 6) Muscovite

Old spruce 32 29 21 15

Beech 32 31 19 12

4.4 Hydrological budget In Figure 1, the model-predicted daily runoff is compared to the runoff derived from the measured stream discharge. The overall shape and amplitude of the model runoff curve is quite satisfactory, although substantial discrepancies with the data occur. For instance, the model produces a peak at the beginning of the simulation (Oct 94) not observed in

Figure 1. Daily runoff (surface runoff + deep drainage) predicted by ASPECTS for the beech and the old spruce stands over the period Oct 1994 to Sep 1995 and compared to runoff values derived from the stream discharge measured at various dates. Negative day numbers refer to the last three months of 1994.

large runoff values measured at the end of January 1995 are not correctly reproduced by the model. Since these high values are associated with snow melt events, this discrepancy suggests that the model parameterization of snow melting may be too simple. 4.5. Weathering and soil solution Table 2 lists soil solution pH and concentration data measured in the Strengbach catchment soils during September 1992 (Probst et al. 2000). The model results for the same month of 1995 can be compared with these data. Although these results are still very preliminary and do not correspond to the same year as the data, the same overall trend is observed in model-predicted concentrations and measurements. However, the concentration of A13+ is highly underestimated by the model. This results is partly linked to a significant overestimation of H4Si04 concentration, but it also indicates that the equilibrium hypothesis of the soil solution with kaolinite (from which A13+ is estimated) is not adequate. A major improvement of such weathering models would consist in making an explicit budget of aluminium in the soil solution and the exchange complex.

5 CONCLUSIONS In this paper, the combined use of models of the waterkarbon cycles and of soil chemistry and

393

Table 2. Comparison of soil solution pH and some concentrations (pmol 1.') predicted by the model with measurements performed at the site in September 1992 (Probst et al. 2000). Measured Model Model 30-70 cm Spruce Beech Sep 1994 Sep 95 Sep 95 4.21 4.22 PH 4.29-5 .O 1 ~ 1 3 + 28-45 2 2 Ca2+ 20-40 11 12 6-8 4 4 Mg2+ Na' 41-61 39 32 K+ 27-1 19 21 19 177 165 H4SiO4 41-71 S04255-62 45 40 c124-98 35 30

weathering has been illustrated through an application to the Strengbach catchment in the Vosges mountains, France. Although the model experiments performed are still preliminary, the methodology looks promising to get a synthetic view of terrestrial ecosystems in a single modelling framework. The model still needs refinements, as well as a more detailed validation on several catchments with different vegetation types and lithologies. A future important step will be the full (two-way) coupling of the vegetation module with the weathering model. Indeed, it is essential to take vegetation uptake into account to deepen our understanding of the dynamics of soil solution chemistry.

REFERENCES Amiotte-Suchet, P., D. Aubert, J.L. Probst, F. Gauthier-Lafaye, A. Probst, F. Andreux & D. Viville 1999. 6I3C pattern of dissolved inorganic carbon in a small granitic catchment: the Strengbach case study (Vosges Mountains, France). Chem. Ceol. 159:129-145. De Pury, D.G.G. & G.D. Farquhar 1997. Simple scaling of photosynthesis from leaves to canopies without errors of bigleaf models. Plant, Cell and Environment 20537-557. El Gh'Mari, A. 1995. Etude mine'ralogique, pe'trophysique et ge'ochimique de la dynamique d 'alte'ration d'un granite soumis au de'pbts atmosphe'riquesacides (Bassin versant du Strengbach, Vosges, France) : m'canismes, bilans et mode'lisations. Thkse de doctorat, UniversitC Louis Pasteur, Strasbourg, 202 p. Fichter, J., M.P. Turpault, E. Dambrine & J. Ranger 1998. Mineral evolution of acid forest soils in the Strengbach catchment (Vosges Mountains, N-E France). Ceodermu 82:315-340. Idir, S., A. Probst, D. Viville & J.L. Probst 1999. Contribution des surfaces saturtes et des versants aux flux d'eau et d'C1tments exportts en pCriode de crue: traqage B l'aide du carbone organique dissous et de la silice. Cas du petit bassin versant du Strengbach (Vosges, France). C.R.A.S. 328:89-96. Leuning, R. 1995. A critical appraisal of a combined stomatalphotosynthesis model for C3 plants. Plant, Cell and Environment 18:339-335. Probst, A., E. Dambrine, D. Viville & B. Fritz 1990. Influence

394

of acid atmospheric inputs on surface water chemistry and mineral fluxes in a declining spruce stand within a small granitic catchment (Vosges massif, France). J. Hydrol. 116:lOl-124. Probst, A., D. Viville, B. Fritz, B. Ambroise & E. Dambrine 1992. Hydrochemical budgets of a small forested granitic catchment exposed to acid deposition: The Strengbach catchment case study (Vosges massif, France). Wat. Air and Soil Poll. 62:337-347. Probst, A., B. Fritz & D. Viville 1995. Mid-term trends in acid precipitation, streamwater chemistry and elements budgets in the Strengbach catchment (Vosges mountains, France). Wat. Air and Soil Poll. 79139-59. Probst, A. & D. Viville 1997. Bilan hydrogtochimique du petit bassin versant forestier du Strengbach h Aubure (Haut-Rhin). In: Rapport scientifique activitds de recherche, 5e rCunion du conseil de direction scientifique IfareIDFIU, Conseil de l'Europe, 3010411998, 59-66. Probst, A., A. El Gh'Mari, D. Aubert, B. Fritz & R. McNutt 2000. Strontium as a tracer of weathering processes in a silicate catchment polluted by acid atmospheric inputs, Strengbach, France. Chem. Ceol. 170:1-4. Rasse, D.P., L. Franqois, M. Aubinet, A.S. Kowalski, I. Vande Walle, E. Laitat & J.-C. GCrard. 2000. Modelling short-term CO2 fluxes and long-term tree growth in temperate forests with ASPECTS. Ecological Modelling, in press.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2

Characteristics of srnectites from nickeliferous laterite in Australia A.Gaudin & Y.Noack CEREGE, UMR 6635 CNRS - University of Aix-Marseille I l l , France

A .Decarreau & S .Petit HYDRASA, UMR 6532 CNRS - University of Poitiers, France

ABSTRACT: Lateritic weathering profiles developed on ultrabasic rocks of the Yilgarn craton (Western Australia) are worked as nickel ore (Murrin Murrin). The upper part of the saprolite consists of a thick (up to 30 m) smectite zone in which nickel is concentrated (up to 2.5 % Ni). These smectites were previously described as nontronites. A refined mineralogical and crystallochemical study (XRD, SEM, TEM, FTIR spectroscopy) gives evidence of a complex nature of these smectites. Their layer charge arises mainly from the octahedral sheet and they must be considered as intermediate between four end-members: Al-montmorillonite, Albeidellite, Fe-montmorillonite, and Fe-nontronite. Moreover these uncommon smectites have a large range of A1 for Fe octahedral substitution, more extended than previously reported for natural smectite. Finally, contrary to other lateritic ores (Brazil, New Caledonia) nickel is not clustered in trioctahedral domains or layers of secondary phyllosilicates, but is randomly distributed within the octahedral sheet of dioctahedral smectites. 1 INTRODUCTION Murrin Murrin Ni-laterite deposit is located at 60 Km East of Leonara at north-eastern Yilgarn of Western Australia. The studied lateritic profiles are developed on ultrabasic rocks constitued of peridotites with a network of serpentine which has a mesh texture. A typical laterite profile of the site comprises at the basis a saprolite zone with serpentine, smectite, maghemite and chlorite, overlained by a smectite zone with smectite, maghemite and chlorite, and, in discordance at the top of the profile, a ferruginous zone with goethite, kaolinite and a little amount of smectite. Smectite can occur in replacement of olivine and serpentine in the saprolite zone; in this case its colour is brown. However, smectite is mainly concentrated in the more altered smectitic zone (up to 30 m thick) or in fractures, and here smectite's colour is green. Initial amount of nickel in the fresh rock is low (usually olivine: 0,3-0,5%, and serpentine: 0,2-0,5% Ni), but weathering processes concentrate nickel in secondary phyllosilicates, here, smectites (1-3% Ni). However, this nickel content is low compared to the one of nickeliferous phyllosilicates from weathering profile of New Caledonia (serpentine: 14.55% Ni, Besset 1978) or Brazil (Jacuba pimelite: 36% Ni, Decarreau et al. 1987).

2 MATERIALS AND METHODS

XRD powder were recorded at 0.046"28/min with a PW3710 Philips diffractometer equipped with a Cobalt tube and a graphite monochromator. Scanning electron microscope and transmission electron microscope observations were made using respectively a SEM 515 Philips (CP2M, Marseille) and a TEM Jeol 2000 FX (CRMC2, Marseille). Using these two microscopes, samples were chemically analysed with an energy-dispersive X-ray system, EDX-EDAX for the SEM and EDX-Tracor Northern for the TEM. Fourier Transform Infrared (FTIR) spectra were recorded in 4000-300 cm-' range on a Nicolet 510 FTIR. The disks were prepared by mixing 1 mg sample with 150 mg KBr. The method of smectites layer charge measurement by infrared spectroscopy is based on the quantitative determination of the amount of NH4' in NH4-saturated smectites before and after Li-fixation (Petit et al. 1998). 3 RESULTS 3. I Isofopic data

Isotopic oxygen and deuterium measurements were made on three samples. Their values vary from 20.1 to 21.4 for 6l80 and from -79.4 to -87 for 6D showing that these smectites originate from low temperature weathering (Savin et al. 1998). o/oo

o/oo

395

o/oo

0/00

end-member. At last in the two groups nickel amount is approximately the same (1-3% Ni), and the occurrence of chrome is noted (sometimes up to 3%).

3.3 Transmission electron microscopy Particles morphology indicates that these smectites are composed of irregularly folded and crumpled sheets. The microdiffraction patterns were always pseudo-annular, sometimes slightly punctuated. Such ring patterns are indicative of turbostratic disorder. EDX chemical analyses show that at the particle scale, smectites with intermediate Fe-A1 compositions can still be observed. This supports the solidsolution hypothesis.

Figure 1. SEM, EDX analyses (% atome).

3.4 Infiared spectroscopy

3.2 Scanning electron microscopy Particles less than 2pm were extracted by centrifugation in order to isolate the smectites from the rock. SEM chemical analyses permit to distinguish between "brown smectite" from the saprolite and "green smectite" from the smectitic zone (Fig. 1). The first group is iron-rich, has a relatively homogeneous composition, and is weakly enriched in magnesium compared to the green smectites. This enrichment may be due to the occurrence of magnesian phyllosilicates (serpentine and chlorite) or of Mgsmectite. The second group reveals a constant content of divalent cations Mg-tNi and a large range of Al/Fe ratio (Fig. 1) without relation with depth of weathering. This large range of compositions can be interpreted either as a mechanical mixing of ferruginous and aluminous smectites or as a solid-solution between a ferruginous and an aluminous smectite

396

We note changes of OH bands in relation with the variation of iron and aluminium content in the smectites. Samples with different Fe-AI contents based on SEM, EDX analyses, are presented in the Figures 2 and 3. In the OH-bending region (Fig. 2): four major absorption bands near 920, 870, 820 and 760-790 cm-' are observed (Fig. 2-b), and respectively attributed to All-OH, A1Fe3+0H, Fe?+-OH, and Mg2+Fe3+-OH vibration (Farmer 1974, Stubican & Roy 1961 Goodman et al. 1976, Russel et al. 1970). Therefore with the increase of aluminium contents we observe a disappearance of Fe?+-OH vibrations, an increase in the relative intensity of the AlFe3'-OH band and the appearance of a shoulder due to Alz3+-OHabsorption band. On the other hand, the increase of aluminium is accompanied by a shift of the Si-0 bands at 495 and at 1025 cm-' towards the higher

Table 1. Distribution charge. Zone Smectite type and associated mineralogy Number of samples YOoctahedral charge Total charge (meq/l OOg)

ferruginous Kaolinite and goethite 2 35-3 1 ?

Smectitic and fracture Green smectite ))

((

wavenumber (Fig. 2-a). In the OH-streching region (Fig. 3): A large assymmetrical band varies with the A1-Fe composition. Indeed, the enrichment in aluminium is accompanied by a shift of the maximum and a widening of this band towards the higher wavenumbers. In fact, this large band is due to the overlapping of several components, particularly Fe23+-OH, A1Fe3+-OH and Al?+-OH located res ectively near 3540 cm-’, 3585 cm-’ and at 3620 cm- (Farmer 1974, Madejova et al. 1994, Grauby et al. 1993). Distribution of octahedral cations: the occurence of a

P

((

saprolite Brown smectite )) serpentine

9 45 to 66 78 to 92

7 37 to 59 72 to 85

2 7 1-63 ?

hedral sheet of smectite but not completely random. Trioctahedral bands are not detected in infrared spectra (3Mg-OH, 2MgNi-OH, Ni2Mg-OH or 3NiOH), and only dioctahedral bands are visible. Trioctahedral clusters are not observed here, contrary to previous data from nickeliferous smectites of Jacuba (Decarreau et al. 1987). From these data, nickel appears diluted in the smectite octahedral sheet. To confirm this result, EXAFS studies are in progress. Layer charge estimation using infrared spectroscopy of NH4smectites is presented in the Table 1. The total layer charge of the whole samples are relatively homogeneous and varies from 72 to 92 meq/l OOg. The octahedral charge is always significant and often predominant and related to (Mg2++Ni2+)/(AI3++Fe3+) ratio. Charge distribution seems to have no relation with Al/Fe ratio. Both “brown” and “green” smectites present an octahedral charge up to 66%. Indeed, on 16 samples observed, only 4 present a dominant tetrahedral charge. Therefore these smectites are intermediate between four end-members: Al-montmorillonite, Al-beidellite, Femontmorillonites and Fe-nontronite. Samples extracted from less altered zone, in the saprolite, present higher relative octahedral charge. These smectites, being mixed with serpentine, have a structural formula that cannot be reached. However, the IR spectra presenting a stronger FeMg-OH bending band, we can suppose that the high level of octahedral charge may be linked to higher Mg+Ni octahedral contents. Smectites from the ferruginous zone present the highest tetrahedral charge percent which could correspond to an increase of tetrahedral substitutions, linked to an enrichment in aluminium in the clay structure.

Figure 3. FTIR, OH-streching region. 4 CONCLUSIONS broad AlFe-OH bending indicates that the iron and aluminium are present in the same sheet, there is not segregation of Fe and Al, these elements seem to be distributed randomly in the octahedral sheet. This confirms, at the sheet scale, the Fe-A1 substitution. There is a good concordance between the SEM chemical analyses and IR data which indicate that the chemical variation in Fe-A1 composition of these smectites occurs within neighbouring octahedra. Concerning magnesium, the FeMg-OH bending is the only visible one. The absence of AlMg-OH band, even for the sample with the higher aluminium content, shows that Mg is distributed within the diocta-

Smectites from weathering profiles developed on ultrabasic rocks in Western Australia have been described previously as nontronite. Furthermore, in Nickel-rich TOT clays from similar weathering profiles in South America, New Caledonia, Ni appears strongly clustered in trioctahedral clusters or layers. The studied smectites must be described as strictly dioctahedral and with most often a dominant layer charge occurring within the octahedral sheet. Chemically, they are characterized by a quite constant Mg content and a wide range of Al/Fe ratios. The chemistry of the octahedral sheet evolves from 397

about (Mgo.5Feo.4All.l) (for 4 Si) to about (Mg0.5Fe1.35A10.15).IR data assess that the 3 octahedral cations are associated within the octahedral sheets, without strong clustering. These smectites must be considered as intermediate between four end-members: Al-montmorillonite, Al-beidellite, Femontmorillonite, and Fe-nontronite. Nickel appears randomly dispersed within the octahedral sheet of these octahedral smectites. Lastly, the chemical changes of smectites cannot be related to depth in weathering profile. These new data about smectites would permit a more accurate approach both in geochemistry of balances and in thermodynamic of reactions occurring during weathering of ultrabasic rocks. AKNOWLEDGEMENTS This study has been supported by GDR Metallogeny of CNRS. Tthe authors thank M. Wells and C. Butt (CSIRO, Perth, Australia) for their help in the field and Anaconda Nickel NL for assistance in providing access to collect the samples. REFERENCES Besset, F. 1978. Localisations et repartitions SUC cessives du nickel au cours de l’alteration lateritique des peridotites de Nouvelle-Caledonie. PhD thesis, Univ. Montpellier. Decarreau, A., F. Colin, A. Herbillon, A. Manceau, D. Nahon, H. Paquet, D. Trauth-Badaud & J. J. Trescases 1987. Domain segregation in Ni-Fe-Mg-smectite. Clays and Clay Minerals 35: 1-10. Farmer, V.C. 1974. The infrared spectra of minerals. V.C. Farmer Edition, Mineralogical Society. Goodman, B.A., J.D. Russel, A.D. Fraser & F.W.D. Woodhams 1976. A Mossbauer and IR spectroscopic study of the structure of nontronite. Clays and Clay Minerals 24: 53-59. Grauby, O., S. Petit, A. Decarreau & A. Baronnet 1993. The beidellite-saponite solid-solution: an experimental approach. European Journal of Mineral09 5: 623-635. Madejova, J., P. Komadel & B. Cicel 1994. Infrared study of octahedral site populations in smectites. Clay minerals 29: 3 19-326. Petit, S., D. Righi, J. Madejova & A. Decarreau 1998. Layer charge estimation of smectites using infrared spectroscopy. Clay Minerals 33: 579-591. Russel, J.D., V.C. Farmer &B. Velde 1970. Replacement of OH by OD in layer silicates and identification of vibrations of these groups in infrared spectra. Min. Mag. 37: 869-879. Savin, S.M. & J.C.C. Hsieh 1998. The hydrogen and oxygen isotope geochemistry of pedogenic clay minerals: principles and theoretical background. Geoderma 82: 227-253. Stubican, V. & R. Roy 1961. Isomorphous substitution and infrared spectra of the layer lattice silicates. Amer. Min. 46: 32-5 1.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Surface area vs mass - which is most important during mineral weathering in soils? M .E.Hodson Department of Soil Science, University of Reading, Postgraduate Research Institute for Sedimentology, Whiteknights,PO Box 233, Reading, RG6 6DW, U.K.

ABSTRACT: Mineral dissolution rate is usually normalised to mineral surface area. It might be expected that the dissolution of finer, high-surface area mineral grain size fractions of soils are more important for release of cations than coarser, low-surface area size fractions. However, this overlooks the relatively greater mass of coarser grain size fractions present in many soils. This paper reports batch experiment measurements of dissolution rates of 2000 - 500, 500 - 250 and 250 - 53 pm grain size fractions from a glacially-derived granitic soil after removal of secondary oxide and organic matter. Surface area- and mass-normalised dissolution rates of the different size fractions vary by a factor of 2 - 3 whereas net contribution of the various grain size fractions to total element release from the soil varies by up to an order of magnitude, with the greater mass of the coarsest fraction contributing the greatest amount of cations.

1 INTRODUCTION The release of elements from the dissolution of primary silicate minerals in soil is an important source of nutrients for the soil biota (Marschner 1995). In addition mineral dissolution can neutralise acid rain and affect water quality (White and Brantley 1995). Thus the study of mineral dissolution is of great importance if we are to develop a true understanding of ecosystem functioning. It is often assumed, either implicitly or explicitly, when considering element release from soils due to mineral dissolution, that, due to their greater surface area (expressed in terms of m2/g) the dissolution of finer grained minerals is more important than the dissolution of coarser grained minerals. There are at least two reasons why this may not be so. In nonglacial soils the finer grain size fractions of a soil are dominated by reaction products generated from the weathering of coarser grained material. These reaction products, usually clay minerals, will be less reactive than the primary silicates from which they are derived. In soils with glacially derived parent material, the mineralogy, and thus reactivity, of the finer and coarser grain size fractions of soils will be similar due to the grinding of parent material by glaciers prior to soil formation. The finer grain size fractions will have a greater surface area expressed as surface area per unit mass, but normally the coarser grain size fractions will have a greater mass. Thus the dominance of either the coarse or fine grain

size fractions in element release during weathering will depend on the total amount of surface present in both size fractions (equal to mass multiplied by surface area per unit mass) and the distribution of the surface area between the various minerals in that size fraction. It is the dissolution of different grain size fractions from a soil with a glacially ground parent material that the present set of experiments were carried out to investigate. 2 METHOD 2.1 Soil An iron-humus podzol with a granitic parent

material was collected from Glen Mharcaidh in Scotland. Soil from the B horizon of this soil was selected for the current experiment. The soil was airdried and sieved to a grain size of < 2 mm. Organic material was removed fiom the soil by oxidation using hydrogen peroxide (Bock 1979). Amorphous secondary precipitates were removed using pH 3 acid ammonium oxalate extraction (McKeague & Day 1966). The remaining mineral component of the soil was dry sieved to yield 2000 - 500 pm, 500 250 pm, 250 - 53 pm and < 53 pm size fractions. The < 53 pm size fraction was separated into a 53 - 2 pm and < 2 pm size fraction by sedimentation. The different fractions were weighed and their surface areas were determined by gas adsorption and application of the BET isotherm (Brunauer et al. 1938) using a Coulter Surface Area analyser (Table 1>. 399

Table 3. Size fraction composition (wt%).

Table 1. Size fraction properties. Size fraction

Surface area (m2) in 1 g bulk soil 0.24

Mass (8) used in experiment 1 and 2 2.52, 2.53

0.12

0.08

2.54, 2.72

0.20

0.32

2.08,2.01

0.11

0.68

0.56, 0.55

Surface % of total area grains by mass (m2/g> 0.43 0.56

2000500 pm 5000.69 250 pm 250 1.63 53 pm 53 - 2 pm 6.38

NazO MgO A1203 SiOz p20.5 K20

CaO Ti02 MnO FezO3

Subsamples of the different fractions were ground to a fine powder; their mineralogy was determined by X-ray diffraction following the method of Hooton & Giorgetta (1977) (Table 2) and their elemental composition by X-ray fluorescence using a Philips PW1480 XRF spectrometer and the suppliers X40 software (Table 3).

Mica Chlorite Quartz Potassium feldmar 2 63 16 2000-500 pm 3 2 62 16 500-250pm 3 4 52 12 250-53 pm 5 53-2pm 5 5 55 11

500 - 250 pm 2.8 0.0 10.6 80.4 0.0 5.2 0.2 0.1 0.0 0.7

250 - 53 Pm 3.3 0.3 16.9 70.5 0.2 4.7 0.6 0.5 0.0 3.0

53 - 2 pm 3.1 0.7 16.3 69.5 0.1 4.3 2.1 0.7 0.1 3.1

3 RESULTS Analysis of solution compositions using PHREEQ indicated that all the solutions were undersaturated with respect to primary and secondary silicates except for the solutions in the experiments using the 2-53 pm fraction. These solutions were supersaturated with respect to kaolinite and K-mica [Kal3Si3010(OH)2] after only 24 hours of dissolution. Consequently a second set of duplicate batch experiments were started using 0.1036 g and 1.224 g of powder. These experiments are on going and the 53 - 2 pm data will not be referred to further in this paper. The majority of elements analysed for showed an initial rapid rise in concentration followed by a slower, linear increase in concentration over time (e.g. Figure 1). The main exception to this was Ca which maintained a roughly constant concentration throughout the experiments (Figure 2). The other exception was Fe concentrations in the 250 - 53 pm experiments which remained below detection levels (5 ppb) until the last sampling date. Element release rates for Na, Mg, K, Fe, A1 and Si were determined in the following fashion:

Table 2. Size fraction mineralogy (W?). Size fraction

2000 500pm 2.4 0.0 10.1 80.1 0.0 6.3 0.3 0.1 0.0 0.7

Plagioclase feldmar 16 17 27 24

2.2 Batch experiments Batch experiments were carried out on duplicate subsamples of the different mineral size fractions. About 2.5 g of material was added to a 250 mL polypropylene, acid-washedybottle, containing pH 4 HC1. The bottles were placed in a shaking water bath set at 25 "C. Every 7 to 14 days the bottles were removed from the shaker and left to stand for 4 hours to give suspended particles time to settle. 25 mL of solution was then removed from the bottles using a pipette, filtered through a 0.2 pm filter, and acidified to a strength of 2.5 % HNO3. 25 mL of fresh, unreacted €€NO3 was then added to the bottles before they were placed back on the shaker. A blank experiment, containing no mineral powder but otherwise identical to the other batch experiments, was also run. After 12 weeks the experiments were stopped. All the solutions were analysed by inductively coupled plasma-optical emission spectroscopy (Ca, Si, AI, Fe, Mg) and atomic adsorption spectroscopy @a, K). Solution compositions were analysed using PHREEQ (Parkhurst & Appelo 1999) to check to see whether they were saturated with any phases. Element release rates from the different grain size fractions were calculated.

E = 0.25m / (A.M) Where E = element release rate expressed as moles of element released per gram of solid per second; m = the slope of a straight line fitted through the element concentration data for the various experimental solutions of the experiment. Except for the 250 - 53 pm Fe data the RSQ of these lines was > 0.8. A = atomic mass (pg) of element X; M = mass (g) of material originally in experiment. To convert E to a release rate expressed in mol/m2/s the mol/g/s release rate was divided by the surface area (m2/g) of the mineral powder that is dissolving. Element release rates normalised to both sample mass and sample surface area are given in Tables 4 and 5 respectively.

400

Table 5. Element release rates normalised to surface area (lOI4 x moi/m2/s). Size fraction Na 2000 - 500 pm 13.60 11.58 500-250 pm 8.21 4.18 4.69 250 - 53 pm 4.93

Mg 7.46 7.47 3.33 2.19 1.11 1.03

K 8.07 8.25 2.64 1.84 0.93 0.61

Fe 2.93 4.94 1.04 0.62

-

AI 23.23 23.89 14.61 13.67 2.46 5.89

Si 70.22 76.69 33.00 28.44 10.84 9.76

systematic trend in the mass normalised release rates is an increase in rate for Si with increasing grain size from the 250 - 53 pm fraction to the 2000 - 500 pm fraction. More systematic variations are seen in the surface area normalised release rates. Here there is an increase in the release rate of all elements from the finer 250 - 53 pm fraction to the coarser 2000 500 pm fraction. The contribution of the different grain size fractions to the element release from 1 g of minerals in the soil (Table 6 ) was determined by multiplying the mass normalised element release rates (E) by the mass fraction of the different grain size fraction present in the soil. These values are approximately the same if the analogous calculation is carried out using the surface area normalised release rates. The 2000 - 500 pm size fraction is seen to contribute most to the release of elements from one gram of soil but contributions of the 500 - 250 pm and 250 53 pm size fractions are very similar.

Figure 1. Changes in Si concentration over time in the batch experiments.

4 DISCUSSION It is not clear what the cause of the Ca behaviour was. Lack of Ca in the control experiment indicates that contamination was not the cause. Calculation of the total mass of Ca released into solution shows that the mineral powders should still contain undissolved Ca so Ca concentration is not limited by bulk Ca availability. Results from PHREEQ indicate that the solutions were not saturated with respect to Ca-bearing phases (which would have buffered the Ca concentrations). To investigate whether Ca concentrations were

Figure 2. Changes in Ca concentration with time in the batch experiments. Table 4. Element release rates normalised to mass (IOi4 x mollgls). Size fraction Na 2000-500 pm 5.85 4.98 500-250 pm 5.66 2.89 250-53 pm 7.65 8.03

Mg 3.21 3.21 2.30 1.51 1.80 1.67

K 3.47 3.55 1.82 1.27 1.52 1.00

Fe 1.26 2.12 0.72 0.43

-

A1 9.99 10.27 10.01 9.43 4.01 9.61

Si 30.19 32.98 22.76 19.62 17.68 15.90

Table 6. Number of moles of element released from the different grain fractions as 1 g of bulk soil dissolves for 1 second (1Oi4 x mol). Size fraction Na 2000- 500 pm 3.27 2.79 500 - 250 pm 0.68 0.35 250 - 5 3 pm 1.53 1.61

The element release rates of the duplicate experiments are similar, generally differences are less than a factor of two. Differences between the normalised release rates show more variation, differing by up to a factor of almost eight. The only

40I

MQ 1.80 1.80 0.28 0.18 0.36 0.33

K 1.94 1.99 0.22 0.15 0.30 0.20

Fe 0.70 1.12 0.09 0.05

-

AI 5.59 5.75 1.21 1.13 0.80 1.92

Si 16.91 18.47 2.73 2.35 3.54 3.18

being buffered by precipitation of a phase not included in the PHREEQ database the concentration of Ca in the two 2000 - 500 pm grain size experiments and in the control was increased to c. 1 ppm by the addition of CaCl2 solution. The pH of the solutions was not unduly affected by this procedure. The concentration of Ca remaining in solution was then determined one day and seven days after the addition of the CaC12 solution. Concentrations remained constant indicating that Ca content is not buffered by the precipitation of a Cabearing phase. Previous studies concerning the dissolution of granite (e.g. White et al. 1999) have shown that 1) many granites contain trace quantities of calcite difficult to detect using X-ray diffraction of whole rock samples and, 2) preferential dissolution of this calcite accounts for the bulk of Ca release from granites. It may be the case that dissolution of small concentrations of calcite in the different grain size fractions resulted in the initial high Ca concentrations in the experimental solutions discussed here. Further Ca release due to the weathering of Ca-bearing plagioclase may have been sufficiently slow or have released relatively small quantities of Ca so that the gradual further accumulation of Ca in solution was not detectable. The similarity of the calculated element release rates (Tables 3 and 4) would be expected given that the mineralogies of the different size fractions are very similar (Table 2). The 250 - 53 pm size fraction contains more plagioclase and less quartz and potassium feldspar than the coarser 2000 - 500 pm and 500 - 250 pm size fractions. Despite the relative amount of uncertainty associated with quantitative X-ray diffraction (Wilson 1987) the differences in quartz and plagioclase content are probably significant. Assuming that grain shape is constant between size fractions the surface area of grains in the finer size fractions will be higher than in the coarser fraction. Plagioclase is generally accepted as being more reactive than quartz (White & Brantley 1995) so if dissolution is a function of bulk surface area it is surprising that dissolution of the 2000 500 pm size fraction yields significantly higher element release rates than the 250 - 53 pm fraction and that the 500 - 250 pm and 250 -5 3 pm fractions have such similar rates. This implies that the plagioclase grains in the coarsest fraction contain more reactive sites than those in finer size fractions and that either reactive surface area is not proportional to BET surface area or the relative surface area of the different minerals in the different size fractions is not constant. This latter possibility would imply that the shape of grains of a given mineral change between grain size fractions. The dominant contribution of the 2000 - 500 pm grain size fraction to the bulk element releasz from the unfractionated mineral powder (Table 6) is due

to the greater mass of that fraction present in the soil compared to the 500 - 250 pm and 250 - 53 pm fractions. The lower element release rates of the 250 - 53 pm size fraction compared to the 500 - 250 pm fraction are offset by the greater proportion of soil mass and surface area that this fraction occupies. 5 CONCLUSIONS Both the surface area of mineral powders and their mass are important for determining the release of elements into solution as the minerals dissolve. In the soil considered here, the more massive, lowsurface area 2000 - 500 pm size fraction would contribute more to the element release from the bulk soil than the less-massive, higher-surface area 500 250 pm and 250 - 53 pm size fractions. Further work is aimed at determining dissolution rates of 53 - 2 pm and < 2 pm grain size fractions and carrying out similar experiments on more glacial and non-glacial soils. ACKNOWLEDGEMENTS Mike Andrews (PRIS), Franz Street (PRIS) and Anne Dudley (Soil Science) are thanked for help with XRD, XRF and ICPOES respectively.

REFERENCES Bock, R. 1979. A handbook of decomposition methods in analytical chemistry. International Textbooks. Brunauer, S., Emmett, P.H. & E. Teller 1938. Adsorption of gases in multimolecular layers. J. Amer. Chem. Soc. 60:309 - 319. Hooton, D.H. & N.E. Giorgetta 1977. Quantitative X-ray diffraction analysis by direct calculation method. X-ray Spectro. 612 - 5. Marschner, H. 1995. Mineral nutrition of higher plants. New York: Academic Press McKeague, J.A. & J.H. Day 1966. Dithionite- and oxalateextractable Fe and A1 as aids in differentiating various classes of soils. Can. J. Soil Sci. 46 13 - 22. Parkhurst, D.L. & C.A.J. Appelo 1999. User’s guide to PHREEQ (version 2.) Water Resources Investigations Report 99-4259. United States Geological Survey. White, A.F. & S.L. Brantley 1995. Chemical weathering rates of silicate minerals. Mineralogical Society of America. White, A.F., Bullen, T.D., Vivit, D.V., Schulz, M.S. & D.W. Clow 1999. The role of disseminated calcite in the chemical weathering of granitoid rocks. Geochim. Cosmochim. Acta 63: 1939 - 1953. Wilson, M.J. 1987. Clay mineralogy. London: Chapman and Hall.

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Wafer-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Geochemistry of a profile at the weathering front in dolomite H.B.Ji, S.J.Wang, Z.Y.Ouyang, C.Q.Liu, C.X.Sun & X.M.Liu State Key Laboratory of Environmental Geochemistry,Institute of Geochemistry, Chinese Academy of Sciences, Guiyang 550002, China

ABSTRACT: This paper presents a complete set of chemical data developed from one dolomitic weathering profile in the upland of a karst terrain in central Guizhou Province, China. Examination of the data from the weathering front indicates obvious transport of the elements in the weathering front. The strong enrichment of REEs at the rock/soil interface is the joint result of weathering of P-rich minerals, waterhock interaction and leaching of REE from the upper part of the profile. In the weathering front, as increasing weathering intensity, 87Sr/s6Srratio shows a general increase tendency and Sm/Nd ratio varies greatly, and the resetting of the isotopic systematics could be probably resulted from. Waterhock interaction is considered as one of the important pedogenic process in carbonate rock areas. 1 INTRODUCTION Weathering profiles are much more permeable than the underlying unaltered crystalline rock and permit waterhock interactions capable to produce significant chemical changes (Panahi et al. 2000). Previous authors thought that the REE could not be mobilized in the process of weathering. Later, as reported by Nesbitt (1979) and evidenced by several studies, REEs are mobilized in the process of weathering. Meanwhile, studies in recent years on the palaeoweathering recorded in sedimentary rocks and till weathering profiles revealed that variations would occur in Sm/Nd ratio and Nd isotopic composition during weathering (Bock et al. 1994, McDaniel et al. 1994, Macfarlane et al. 1994, Ohlander et al. 2000). The Sr isotopes have been used on a local scale, in an attempt to determine the losses of exchangeable cations and weathering rates (Miller et al. 1993). In using Sr isotope in the study of supergenic weathering and waterhock interaction the key problem is to fully understand the mechanism to cause variations in Sr isotopic composition of fluids and to identity the contributions of different horizons of soil sequences or different minerals to the source rocks (Innocent et al. 1997). Laterites in tropical and subtropical areas are of particular significance in this aspect because great variations in chemical composition would occur during their formation. The aim of this paper is to make a systematic

geochemical study on dolomitic weathering front in an attempt to shed light on geochemical processes in carbonate rock alteration, fractionation between elements and isotopes and their redistribution during supergenic weathering. 2 LOCATION OF STUDY AND ANALYTICAL METHOD The Pingba profile is situated on the upland of a karst terrain in central Guizhou Province, China (subtropical warm-humid area). The basement rocks are composed of pure (about 1% acid-insoluble residue) Triassic Anshun Formation dolomite (Tla), showing a gentle attitude, and measuring 712m in total thickness. The weathering profile is considerably thick and can be divided, from the top downwards, into the cultivated layer, the laterite layer, the ferrous crust, the reddish soil layer, the yellow soil layer and the bedrock layer. The weathering front includes the lower part close to soilhock interface in the weathering profile and the upper part of the bedrock layer and can be divided into the lower soil layer (Sample Nos. T-1 to T-10), the flour rock layer (Sample No. Y-3), the cracked rock layer (Sample No. Y-2) and the bedrock layer (Sample No. Y-1). Studies of acid-insoluble residues and comparisons of mineral compositions in the profile in conjunction with the REE distribution patterns and trace element ratios provide strong

403

Figure 1. Weathering characteristics as a function of depth at the weathering front. The detailed explanation see text.

evidence for in-situ weathering characteristics of the profile (Wang et al. 1999). Samples were collected from man-dug holes along a natural profile. Sample YT represents the residue leached with 1N chloric acid solution from the bedrock dolomite (Y-1) according to the extracting procedure described in Wang et al. 1999. The errors involved in measurements are: f 2% for major elements, rt 15% for trace elements such as Li, CO, Cr, Sr, Zr, Ba, Hf, and Ta, and k 10% for other trace elements. Repeated measurement of 87Sr/86Sr give a precision better than 0.0001; the recision of Rb/%r is within 1% and that of '43NdJ44Ndbetter than 0.00001; the precision of 147Sm/'44Ndvalues is within 0.5%.

3 RESULTS AND DISCUSSION 3.1 Transport of the elements According to the formulae presented in White et al. (1998), assuming Zr as the least mobile element, we have worked out the mass transport coefficient z j , w (Fig. 1), which reflects that at the weathering front from Y-1 to Y-2 and Y-3 the major elements Al, K and Mn are of extensive enrichment whereas Si, Mg and Na show a slight depletion; the trace elements Ga, Cs, Y and REEs are considerably enriched. Above soil/rock interface the major element A1 is greatly enriched while K is slightly enriched, and all other major elements show a depletion; Mg, Ca and Na are almost completely lost at the interface; the trace elements Ga and Cs are considerably enriched and other elements show varying degrees of

depletion; the LREEs abnormal enrichment is observed in soil samples near the interface with Ce showing no obvious variation; the other soil samples show a remarkable REEs depletion and a Ce enrichment. From the above results, we can see that both rocks and soils show a tendency to be enriched or depleted in the same elements; this tendency depends from the rapid dissolution of the major mineral dolomite and the gradual accumulation of acid-insoluble residues during dolomite weathering, as well as from the increasing degree of oxidation in going upwards the interface. The enrichment of Mn in flour rocks is due to waterhock interaction, that of the major elements A1 and K is the result of accumulation of acid-insoluble residues. The trace element Ga maintains a close relation with A1 in weathering and pedogenic environments (Panahi et al. 2000) while Cs is largely enriched in feldspar minerals and can be easily adsorbed on clay minerals. That is why the enrichment of both Ga and Cs may be related with the accumulation of acidinsoluble residues. The inconsistent variation of the REEs seems to be related with the change of redox environment at the interface. Ce tends to be separated from the other REEs because of its inactivity in the supergenic oxidation environment. 3.2 Fractionation process of the REEs From Y-1 to Y-2 and to Y-3 the total REE amount and GdN/YbN ratio tend to increase whereas the 6Ce values tend to decrease remarkably (Fig.2). The above results may be explained by: (1) the sedentary accumulation of leaching residues of dolomite and of minerals with MREE- and HREE-rich distribution

404

patterns; or/and (2) the occurrence of supergenic waterhock interactions at the weathering front of dolomite. The obvious reduction of 6Ce values indicate that supergenic oxidation during the weathering process has caused the loss of Ce (Banfield & Eggleton 1989). From Y-3 to T-1 the total REE amount tends to increase drastically, especially LREEs and MREEs, and the 6Ce values are so abnormal as to reach 0.01. This phenomena is jointly resulted from great change in volume and waterhock interaction from the alteration of flour rock to soil as well as the accumulation of REEs leached from upper profile. The remarkable negative anomalies are indeed in conformity with the weathering characteristics of P-rich minerals (Banfield & Eggleton 1989). For the other soil samples, with the exception of Ce, the other REEs were lost greatly as a result of leaching. Comparatively, the loss of the HREE is greater than that of the LREE (Fig.2).

Figure 3 . Diagram of Nd versus Sr isotopic composition in rocks and soil samples.

represent the same process implied by line (I). As derived from our data, both primary rocks and soil samples from the profile have the identical model ages (T=183 Ma), reflecting that the two different isotopic systems have undergone the same resetting process and also reflecting that this profile is an insitu weathering profile. So we think that it is the waterhock interaction which took place at the modern weathering front of dolomite that led to RbSr isotope variations and REE fractionation.

4

CONCLUSIONS

1. At the present-day weathering front of dolomite the elements have experienced obvious transport, which agrees with the process of accumulation of “insoluble” constituents in carbonate rocks at the weathering front. 2. Fractionation of REEs is noticed during weathering processes; the REEs in the soil layer Figure 2. The chondrite-normalized REE distribution patterns in the Pingba profile.

3.3 Resetting of Nd- Sr isotopic systematics A good linearity for two component mixing in the profile from protolith+soil-+leaching-dissolution residues occurs in Nd-Sr isotope space (Fig.3), which provides strong evidence for in-situ weathering characteristics of the profile. Meanwhile, the depletion of Nd in Y-3 is obvious to be mirrored by its enrichment in Y-2, indicating the preferential removal of some components from dolomite. In Figure 4, at the weathering front of dolomite Nd and Sr isotopes show obvious variations and two approximately parallel lines occur between the profile of Y-l+Y-2+YT and a soil sample (line I) and between Y-3 and soil samples close to the interface (line 11). The first line (I) is indicative of the process of acid dissolution of dolomite and the process of accumulation of acid-insoluble residues; the second parallel line (11) is considered to

Figure 4. Rb-Sr and Sm-Nd isochron diagrams of rocks and SOi1 samples.

405

above the rockkoil interface are abnormally enriched; remarkable negative Ce anomalies are the result of weathering of P-rich minerals, interaction and leaching of REE from the upper part of the profile. 3. The dolomite weathering process in this study is accompanied with waterhock interaction, probably causing the opening of the Nd-Sr system and the resetting of these two kinds of isotopes in the supergenic environment.

1998. Chemical weathering in a tropical watershed, Luquillo Mountains, Puerto Rico: I. Long-term versus short-term weathering fluxes. Geochimica et Cosmochimica Acta 62: 209-226.

ACKNOWLEDGEMENTS This research project was granted jointly by the National Natural Science Foundation of China (Grant No. 49833002) and the State Climbing Program (95-pre-39).

REFERENCES Banfield, J.F. & R.A. Eggleton 1989. Apatite replacement and rare earth mobilization, fractionation, and fixation during weathering. Clay and Clay Minerals 37: 113- 127. Bock, B., McLennan, S.M. & G.N. Hanson 1994. Rare earth element distribution and its effect on the neodymium isotope system in Austin Glen Member of the Normanskill Formation, New York, USA. Geochimica et Cosmochimica Acta 58: 5245-5253. Innocent, C., Michard, A., Malengreau, N., Loubet, M., Noack, Y., Benedetti, M. & B. Hamelin 1997. Sr isotopic evidence for ion-exchange buffering in tropical laterites from the Parana, Brazil. Chemical Geology 136: 219-232. Macfarlane, A.W., Danielson, A., Holland, H.D., & S.J. Jacobsen 1994. REE chemistry and Sm-Nd systematics of late Archean weathering profiles in the Fortescue Group, Western Australia. Geochimica et Cosmochimica Acta 58: 1777-1794. McDaniel, D.K., Hemming, S.R., McLennan S.M. & G.N. Hanson 1994. Reseting of neodymium isotopes and redistribution of REEs during sedimentary processes: The early Proterozoic Chelmsford Formation, Sudbury basin, Ontario, Canada. Geochimica et Cosmochimica Acta 58: 93 1-94 1. Miller, E.K., Blum, J.D. & A.J. Friedland 1993. Determination of soil exchangeable-cation loss and weathering rates using Sr isotopes. Nature 362: 438-441 Nesbitt, H.W. 1979. Mobility and fractionation of rare earth elements during weathering of a granodiorite. Nature 279: 206-2 10. Ohlander, B., Ingri, J., Land, M., & H. SchOberg 2000. Change of Sm-Nd isotope composition during weathering of till. Geochimica et Cosmochimica Acta 64: 8 13-820. Panahi, A., Young, G.M. & R. H. Rainbird 2000. Behavior of major and trace elements (including FEE) during Paleoproterozoic pedogenesis and diagenetic alteration of an Archean granite near Vil!e Marie, Quebec, Canada. Geochimica et Cosmochimica Acta 64: 2 199-2220. Wang, S.J., Ji, H.B., Ouyang, Z.Y., Zhou, D.Q., Zheng, L.P., & T.Y. Li 1999. Preliminary study on weathering and pedogenesis of carbonate rock. Science in China (Ser.D) 42: 572-581. White, A.F., Blum, A.E., Schulz, M.S., Vivit, D.V., Stonestrom, D.A., Larsen, M., Murphy, S.F. & D. Eberl

406

Water-Rock lnferaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

The use of U-isotopes on the study of a weathered cover in Paran6 basin, Brazil J .R .Jim&ez-Rueda & D .M .Bonotto Instituto de Geoci&cias e Ci2ncias Exatas, UNESP, Rio Claro, SGo Paulo, Brazil

ABSTRACT: This paper describes the results of a study involving the sampling of a soil profile developed over sandstone from Itarark Sub-Group at the Parana sedimentary basin, Brazil. It was carried out to integrate chemical and U-isotopes data in order to improve the knowledge of the weathering processes acting in the area. 238U and its daughter 234Uproved to be important tools for evaluating physical and chemical alteration, allowing to suggest a possible timescale for the development of the more superficial soil horizons. isotopes data were integrated for evaluating the processes affecting the chemical weathering of an important lithofacie occurring at S5o Paulo State, Brazil, and widely disseminated in Parana sedimentary basin.

1 INTRODUCTION Chemical weathering of rocks has been much investigated since alteration is one fundamental phenomenon leading to the present geomorphology of the continents. Many approaches have been realized to group elements according to their relative mobility in the weathering zone, to evaluate dissolutioddeposition processes, and to estimate chemical weathering rate of rocks, where the natural radionuclides belonging to the U and Th decay series proved to be useful for these purposes (Hansen & Stout 1968, Moreira-Nordemann 1980, Michel 1984, Colman & Dethier 1986). In particular, 238Uand its daughter nuclide 234Uallow to generate 2 3 4 ~ / 2 3 activity 8~ ratios (ARS) suggestive of processes occurring in soil profiles since the present climatic conditions up to the last one million years ago (Latham & Schwarcz 1987). This evaluation is based on the condition that secular equilibrium is established between 234Uand 238Uin all rocks and minerals that are closed systems for U, i.e. AR=l within the bulk of most rock matrices. However, dissolution processes at the rocWsoi1-water interface frequently result in AR >1 for uranium dissolved in the liquid phase and AR < 1 for uranium in the solid phase (Cherdyntsev 1971, Osmond & Cowart 1976). Furthermore, uranium has been considered very insoluble under reducing conditions, occurring its most active etch solution where oxidizing conditions generally prevail (Osmond & Cowart 1976, Langmuir 1978). This paper evaluates the chemical weathering responsible for the generation of a thick soil profile developed under tropical climate dominated by intensive wet-dry seasons. Geochemical and U-

2 SAMPLING AND RESULTS It was investigated a soil profile developed over the Itarare Sub-Group and located in the 50-km wide Depress50 Periferica geomorphological province. The studied area is situated in Limeira city, about 150 km distant from S5o Paulo city. Several detritic lithofacies comprise the Itarare Sub-Group, where the main lithological types are massive feldspathic or arkoseous sandstones of fluvial, marine, lacustrine, deltaic or eolic origin, varying considerably in texture and fossil content. It was sampled the weathered cover developed over an altered sandstone matrix 17-m deep and 6 km distant from Limeira city. Some 2 kg of residual rock and soil samples representative of different horizons in the profile developed over Itarare Sub-Group were collected with a 30-cm long, 10-cm diameter galvanized steel auger bucket connected to a cross handle by a 1-m long extension. The samples were described in the field, placed in plastic bags, sealed, and after drying and mixing, aliquots were separated for mineralogical, physical and chemical identification. X-ray diffraction was utilized in the mineralogical identification of the samples, whereas the major element analysis was performed by X-ray fluorescence spectrography, and the organic matter by colorimetry (Allman & Lawrence 1981). 407

Table 1. Physical and chemical evaluation of the soil profile developed over sandstone from ItararC Sub-Group, Brazil. PARAMETER

HORIZON CODE

UNIT

Residual Rock Bol Bo2 Bo3 BC Cr Oxic Oxic Oxic Transitional Altered Rock Horizon Name Reddish Reddish Reddish Red to light Red to light Light Color brown brown brown red reddish brown reddish brown 3.20-4.50 m 0.00-0.26 0.26-0.84 0.84-1.50 1 SO-3.20 4.50-7.00 15.00-17.00 Interval of Depth 0.55 1.17 2.35 3.85 0.13 5.75 16.00 m Average Depth 1.61 1.73 1S O 1.58 1.80 g/cm3 1.65 2.1 1 Bulk Density 2.54 2.58 2.63 g/cm3 2.60 2.56 2.57 2.50 True Density 86.07 83.46 87.83 79.78 86.82 wt. % 80.40 Total Sand 7.84 9.18 6.25 10.00 7.31 10.20 wt. % Total Silt 13.92 16.53 12.67 20.22 wt. % 13.17 19.59 Total Clay 41 39 39 16 33 wt. % 29 36 Total Porosity 72.68 80.65 81.64 79.59 82.23 67.77 85.27 wt. % SiOz 4.94 7.38 4.96 8.39 2.82 5.78 wt. % 3.83 A1203 0.076 0.040 0.045 0.057 0.073 0.069 0.016 wt. % NazO 0.12 0.21 0.10 0.33 0.44 0.31 wt. % 0.10 K20 0.066 0.092 0.048 0.052 0.053 0.068 0.056 wt % CaO 0.084 0.127 0.092 0.139 0.054 0.191 0.065 wt. % MgO 1.65 1.22 2.15 0.44 2.54 wt. % 1.60 1.40 Fe203 0.008 0.004 0.012 0.006 wt. % 0.008 0.008 0.007 MnO 0.40 0.58 0.36 0.05 0.44 0.53 wt. % 0.35 TiOz 0.022 0.018 0.016 0.013 0.007 0.013 0.020 wt. % p205 0.82 0.28 1.29 0.37 0.34 0.62 0.47 wt. % Organic Matter 6.78 13.81 8.19 19.98 8.79 11.07 26.39 wt.% Other' 6.7 1 1.19 7.60 1.40 2.3 1 9.10 7.88 U2 P%/g 2 3 4 ~ 1 2 3 AR 8~ 1.06kO.10 0.97M.09 1.06+0.11 1.17M.18 0.80M.08 0.87M.08 0.90M.07 Other elements and compounds not analyzed, with predominance of adsorbed water, water in crystal lattices and fluid inclusions, COz of carbonates, and SO2 of sulfides; Uncertainty k10% corresponding to 10 standard deviation. AD Ocric Dark gray

'

Measurements of U content and AR were realized by alpha spectrometry after complete dissolution of the samples in a mixture of HN03+HF+HC104 followed by a radiochemical procedure for extraction and electrodeposition of uranium isotopes on stainless-steel planchets (Veselsky 1974, Osmond & Cowart 1976, Ivanovich & Harmon 1982). The description of the samples and their physical, chemical and isotope characterization is shown in Table 1.

presence of ferrite, bauxite, and kaolin (Fig. 1) shows that kaolinization processes are taking place in the soil horizons and altered parent rock. Such finding was confirmed by the use of X-ray diffraction in the mineralogical identification of the samples, which indicated the presence of kaolinite in all them. Because A1203 is the constituent most often chosen for identifying real gains and losses during chemical weathering (Colman 1982, Faure 199 l), this oxide was considered to remain constant in

3 MAJOR COMPOUNDS RELATIONSHIPS The analysis of the particle size distribution indicated the predominance of sand in all samples (80-88 wt.%), and an excellent inverse correlation (r = -1) between the sandy and clayey fractions. The insertion of the chemical data obtained for the analyzed samples in the Si02-AI203-Fe203 triangle proposed by Schellmann (1979) for evaluating the degree of weathering indicated that all horizons fall into the field of the sandstone matrix dominantly composed by Si02. Furthermore, the insertion of the same data in the diagram proposed by Balasubramaniam et al. (1983) for defining the

Figure 1. Data for the analyzed samples plotted on the Si02-A1203Fe2O3diagram proposed by Balasubramaniamet al. (1983).

408

correlate reasonably (r = 0.82), with the organic matter tending to decrease with increasing depth in the soil horizons (r = - 0.85) (Fig. 3). The same trend was found for uranium (r = -0.86) (Fig. 3), despite of the non-existence of correlation between organic matter and uranium. A possible explanation for such discrepant behavior is the different particle size distribution along the soil profile, also caused by influence of physical factors like humidity and temperature in addition to chemical processes acting there. Thus, there is a decrease of the sand fraction with increasing depth (r = -0.92), whereas an increase of the clay fraction with increasing depth (r = 0.93) (Fig. 3). As expected, K20 correlates very well with the amount of clay in the soil horizons (r = l), and, consequently, K2O content also increases with increasing depth (r = 0.93) (Fig. 3). In contrast, U content decreases with increasing K20 content (r = 0.88), as well with the increase of the clay fraction (r = -0.84). Uranium only correlated directly with the amount of sand fraction (r = 0.82) that increases towards the more superficial soil horizons. This suggests that it is contained in resistatedminor refractory minerals (such as zircon) which are highly resistant to weathering, so that they are incorporated intact in the coarser grained fraction of the soil horizons. The four more superficial soil horizons exhibit A R s corresponding to unit, within experimental errors, indicating radioactive equilibrium between 234Uand 238Uin the detritic matrix at least over the last one million years (Latham & Schwarcz 1987). Thus, there is no occurrence at these sites of extra 234U-losseither due to a-recoil effects (Kigoshi 1971) or by preferential leaching/etching of recoil-damaged sites

amount in the investigated soil profile, despite its concentration appeared to have changed. Under this assumption, it was possible to verify that Mg, P, Fe, Mn, and Ti also accumulated in all soil horizons, yielding significant correlations between A1203 and MgO (r = 0.97), A1203 and P205 (r = 0.92), AI203 and Fe203 (r = 0.90), A1203and MnO (r = 0.88), and A1203 and Ti02 (r = 0.87) (Fig. 2). Micas, feldspars, clays, apatite, and amorphous Fe-Mn oxi-hydroxides are the main probable sources of these constituents. Profuse rainfall (about 1.6cdyear) characterizes the studied area, favoring the occurrence of substantial losses of silica and bases in all soil horizons relatively to the slightly altered parent rock, i.e. 3967 wt.% Si02, 55-84 wt.% Na20, 66-87 wt.% K20, and 39-63 wt.% CaO. The remaining Na2O correlated significantly with Si02 (r = 0.85) and K20 (r = 0.78), possibly reflecting the presence of feldspars, micas and clays as their source minerals. 4 URANIUM RELATIONSHIPS Uranium does not accumulate in the soil horizons associated with AI$&, MgO, P205, Fe203, MnO, and Ti02, since no significant correlation was found among their concentrations. Organic matter has often be considered an important complexing agent for uranium (Szalay 1964), however, in the studied soil profile the correlation between these parameters also is not significant (r = 0.64). As expected, the data obtained for organic matter and loss on ignition

Figure 2. Chemical data from the studied area plotted against depth.

Figure 3. Chemical and U-isotopes data from the studied area plotted against depth.

409

(Fleischer 1975). However, preferential removal of 234Uhas occurred at the slightly altered parent rock, and Cr and BC soil horizons, generating ARsl for uranium dissolved in the liquid phase, and, consequently, ARs POE

E

2.67

1.60

40.0

35.0

87.5

0.75

5.26

-7.2

-6.2

0.7

2.55

1.50

41.5

35.9

87.3

variations weakly welded

I-E

A -15 -17 -15 -23 47 75 40 217 -15 -18 283 117

-

ppm

V Ni Zn Rb Sr Ba Zr Pb Nb Y La Ce

I 29.5 3.10 25.6 223 49.2 325 144 26.5 6.50 18.6 24.2 51.5

E A 30.0 2 18.1 484 405 1481 255 14 162 228 556 71 151 5 73.3 177 nd -1.30 -93 24.5 I 36.8 -29

The formation of reprecipitation crusts is facilitated in those areas where water is able to penetrate and subsequently evaporate (lower porosity). On the other hand, mechanical alteration phenomena such as thermoclastism and crioclastism enhance surface porosity and, depending on exposure, lead to a prevalence of physical breakdown over chemical transformation. For strongly-welded pyroclastics, a positive correlation has been observed between the imbibition coefficient related to volume and the saturation index in slightly-vitreous facies (macropores, generated by microfractures). On the other hand highlyvitreous facies show an inverse correlation due to of their greater resistance to physical weathering. In moderately-welded facies, hewn stones are weathered to greater depth. They are not so compact and contain lower percentages of massive glass. Samples analysed showed greater mean porosity and lower apparent density, due to lower welding temperature and the presence of evenly-distributed microporosity, and higher imbibition coefficients (Table 1). In some cases, the saturation index was over 100%; this mechanism may be explained by chemically induced permeability through pre-existing barriers (Bralia et al. 1995). Based on the higher or lower open porosity in the outer part, weakly-welded pyroclastic rocks have been divided into two groups. Those in the first group have higher open porosity in the uppermost (41.0%) than in the innermost part (37.6%). The simultaneous reduction in bulk and real density indicates loss of material due to physical breakdown. The largest variations measured show an increase in open porosity of up to 28% and a decrease in apparent density of 15.5%. Pyroclastic rocks in the second group have lower porosity in the outer (40.0%) than in the inner part (43.1%). This behaviour can be attributed to reprecipitation of material of diverse origin (Manganelli Del Fa et al. 1989) within the surface crust. XRD

Table 1. Mean values of physical properties. d = real density, g/cm3; a,= bulk density, g/cm3; PO=open porosity, %; IC,= imbibition coefficient, volume%; SI = saturation index, %. I = inner; E = outer. facies

E 55.29 0.40 10.58 2.81 0.81 0.14 0.42 6.81 1.05 7.20 8.00 6.48

417

ACKNOWLEDGMENTS

analyses revealed wheddellite in the surface crust of some of the samples, which may have originated from: 1) microbial activity which produces oxalic acid from which oxalate is formed; 2) substances used in the past as preservatives or for other purposes. As in the first group, the different apparent density marks the passage from the surface to the inner part of the stone, showing higher values in the reprecipitation crust (1.60 g/cm3) and lower values in the less weathered inner matrix (1.52 g/cm3). In the first group, the saturation index generally decreases outwards, whereas in the second group, mean variation is not significant. Compared to the other facies, weakly-welded pyroclastic rocks differ in other aspects. They are only slightly coherent and are weathered to a greater extent. They contain abundant pumice and lithic fragments, and no significant differences have been observed between the outer and inner part. They are normally very porous, having higher mean porosity than the other facies. This behaviour is associated with the presence of syngenetic macropores, as the result of low lithification and slight thickening.

This work was carried out with the support of the National Scientific Research Programme: "Italy's Historical Heritage of Stones: Knowledge Aimed at Conservation. Method Checking and Application to Significant Urban and Territorial Cases". COFIN 1999 (National Coordinator C. D'Amico; Local Coordinator G. Macciotta).

REFERENCES

4 CONCLUSIONS The volcanic rocks used to build the S. Antioco of Bisarcio Basilica outcrop in the vicinity of the church and can easily be distinguished from other pyroclastic rocks nearby. In general, the former are more compact and resistant, and as no earlier settlements are known to have existed in that area, it is plausible that the location of the Basilica was dictated by the availability of good building material nearby. Water-rock interaction in construction materials can be summed up into two extreme types of behaviour: 1) physical disaggregation of pyroclastic rocks prior to the onset of geochemical-mineralogical alterations precluding new isochemically-formed reaction phases. In this context, the scattering of Na20 contents warrants further investigation; 2) disaggregation and increased porosity in those parts more exposed to atmospheric agents with reprecipitation of salts of varying origin and consequent allochemical transformation. In conclusion, low chemical and high physical decay can be attributed to medium-to-low lithification and thickening of the material and to a limited period of water-rock interaction.

418

Beccaluva, L., Civetta, L., Macciotta, G. & C.A. Ricci 1985. Geochronology in Sardinia: results and problems. Rend. Soc. It. Min. Petr. 40: 57-72. Bralia, A., Ceccherini, S., Fratini, F., Manganelli Del Fa, C., Mellini, M. & G. Sabatini 1995. Anomalous water absorption in low-grade serpentinites: more water than space?. European Journal of Mineralogy 7 : 205-215. Cherchi, A. & L. Montadert 1982. The Oligo-Miocene riR of Sardinia and the early history of the western Mediteranean basin. Nature 298: 736-739. De La Roche, H., Leterrier, J., Grand Claude, P. & M. Marchal 1980. A classification of volcanic and plutonic rocks using RI-R2 diagram and major element analysis. Its relationships with current nomenclature. Chem. Geol. 29: 183-210. Ghiara, M. R., Lonis, R., Petti, C., Franco, E., Luxoro, S. & G. Balassone 1997. The zeolitization process of Tertiary orogenic ignimbrites from Sardinia (Italy): distribution and meaning importance. Per. Mineral. 66: 21 1-23 1. Lecca, L., Lonis, R., Luxoro, S., Melis, F., Secchi, F. & P. Brotzu 1997. Oligo-Miocene volcanic sequences and rifting stages: a review. Per. Mineral. 66: 7-61. Manganelli Del Fa, C., Camaiti, M., Borsell., G. & P. Tiano 1989. Variazioni della quantita di acqua di cristallizzazione dell'ossalato di calcio in funzione delle condizioni termoigrometriche. Atti del Convegno: Le pellicole ad ossalati: origine e signijicato nella conservazione delle opere d'arte. Centro C.N.R. G. Bozza Ed., Milano: 91-97. Morbidelli, P., Ghiara, M.R., Lonis, R. & A. Sau 1999. Zeolitic occurrences from Tertiary pyroclastic flows and related epiclastic deposits outcropping in northern Sardinia (Italy). Per. Mineral. 68: 287-3 13. Savelli, C., Beccaluva, L., Deriu, M., Macciotta, G. & L. Maccioni 1979. WAr geochronology of the Tertiary "CalcAlcalik" volcanism of Sardinia (Italy). Journ. Volc. Geoth. Res. 5: 257-269.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Characterization of oxidation products onto pyrite: coupling of XPS and NMA F.Mercier UMR 8587 "Analyse et Environnement ", CEAICNRSIUniversite' d 'Evry Val-d'Essonne, France

M.Descostes & C.Beaucaire Lnboratoire d'e'tudes des lnte'ractions Roche Eau, DCCIDESDISESD, CEA, 91191 Gif sur Yvette, France Lnboratoire de Ge'ochimie des E a u , Universite' Paris 7 & lnstitut Physique du Globe de Paris, France "Present address: IPSNIDPREISERGD, CEA-FAR,BP n06, 92265 Fontenay aux Roses, France

P.Trocellier

Laboratoire Pierre Sue, CEA, France

P.Zuddas Laboratoire de Giochimie des E a u , Universite' Paris 7 & lnstitut Physique du Globe de Paris, France

ABSTRACT: X-ray Photoelectron Spectroscopy (XPS) has been coupled to Nuclear Microprobe Analysis (NMA) to characterize thin oxidation layers onto pyrites. XPS evidences both oxidation state and chemical environment of S and Fe. NMA informs about spatial distribution and chemical composition heterogeneity of oxidation products. Pyrites oxidized in acidic medium produces few solid phases. Only Fe" sulfate is detected on oxidized pyrite surface. In carbonate medium, oxidation layer is more complex. Iron is mainly in an oxidation state (11) under siderite or Fe" sulfate form. Sulfur oxidation induces intermediate species (polysulfides and sulfoxyanions as S ~ 0 3also ~ - evidenced in solution) indicating that oxidation occurs at solid state before dissolution. NMA has shown that oxidation occurs only on localized points of pyrite surface, with oxidation layers showing spatial distribution and thickness heterogeneities. literature concerning Rutherford Backscattering Spectrometry (RBS) studies of sulfide surfaces oxidation is very sparse and only concerns use of macrobeams with a scientific approach different from ours. Hence Ponsot et al. 1998 have used RBS with incident 160sc ions at 6 MeV to investigate the thermal oxidation kinetics of galena (PbS) in air and have shown that the oxidized layer thickness varies from some tens to few hundreds nanometers. Pratesi & Cipriani (2000) have applied RBS with incident H+ ions at 2.4 MeV to study surface alteration of iron sulfides strongly oxidized after two years of air exposure. Our paper is devoted to evidence spatial distribution (homogeneity or heterogeneity) and chemical nature of oxidation products derived from pyrites oxidized in different media (acidic and carbonated) by coupling XPS and NMA studies.

1 INTRODUCTION Pyrite (FeS2) oxidation knows a large and consequent bibliography (Lowson 1982, Evangelou 1995). Pyrite is involved in acid mine drainage when oxidative conditions are encountered. The oxidation of 1 mole of FeS2 leads to 2 moles of sulfuric acid, according to the overall reaction 1, FeS2 + 7/2 0

2

+ H2O -+ Fe2++ 2 SO-:

+ 2 H+ (1)

The important number of pyrite oxidation studies illustrates both its importance in environmental damage and the fact that oxidation mechanism is not yet well understood. X-ray Photoelectron Spectroscopy (XPS) is suitable to determine both oxidation state and chemical environment of S and Fe. Many XPS studies are available (Bonnissel-Gissinger et al. 1998, Brion 1980, Sasaki et al. 1995), but the nature of oxidation products depends on preparation conditions. Moreover, most of papers issued from literature reported results acquired from FeS2 surfaces already oxidized before their treatment. Nevertheless, spatial repartition of oxidation products remains unknown by XPS. The same remark applies to thickness and elementary composition of oxidation products. Nuclear Microprobe Analysis (NMA) by using resonant nuclear reactions on light elements as C and 0, permits to determine if the spatial distribution is homogeneous or heterogeneous. To our knowledge,

2 EXPERIMENTAL AND METHODS 2.1 Pyrite preparation Cubic samples of FeS2 (Logroiio, Spain) were first dipped with concentrated HCl (37%) during several hours to eliminate any oxidation products present at the mineral surface. The pyrite was then introduced in a glove box (p(H20) and p(02) < 1 vpm) and rinsed with acetone. The mineral was ground in an agate mortar and sifted with ethanol (grain sizes in the 150-250 pm fraction). FeS2 was then washed in an ultra-sonic bath to remove any fine particles adhering to the grains surface. Samples were kept in 419

glove box for drying until experiments. Batch experiments were run in a glass electrochemical cell used as reactor in contact with atmospheric oxygen (20%). Temperature was regulated through a heating-bath circulator at 25.0kO.l "C. Agitation by a TeflonB stirring bar guaranteed a continuously homogeneous solution. The water to solid ratio was 150 mL.g-'. Time course begun with FeS2 introduction in solution. Dissolution experiments were carried out in two different media: perchloric acid (1Om2 mol/L) and carbonated solution ([HCO3]=I, 12.10-~ mol/L). TWO contact times were selected: 6 and 24 hours. The final samples were kept in glove box before XPS and NMA investigations.

Figure 1. (aa)spectrum of a freshly prepared FeS2.

2.2 Analytical techniques

X-Ray Photoelectron Spectroscopy (XPS) Experiments have been carried out at CEA Saclay (France), using a VG Escalab MKII spectrometer with an unchromated AlK, (1486.6 eV) radiation. Binding energies (BE) positions of Cls, S2p and Fe2py2 are presented in this paper to evidence of chemical nature of oxidation products (see Descostes et al. 2000 for experimental procedures). Nuclear Microprobe Analysis (IVMA) Nuclear reactions 12C(pp)'2C and '60(aa)'60at resonance energies of 1.725 and 3.05 MeV respectively have been applied to determine C and 0 in oxidation products located in surface thin layers. The first nuclear reaction permits to evidence C in an oxidation layer of 1 pm maximum thickness whereas the second one gives informations about 0 contained in a layer of 33 nm maximum thickness. (Mayer & Rimini 1977). Moreover, the use of the particle backscattering permits to evidence Fe and S edges and to measure elementary ratios. Analyses have been performed using the microprobe of the LPS Laboratory (CEA Saclay France). Punctual measurements and elementary map ings from surfaces beam scanning of 115x 155 pmPdimensions have been realized. A beam diameter of 15 pm and a beam current density of 1 pA/pm2 were chosen.

traducing the existence and the diversity of oxidation products onto FeS2 surface. These spectra suggest that oxidation layer presents elementary composition variations. Indeed, some configurations are evidenced: presence of 0 without C @resence onto surface of iron oxides or oxidation products of S), presence of both 0 and C (presence onto surface of carbonate species). As illustrated on figure 3, atomic S/O ratios issued from two (pp) spectras of a FeS2 prepared in carbonated medium is of 0.6 (intermediate species such as S ~ 0 3 ~and - ) of 0.2 (SO?- species). In addition to elementary composition variations, a heterogeneous repartition of oxidation layer onto FeS2 surfaces is evidenced for all pyrites (acidic and carbonated media), as seen on figure 4 displaying 0 mapping of a FeS2 prepared in acidic medium. The darker parts of this mapping are attributed to highest 0 concentrations. A cluster richer in 0 appears in the top at the right of mapping. The comparison of (aa)spectra inside and outside the cluster, confirms the heterogeneous repartition of the oxidation layer: presence of 0 signal in cluster and absence of 0 outside with only Fe and S edges corresponding to Fe&. As seen in XPS, no C signal has been observed onto this pyrite.

3 RESULTS AND DISCUSSION 3.1 NMA informations Figure 1 presents the (aa) spectrum of a freshly FeS2 prepared in glove box. No oxidation product is detected and the mean of Fe/S atomic ratio is of 2.09kO. 19, respecting the FeS2 stechiometry. In contrary, for a FeS2 oxidized in carbonated medium, the spectra of the same pyrite at two different locations reveal a sharp 0 signal on (aa) spectrum associated to the C signal on (pp) spectrum (Fig. 2a) or not correlated to the C signal (Fig. 2b), 420

Figure 2. (pp) and (aa)spectra obtained at two different locations onto a same FeS2 surface oxidized in carbonated medium (contact time=24h).

Figure 3. Examples of (pp) spectra of a FeS2 oxidized in carbonated medium (contact time=6h).

3.2 XPS informations C1s peak The comparison of C 1s peaks for oxidized FeS2 and siderite C l s peaks is displayed on figure 5. Oxidized FeS2 in carbonated medium presents a component close to 289.00 eV, which confirms the presence of Fe carbonate complex at the mineral surface. This component does not appear in acidic media, which is normal since PKA~(H2CO3 /HCO3') is 6.46 and no Fe carbonate complex can form at pH=2.

S2p peak Figure 6 compares oxidized pyrites and S standards S2p region. S2p peak shows two components at 162.30 eV (FeS2) and 169.00 eV attributed to sulfate. SO-: contribution tends to raise with pH and time. In carbonated medium, two contributions near 163.50 and 165.00 eV appear with S oxidation state comprised between (-1) and (IV), i.e. between FeS2 and sulfite

Figure 4. 0 mappings by (aa) of a FeS2 oxidized in acidic medium (a) and comparison of (aa)spectra inside the cluster (b) and outside (c) (contact time=24h).

Figure 5. Comparison of BE of C l s photoelectron peaks of oxidized FeS2 in carbonated medium to siderite.

Fe2p peak Figure 7 compares oxidized FeS2 and Fe standards Fe2p peaks. The Fe2p peak shows two main components at 707.15 (FeS2) and a shoulder close to 71 1.OO eV. Its intensity raises with dissolution time meanwhile BE shifts to higher energy with higher pH, traducing a Fe chemical environment containing more 0 atoms. Thus, XPS shows that acidic oxidation produces few solid components. Only Fe"SO-: is detected on oxidized pyrite surface. Sulfate and Fe (Fe", Fe"') are very soluble at acidic pH. Neither aqueous nor solid sulfoxyanion is discerned. In carbonated medium, XPS spectra are more complex. Fe shows several chemical environments and two oxidation states. Fe is mainly with an oxidation state (11), apart from the oxidized FeS2 first layers where Fe"' may exist as illustrated by the comparison of Fe3p and Fe2p3/2 peaks (Descostes et al. 2000). S shows several oxidation states as the pH increases (SO:-, sulfox anion and polysulfoxyanion close to S032-and S203E). Besides, SO:-, S032-and

Figure 6. Comparison of BE of S2p photoelectron peaks of oxidized FeS2 in acidic and carbonated media to S standards.

421

NMA has shown that oxidation occurs only on certain points of pyrite surface. The oxidation layer displays spatial repartition and thickness heterogeneities. Besides, in carbonated medium, the nature of oxidation products varies, traducing as in XPS the formation of intermediate species of sulfur as S2032-and the presence of FeC03. ACKNOWLEDGMENTS The support of the L'Agence Nationale pour la gestion des Dechets Radioactifs" through grant FT 00- 1-066 is gatefully acknowledged.

REFERENCES

Figure 7. Comparison of BE of Fe2p photoelectron peaks of oxidized FeS2 in acidic and carbonated medium with Fe standards.

S2032-have been detected in solution (Descostes et al., personal communication). In carbonated medium, FeS2 knows a solid state oxidation. S2032by its redox stability and low oxidation state seems to be the first sulfur aqueous intermediate species produced during pyrite oxidation, as suggested by Luther (1987). This species is not observed in acidic medium, where either the layers concerned are under the XPS analyzed depth (30-50 A, i.e. 5-10 atomic layers), either the dissolution of oxidation products is too rapid to allow this species to be observed.

4 CONCLUSION XPS and NMA are complementary techniques to characterize pyrite oxidation layers. XPS permits to determine both oxidation state and chemical environment of S and Fe. NMA gives informations about spatial distribution and chemical composition heterogeneity of oxidation products. FeS2 oxidized in acidic medium produces few solid components. Only Fe" sulfate is detected on oxidized pyrite surface. In carbonate medium, oxidation layer is more complex. Iron is mainly with an oxidation state (11) under siderite or Fe" sulfate form, and in a minor extent with an oxidation state (111) under FeOOH form and from the first oxidized pyrite layers where Fe"' may exists as illustrated by the comparison of Fe3p and Fe2p312 peaks. Sulfix oxidation induces intermediate species (polysulfides and sulfoxyanions as S2032- also evidenced in solution) indicating that oxidation occurs at solid state before dissolution.

422

Bonnissel-Gissinger, P., Alnot, M., Ehrhardt, 3.-J. & P. Behra 1998. Surface Oxidation of Pyrite as a Function of pH. Environ. Science and Technology 32: 2839-2845. Brion, D. 1980. Etude par specnoscopie de photoelectrons de la degradation supeficielle de FeS2, CuFeS2, ZnS et PbS a I'air et dans I'eau. Appl. of Surf: Sci. 5 : 133-152. Descostes, M., Mercier, F., Beaucaire, C. & P. Zuddas, in prep. Descostes, M., Mercier, F., Thromat, N., Beaucaire, C. & M. Gautier-Soyer 2000. Use of XPS to the determination of chemical environment and oxidation state of iron and sulfur samples: Constitution of a data basis in binding energies for Fe and S reference compounds and applications to the evidence of surface species of an oxidized pyrite in a carbonate medium. Appl, SurJ Science 165: 288-302. Evangelou, V.P.B. 1995. Pyrite oxidation and its control. CRC Press. Lowson, R.T. 1982. Aqueous oxidation of pyrite by molecular oxygen. Chemical Reviews 82: 46 1-497. Luther 111, G.W. 1987. Pyrite oxidation and reduction: Molecular orbital theory considerations. Geochim. et Cosm. Acta 51: 3193-3199. Mayer, J.W. & E. Rimini 1977. Ion beam handbook for material analysis. New York: academic press. Ponsot B., Salomon, J. & P. Walter 1998. RBS study of galena thermal oxidation in air with a 6-MeV I6O3+ion beam. Nucl. Inst. and Meth. in Phys. Res. B 136-138: 1074-1079. Pratesi, G. & C. Cipriani 2000. Selective depth analyses of the alteration products of bornite, chalcopyrite and pyrite performed by XPS, AES, RBS. Eur. J. of Mineralogy 12: 397-409. Sasaki, K.,. Tsunekawa, M., Tanaka, S. & H. Konno 1995. Confirmation of a sulfur-rich layer on pyrite after oxidative dissolution by Fe(II1) ions around pH 2. Geoch. et Cosm. Acta 59: 3155-3158.

Wafer-Rocklnteracfion2001, Cidu (ed.),0 2001 Swets & Zeitlinger,Lisse, ISBN 90 2651 824 2

Local structure of uranium (Vl) sorbed on clinoptilolite and montmorillonite R .J .Reeder State University of New York at Stony Brook, Stony Brook, New York, USA

M .Nugent Center for Nuclear Waste Regulatory Analyses, San Antonio, Texas, USA

R.T.Pabalan Center for Nuclear Waste Regulatory Analyses, San Antonio, Texas, USA

ABSTRACT: X-ray absorption fine structure spectroscopy was used to determine the structure and oxidation state of uranium species sorbed onto the clay montmorillonite and the zeolite clinoptilolite. Samples were prepared at solution pH -3 and -6 so the effect of pH on the sorption mechanism could be evaluated. The results demonstrate a difference in the equatorial coordination of the uranyl sorbate as a function of pH for both minerals. Split equatorial shells are evident for both samples at pH -6, whereas primarily a single shell exists at pH -3. The split equatorial shells probably indicate that discrete equatorial oxygens form chemical bonds at surface functional groups, as would be expected for an inner-sphere-type surface complex. In contrast, the single equatorial shell for samples at pH -3 suggests a more uniform bonding environment for the oxygens as would be expected for an outer-sphere-type complex. Such an environment is consistent with ion exchange at cationexchange sites of the sorbents. 1 INTRODUCTION Sorption is an important mechanism for attenuating the migration of radionuclides from nuclear waste repositories, such as the proposed geologc repository at Yucca Mountain, Nevada, to the accessible environment. Sorption experiments can provide mformation on radionuclide uptake as a function of the physicochemical characteristics of the mineral sorbent (e.g. composition, structure, surface area, surface charge) and the chemistry (e.g. pH, Eh, ionic strength, complexing ligands) of radionuclide-bearing water. However, these experiments give no direct information on the structure and local chemical environment of the sorbed species. In ths study, x-ray absorption fine structure (XAFS) spectroscopy was used to elucidate the structure and oxidation state of uranium species sorbed onto montmonllonite and clinoptilolite, two important minerals found at Yucca Mountain. Montmonllonite is a smectite clay with a 2: 1 layered structure, characterized by a layer of octahedrally coordinated aluminum atoms sandwiched between two layers of tetrahedrally coordinated silicon atoms. Choptilolite is a zeolite mineral characterized by open intracrystalline channels parallel to the c-axis that allow easy movement of some ions and molecules into and out of the structure. Both minerals contain two types of sorption sites: (i) permanently charged cation-exchange sites and (ii) variably charged surface hydroxyl groups. The former is due to isomorphic substitutions in the structure, e.g. w i t h the octahedral (Mg2+for A'+) or tetrahedral (Al" for Si4') layers of montmonllonite or w i t h the

zeolite framework (A3'' for Si"), causing a permanent negative charge that is compensated externally by cations interacting with the interlayer or intracrystalline exchange sites. The latter is due to partially coordinated aluminum and silicon exposed at the clay crystallite edges or zeolite surfaces that hydrolyze to form aluminol (AlOH) and silanol (SiOH) groups. These hydroxylated sites exhibit acidbase behavior and coordinative properties similar to those of oxide (e.g. Si02) surfaces. Uranium(V1) sorption onto mineral surfaces is a strong function of pH. As shown in Figure la, at a PCOZ(g) equal to 10-3.5atm, uranyl sorption on Namontmonllonite, Na-clinoptilolite, and quartz is highest at near-neutral pHs and decreases towards more acidic or more alkahe conditions (Pabalan et al. 1998). A comparison of the pH-dependence of uranyl sorption (Fig. la) with that of uranyl aqueous speciation (Fig. 1b) indicates that uranyl sorption is favored in the pH range where hydroxy-complexes are the predominant aqueous species. Uranyl sorption is low at alkaline pHs where uranyl-carbonate species are predominant. For minerals with no permanently charged cation-exchange sites such as quartz (Fig. la), uranyl sorption is also low at acidic pHs where aqueous uranyl exists primarily as the mononuclear, ionic species UOZ2+.However, uranyl sorption onto negatively-charged cation-exchange sites of montmorillonite and clinoptilolite from solutions of low pH and low ionic strength has been shown to be important (Zachara & McKinley 1993, Andreeva & Chemyavskaya 1982). The results plotted in Figure l a for Na-montmorillonite and Naclinoptilolite do not show much sorption at low pH

423

powder was not measured, but the value reported previously for the 100-200 mesh size (150-75 p) fraction is 10 mz/g (Pabalan 1994). The cation exchange capacity reported for the SAz-1 montmorillonite is 1.2 meq/g (Van Olphen & Fnpiat 1979) and 2.04 meq/g for the clinoptilolite (Pabalan 1994). Uranium(V1)-loaded samples for XAFS analysis were prepared by reacting 1 gram of mineral powder with 450 mL of uranyl nitrate solutions. For each mineral, samples were prepared at two values of solution pH (-3 and -6) so the effect of pH on the sorption mechanism could be evaluated. Uranyl concentrations of 9.2 x 10-’ and 4.6 x 10-5 M, respectively, were used at pH -3 and -6. After mixing the solid with the uranyl solution, the solution bottles were capped and placed on a gyratory shaker. I r was bubbled through the solutions to maintain atmospheric K O 2 . The solution pH was monitored and adjusted with HN03 or NaOH to maintain a pH of -3 or -6. After about two weeks, the mixtures were centrifkged and the moist pastes were loaded onto XAFS sample holders. The uranium concentrations before addition of solid and after sorption equilibrium was achieved were analyzed by inductively coupled plasma emission spectrometry to determine the uranyl uptake. The relative amount of uranium sorbed by clinoptilolite is 9 YOat pH-3 and 95 % at pH-6. The montrnorillonite sorbed 92 % at pH-3 and 82 % at pH-6.

Figure 1. Comparison of (a) uranyl sorption on Namontmorillonite, Na-clinoptilolite, and quartz and (b) uranyl aqueous speciation (CU, = 2 . 1 ~10-7 M; PCO, = 10”’ atm; 0.1 M NaNO, matrix) (from Pabalan et al. 1998).

2.2 X-ray absorptionfine structure spectroscopy

because the 0.1 M NaN03 background electrolyte of the experiments suppressed potential ion exchange between the uranyl ion in solution and the Na’ in the solid (Pabalan et al. 1998).

XAFS spectra were collected on the moist paste samples at beamline X11A of the National Synchrotron Light Source at Brookhaven National Laboratory. Multiple scans were taken over the uraniumL2(20948 eV) and L3-edges (17166 eV) using a pair of sacon( 111) monochromtor crystals, with one crystal detuned by -30% for harmonic rejection. Absorption spectra were obtained at room temperature using fluorescence detection with a 13-element solidstate germanium detector. The L3-edge spectra of the clinoptilolite @H-3) revealed fluorescence interferences from rubidium and strontium. Therefore, spectra subsequently were obtained at the L e d g e to avoid the interferences. Data analysis included subtraction of pre-edge background, nomahation, and conversion to k-space, followed by p ~fitting , using a cubic spline. The ~ ( k h)c t i o n was Fourier transformed using k3 weighting (typical k-range: 2.9-12.8 A-’).All fitting was done in R-space using WinXAS (Ressler 1997), and theoretical backscattering amplitudes and phase shlRs were calculated using FEFF7 (Zabinsky et al. 1995). Reference spectra from previous studies (Reeder et al. 2000) were used to assess the fit quality with the theoretical amplitudes and phases. Several starting models were used for the FEFF7 calculations, including soddyite [ (U02)2(Si04).2H,O] and sklodowskite [Mg(U02)2Si206(0H)2.5H20]. A single threshold enwas allowed to vary during fitting. ergy value (MO) Errors in the fit parameters, estimated from fits of

2 EXPERIMENTAL 2.1 Sample preparation The clinoptilolite powder used in the study was prepared from a sample of clmoptilolite-rich tuff (source locality: Death Valley Junction, California). The preparation method, described in Pabalan (1994), involved grinding and sieving to certain size ranges, heavy liquid separation of mineral impurities, dissolution of carbonate minerals with a sodium acetate buffer, dissolution of iron oxides with a dithionitecitrate-bicarbonate mixture, and exchange with NaCl solution to form homoionic Na-clinoptilolite. The choptilolite powder in the size range 200-270 mesh (75-53 pn) was used in the preparation of uranylloaded clinoptilolite. The montmorillonite was obtained in powder form (SAz-1; C a - f o p source locality: Apache County, Anzona) from the Source Clays Mineral Repository. The montmonllonite was also pretreated to remove carbonates and iron oxide minerals. The external surface area of the pretreated montmorillonite sample measured with an N2-BET Coulter SA3100 surface area analyzer is 109 m2/g. The surface area of the 200-270 mesh size choptilolite 424

Clinoptilolite pH 3.2

well-characterized model compounds, are -120 % for coordination number (CN) and -10.02 8, for first and second shell distances (R).Debye-Waller type factors ( G ~have ) approximate errors of -10.002-0.003 8,'.

0101

3 RESULTS

a~

(a)

-0.10

- Magn

I1

-0.05

The near-edge regions of the uranium L2-edge absorption spectra are shown in Figure 2a For all samples, the edge positions, relative to U(1V) and U(V1) reference standards, and the presence of a characteristic shoulder above the absorption maximum confirm the uranium remains in the 6+ oxidation state as the O=U=O moiety. No obvious difference in the near-edge structure is observed among the samples. Fourier transform (FT) magnitudes are shown in Figure 2b. The principal peak at 1.4 8, in all FT magnitudes (not corrected for phase shifts) corresponds to the two axial oxygens at -1.8 8,in the uranyl moiety. Peaks at -2 8, in the FT correspond to equatorial oxygen shells. For the clinoptilolite and montmorillonite reacted at pH -3, the FT magnitudes show a single peak well separated from the axial oxygen component. In contrast, both minerals at pH -6 show separate and weaker peaks, suggesting split equatorial oxygen coordination. No signlficant peaks are apparent at higher R, indicating no significant backscattering from silicon, aluminum, or other uranium atoms and unlikely formation of a uranyl-containing

Uranium L,-edge (20948eV)

A

U

i 0

1

(data)

_ _ _ Magn (fit) (data) _ - lmag lmag (fit)

I1 2

3

R

4

I

5

I 6

(4

Clinoptilolite pH 6.4 Uranium L,-edge (17166 eV)

Magn (data)

~

0

1

2

3

4

5

6

R (A)

Figure 3. R-space fits showing the real and imaginary parts for uranyl sorbed onto clinoptilolite at pH -3 and -6.

second phase. A weak peak at -3 8, in the FT reflects a multiple-scattering path within the U 0 2 moiety. For all fits, a multiple-scattering contribution at -3.58 A was fitted with the four-legged OaxI-U-OaxZU path. The real and imaginary components of the Rspace fits for clinoptiloiite are shown in Figure 3, and

Uranium L, absorption edges

Clinopt pH 3.2

Table 1. Best-fit X A F S parameters for uranyl sorbed on clinoptilolite and montmorillonite at different pHs Clinoptilolite pH 3.2 Shell U-0, U-O,,, U-O,,, MS'

0.0 20900

210w

211w

Photon energy (ev)

(b)

Fourier transforms

AEof

CNa R (A)b 2d 1.78 1.4 2.21 4.4 2.42 2.6 3.56 4.2 eV

Clinoptilolite pH 6.4 o2(A2)' 0.001 0.002 0.004 0.005

Montmorillonite pH 3.3 Shell IJ-0, U-O,,l

Clinopt pH 6 4

-

Mont p H 3 3

MS' MO

Mont pH6.3 0

1

2

3

4

5

6

a

R (A)

Figure 2. (a) Near-edge regions of uranium L,-edge absorption spectra and (b) Fourier transform magnitudes (,@-weighting; uncorrected for phase shifts) for uranyl sorbed on clinoptilolite and montmorillonite at different pH values.

425

CNa R (A)b 2 1.77 5.5 2.41 _ _ _ 2.6 3.54 4.5 eV

o2(A')' 0.001 0.007 -

0.003

Shell U-0, U-O,,, U-O,,, MS" AEo

CN R ( A ) 2 1.79 2.0 2.28 3.1 2.45 2.3 3.57 5.1 eV

oZ(A2) 0.001 0.003 0.005 0.004

Montmorillonite pH 6.3 Shell U-0, U-O,,, U-O,,, MS" AEo

CN R(A) 2 1.79 3.5 2.28 3.4 2.44 2.4 3.59 4.4 eV

02(Az) 0.002 0.010

0.010 0.002

Coordination numbers have errors off 20 % Errors on distance are f 0.02 A Errors on DebyeWaller type disorder parameters are f 0.002-0.003 A' Fixed at 2 for most fits, but refined to a value near two for test fits Four-legged multiple scattering path O,,-U-O,,-U Global energy threshold varied during fitting

the best-fit results for both minerals are reported in Table 1. All fits show two axial oxygens in the range 1.77-1.79 A. The uranyl sorbed onto montmorillonite at pH 3.3 shows a single equatorial shell of 5-6 oxygens at 2.41 A. At pH 6.3, the montmorillonite is best fitted with two equatorial shells (each with -3 oxygen atoms) at 2.28 and 2.44 A. Generally sirmlar fit results were found for uranyl sorbed on clinoptilolite. The sample of clinoptilolite reacted at pH 6.4 shows two equatorial oxygen shells at 2.28 and 2.45 A. For the clinoptilolite reacted at pH 3.2, the single peak in the FT suggested a single equatorial oxygen shell, and fitting yielded -5 oxygen atoms at 2.45 k However, the fit quality was not as good as for the montmorillonite at pH 3 . 3 . The fit was improved by an additional, weak equatorial oxygen component, giving best fit results of -1 oxygen at 2.21 A and -4 oxygens at 2.42 A. The significance of the weak contribution at 2.21 8, needs additional clardication through additional experiments. Attempts to fit U-& and U-Si paths generally resulted in very small CN values that were significantly less than the estimated errors. The lack of uranium backscattering also suggests that sorbed uranyl species are mononuclear. 4 CONCLUSIONS

The results demonstrate a difference in the equatorial coordination of the uranyl sorbate as a h c t i o n of pH for both montmorillonite and clinoptilolite. Split equatorial shells are evident for both samples at pH -6, whereas primarily a single shell exists at pH -3. The results for montmorillonite are consistent with the EXAFS data reported by Sylwester et al. (2000). The difference in U-0 distances in the equatorial oxygen shell for the samples at pH -6 must reflect a difference in the bonding of these oxygens with the sorption sites of the minerals. Hence, the split equatorial shells probably indicate that discrete equatorial oxygens form chemical bonds at surface hctional groups, as would be expected for an inner-spheretype surface complex. In contrast, the single equatorial shell for samples at pH -3 suggests a more uniform bondmg environment for the oxygens as would be expected for an outer-sphere-type complex. Such an environment is consistent with ion exchange at cation-exchange sites of the sorbents. However, the experiments do not provide direct evidence that the uranyl ion is located in the interlayer exchange sites of montmorillonite or in the intracrystahe channels of choptilolite. The alternate X A F S fit indicating equatorial splitting for choptilolite reacted at pH -3 is consistent with distortion of the equatorial shell due to steric limitations imposed by the zeolite structure. Further investigations are needed to refine the interpretation of the sorption mechanisms.

ACKNOWLEDGMENTS This work was f h d e d by the National Science Foundation grant EAR9706012 (R. Reeder) and by the U.S. Nuclear Regulatory Commission (NRC) under Contract Number NRC-0297-009 (M. Nugent and R. Pabalan). This paper does not necessarily reflect the views or regulatory position of the NRC.

REFERENCES Andreeva, N.R & N.B. Chernyavskaya 1982. Uranyl ion sorption by mordenite and clinoptilolite. Radiokhimiya 24:9-13. Pabalan, RT. 1994. Thermodynamics of ion-exchange between clinoptilolite and aqueous solutions of Na"/K' and Na+/Ca2+. Geochim. Cosmmhim.Acta 58: 4573-4590. Pabalan, R.T., D.R. Turner, F.P. Bertetti & J.D. Prikryl 1998. Uranium(VI) sorption onto selected mineral surfaces: Key geochemical parameters, In E. Jenne (ed.), Adsorption of Metals by Geomedia: 99-130. San Diego, California: Academic Press. Reeder, R.J., M. Nugent, G.M. Lamble, C.D. Tait & D.E. Morris 2000. Uranyl incorporation into calcite and aragonite: XAFS and luminescence studies. Environ. Sci. Technol. 34: 638-644. Ressler, T. 1997. WinXAS: A new software package not only for the analysis of energy-dispersive XAS data. J . Physique IV 7: C2-269. Sylwester, E.R., E.A. Hudson & P.G. Allen 2000. The structure of uranium(VI) sorption complexes on silica, alumina, and montmorillonite. Geochim. Cosmochim. Acta 64, 243 1-2438. Van Olphen, H. & J.J. Fripiat (eds.) 1979. Data Handbookfor C l q Minerals and Other Non-Metallic Materials. Oxford: Pergamon Press. Zabinsky, S.I., J.J. Rehr, A. Ankudinov, R.C. Albers & M.J. Eller 1995. Multiple-scattering calculations of X-ray absorption spectra. Phys. Rev. B 52: 2995-3009. Zachara, J.M. & J.P. McKinley 1993. Influence of hydrolysis on the sorption of metal cations by smectites: Importance of edge coordination reactions. Aquatic Sci. 55: 250-161.

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Surface composition of enargite(CU3AsS4) A .Rossi. D .Atzei & B .Elsener Dipartimeizto di Chiinica Irzorganica ed Analitica, Universita’ di Cagliari, Italy

S.Da Pelo, F.Frau & P.Lattanzi Dipartimento di Scienze della Terra, Universita’ di Cagliari, Italy

P.L.Wincott & D.J.Vaughan Department of Earth Sciences, The University, Manchester, UK

ABSTRACT: Freshly ground powders and crystals of natural enargite samples from a Sardinian mine (Furtei) were analysed by XPS at liquid nitrogen temperature. Binding energies of copper Cu2p3/2, arsenic As3d5/2 and sulfur S2p at 932.4 f 0.2 eV, 43.9 0.1 eV and 162.2 f 0.2 eV, respectively, were found. On ground samples (powder), a second component of the S2p peak at 163.8 f 0.2 eV was found. This could be attributed to either elemental sulfur or to a polysulfide. Quantitative evaluation of the XPS intensities showed a composition of 47 wt. % copper, 19 wt. % arsenic and 34 wt. % sulfur, in very good agreement with the bulk electron microprobe analysis. By contrast, on “as received” crystals of natural enargite, that had been exposed for a long time to the atmosphere, an oxidised layer of 0.5 nm thickness enriched in arsenic was found. This outer layer is very likely to influence the interaction of enargite with the natural environment and its behaviour in mineral processing plants, and should be taken into account in the assessment of the potential impact of enargite-bearing ores on the environment.

*

1 INTRODUCTION Enargite is a sulfide mineral containing copper and arsenic. It is present in some precious metal deposits, including the Furtei mine in Sardinia (Italy). The extraction of gold from these ores involves exposure to the exogenous environment of significant amounts of enargite. This is of concern because of the increasing demands for novel, more environmentally-friendly methods of mining and metal extraction. The outer layers present on mineral surfaces determine the reactivity of these minerals in contact with water, aqueous solutions of inorganic or organic species, and bacteria (McCarron et al. 1990, England et al. 1999). It is acknowledged that the characterisation of mineral surfaces can be accomplished best with surface analytical techniques, such as X-ray Photoelectron Spectroscopy (XPS), which allows the determination of the chemical state of the elements present on the surface. This information is important to establish the changes in the chemical state of the elements due to surface reaction. Significant progress has been made in monitoring changes in the chemical state of surface elements, also thanks to the availability of curve-fitting routines for processing the spectra. Less has been published on the changes in composition of the outer layers present on mineral surfaces, despite a big demand for the results of

quantitative analysis. Most of the papers published in the literature report the results of average surface composition, that is calculated assuming that the analysed sample has a homogeneous chemical profile from surface towards bulk. This might be correct for the surfaces of minerals that have been fractured in vacuum, but as soon as a surface has been in contact with the atmosphere or a solution, it reacts chemically, and very often the resulting surface layer not only has a different chemical state, but also a different composition with respect to the bulk material. To calculate the composition of the reaction layer and of the bulk material, it is thus mandatory to work under the assumption that a layered structure is present on the surface. In this work XPS qualitative and quantitative results obtained on “as received” enargite samples from a Sardinian mine (Furtei) are presented. These results are the basis for ongoing experiments on enargite oxidation in an acid environment. 2 EXPERIMENTAL Samples of natural enargite collected at the Furtei mine in Sardinia (labelled ENFurtei) were studied. They were characterised by X-ray powder diffraction and electron probe microanalysis before XPS analysis. Prior to introduction in the spectrometer, 427

the crystals or powders obtained after grinding in air were mounted on a nickel stub. XPS analyses were performed at the Laboratory for Surface Analysis at the University of Cagliari on an ESCALAB200 spectrometer using AlKa and MgKa sources run at 15 keV and 20 mA. Calibration of the spectrometer has been accomplished according to Seah (1989). Charge shift was corrected by using the C l s peak position at 285.00 eV. All XPS measurements were carried out at liquid nitrogen temperature. 3 RESULTS AND DISCUSSION 3.1 Qualitative analysis

Figure 2. High-resolution spectrum of As3d region of "as received" enargite sample ENFurtei (powder on scotch).

The XPS survey scan of natural enargite is shown in Figure 1. Only signals due to copper, arsenic and sulfur are detectable together with carbon and oxygen signals. No signals attributable to other elements were detected, indicating that minerals other than enargite are not present.

synthetic enargite samples studied. Only a very small component at 45.5 eV was resolved. This signal might suggest the presence of a small amount of arsenic bound to oxygen in the outermost part of the surface (Fullston et al. 1999). Copper shows a Cu2p doublet at 932.4 f 0.2 eV and at 952.6 f 0.2 eV (Cu2p3/2 and Cu2p1/2). The energy difference between Cu2p3/2 and Cu2p1/2 is thus 19.8 eV. The Cu2p signals (Fig. 3) were fitted with one single Gaussian-Lorentzian curve for the whole spectrum. The full width at half maximum (FWHM) for the Cu2p3/2 signal was 1.6 f 0.1 eV, and the binding energy value found is in agreement with that reported in the literature for monovalent copper in enargite and luzonite (Velasquez et al. 2000). In the case of this natural enargite sample, no satellite structure, which would be typical for the presence of copper in a higher oxidation state on the surface, was detected. The most intense sulfur line, S2p, is asymmetric due to spin - orbit splitting. This fact was taken into account when processing the spectrum (Fig. 4). Specifically, the energy separation S2p3/2 - S2p1/2 was held constant at 1.2 eV and the relative intensities of each doublet peak were fixed to be

3.2 Chemical state analysis The spectra were processed after background subtraction (Sherwood 1983) by fitting of GaussianLorentzian curves. The curve-fitting parameters were obtained using, as a model compound, synthetic enargite (Rossi et al. 200 l), investigated with the same spectrometer under the same analysis conditions. The spectra of arsenic and s u l k revealed multiple chemical states at the surface of the enargite samples. The representative arsenic and sulfur high-resolution spectra are shown in Figures 2 and 4. Arsenic shows the most intense XPS line, As3d, at 43.9 f 0.1 eV in the freshly ground samples. The As3d spectrum was fitted with an As3d5/2 and As3d312 doublet, with a binding energy difference of 0.7 eV and a ratio of 3:2. The shape of the signals (Fig. 2) did not change for natural or

Figure 1. XPS survey scan of a natural enargite sample (powder on scotch). Elements detected are copper, sulfur, arsenic, carbon and oxygen.

Figure 3. High-resolution spectrum of Cu2p region of L'as received" enargite sample ENFurtei (powder on scotch).

428

Table 1. Results of the quantitative analysis (wt. %) of enargite samples from the Furtei mine, Sardinia, obtained by XPS surface analysis. Measurements on three independent samples. Element

ENFurtei powder

Figure 4: High-resolution S2p spectrum of “as received” enargite sample ENFurtei (powder on scotch).

ENFurtei

ENFurtei

crystal

crystal

oxidised layer

bulk* 46*2

cu

472 1

35* 1

As

19* 1

49k 1

19*2

S 34h 1 16* 1 * composition beneath the oxidised layer

355 1

On the other hand, crystals of ENFurtei exposed to the atmosphere, show, beneath the contamination layer, an oxidised layer of a thickness of ca. 0.5 nm. The composition of this oxidised layer is enriched in arsenic, while the composition of enargite beneath the oxidised layer was found to be nearly identical to that of powdered enargite (Table 1). These results confirm the findings of another paper (Rossi et al. 200 1), where synthetic enargite powders and natural enargite crystals “as received” were analysed under the same experimental conditions. Synthetic enargite powders showed qualitative and quantitative results in agreement with the bulk formula, whereas on enargite crystal surface an As-enriched oxidized layer was observed. Thus, a clear difference between freshly ground natural and synthetic powders of enargite and the “as received” enargite crystals is found. This might indicate that a short time exposure to the air does not alter the surface composition, whereas an alteration does occur after exposition to the atmosphere for long period of time.

equal to the ratio of their respective degeneracy (2: 1). As shown in Figure 4, two sulfur S2p signals were revealed in the detailed sulfur spectrum. The more intense signal at 162.2 f 0.2 eV was assigned to sulfur in the sulfide chemical state (formal oxidation state -2). The weak signal at 163.8 f 0.2 eV can be interpreted as originating only from the surface (or near surface) of the sample, where the electronic state of the sulfur atoms is different from the bulk. In the study of oxidised sulfide mineral surfaces, the signal at about 163.8 eV has been variously interpreted as either elemental sulfur or polysulfide (Cordova et al. 1997, Nesbitt et al. 1995, England et al. 1999, Pratesi & Cipriani 2000; and references therein). 3.3 Quantitative analysis Based on the integrated intensities determined from curve fitting of the As3d, Cu2p3/2 and S2p signals, the surface chemical compositions of the different enargite samples from ENFurtei were determined with a “three layer model” adapted for enargite samples from earlier work on amorphous metals (Rossi & Elsener 1992) and stainless steels (Elsener & Rossi 1995). The peak areas were corrected with the photoionization cross sections from Scofield (1976), for the asymmetry function (Reilman et al. 1976), for attenuation length (Seah & Dench 1979), and for the transmission function Q(E) (Seah 1993). Calculations were performed taking into account the attenuation of the photoelectron signals due to the presence of a contamination layer. This arises from exposure to the natural environment. Its thickness is calculated between 1 and 1.5 nm. The results of the quantitative XPS analysis of the freshly powdered enargite samples of ENFurtei (Table 1) are in very good agreement with the bulk analysis obtained from the electron microprobe (Cu 48.9 wt. %, AS 18.3 wt. %, Sb 0.5 wt. %, S 32.8 wt. %; Rossi et al. 2001), corresponding to a formula of Cu3,olAso.gsSbo.o2S4(calculated on basis of S = 4).

4 CONCLUSIONS From the results of this surface chemical investigation on natural enargite samples from the Furtei mine, Sardinia, it can be concluded that: 1) The surface composition determined by XPS analysis of natural enargite samples, analysed as freshly ground powders, is in good agreement with the bulk analysis obtained by electron microprobe. Thus, a short time exposure to the air does not alter the surface composition. 2) The high-resolution spectrum of sulfur clearly indicates the presence of a sulfur compound at higher binding energies (163.8 eV) with respect to bulk sulfur. This species might be assigned either to elemental sulfur or polysulfide present at the surface, or alternatively by atomic rearrangement of the surface upon grinding of the crystals. 3) By contrast, data obtained for enargite crystals exposed for long time to the atmosphere indicate the 429

presence of a thin surface layer enriched in arsenic. This oxidised layer is likely to affect the interaction of enargite with the environment (and the behaviour of enargite in mineral processing operations), and should be therefore taken into account in the assessment of the potential impact of enargite-bearing ores on the environment. ACKNOWLEDGEMENTS

Seah, M.P. & W.A. Dench 1979. Quantitative electron spectroscopy of surfaces: Standard data base for electron inelastic mean free path in solids. Surface Interface Analysis 1 : 2-1 1 . Sherwood, P.M.A. 1983. Appendix 3. In D. Briggs & M.P. Seah (eds.), Practical Surface Analysis: 445-475. Chichester: Wiley. Velasquez, P., Ramos-Barrado, J.R., Cordova, R. & D.Leinen 2000. XPS analysis of electrochemically modified electrode surface of natural enargite. Surface Interface Analysis 30: 149-153.

The financial support of Italian MURST ( e x 4 0 % grants to L. Fanfani and A. Rossi, and ex-60 YOto P. Lattanzi) is gratefully acknowledged. The work in Manchester was supported by NERC.

REFERENCES Cordova, R., Gomez, H., Real, S.G., Schrebler, R. & J.R. Vilche 1997. Characterization of natural enargite/aqueous solution systems by electrochemical techniques. J. Electrochem. Soc. 44 (8): 2628-2636. Elsener, B. & A. Rossi 1995. Effect of pH on electrochemical behaviour and passive film composition of stainless steels. Materials Science Forum 192-194: 225-236. England, K.E.R., Pattrick, R.A.D. & D.J. Vaughan 1999. Surface oxidation studies of chalcopyrite and pyrite by glancing angle X-ray absorption spectroscopy (REFLEXAFS). Mineral. Mag. 63: 559-566. Fullston, D., Fornasiero, D. & J. Ralston 1999. Oxidation of synthetic and natural samples of enargite and tennantite: 2. X-ray Photoelectron Spectroscopic Study. Langmuir 15: 4530-4536. McCarron, J.J., Walker, G.W. & A.N. Buckley 1990. An X-ray photoelectron spectroscopic investigation of chalcopyrite and pyrite surfaces after conditioning in sodium sulfide solutions. Internationai Journal Of Mineral Processing 30 (1-2): 1-16. Nesbitt, H.W., Muir, I.J. & R. Pratt 1995. Oxidation of arsenopyrite by air and air-desaturated, distilled water and implications for mechanism of oxidation. Geochim. Cosmochim. Acta 59: 1773-1786. Pratesi, G. & C. Cipriani 2000. Selective depth analyses of the alteration products of bornite, chalcopyrite and pyrite performed by XPS, AES, RBS. Eur. J. Mineral. 12: 397410. Reilman, R.F., Msezane, A. & S.T. Manson !976. Relative intensities in photoelectron spectroscopy of atoms and molecules. J. Electron Spectroscopy 8: 389-394. Rossi, A. & B. Elsener 1992. XPS analysis of passive films on the amorphous alloy Fe70CrIOP13C7: Effect of the applied potential. Surface Interface Analysis 18: 499-504. Rossi, A., Atzei, D., Da Pelo, S., Frau, F., Lattanzi, P., England, K.E.R. & D.J. Vaughan 2001. A quantitative Xray photoelectron spectroscopy study of enargite (Cu,AsS,) surface. Surface Interface Analysis (accepted). Scofield, J.H. 1976. Hartree-Slater subshell photoionisation cross-sections at 1254 and 1487 eV. J of Electron Spectroscopy and Related Phenomena 8: 129-137. Seah, M.P. 1989. Reference binding energies. Surface Interface Analysis 14: 488. Seah, M.P. 1993. XPS reference procedure for the accurate intensity calibration of electron spectrometers. Results of a BCR intercomparison co-sponsored by the VAMAS SCA TWA. Surface Interface Analysis 20: 243-266.

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Orthoclase surface structure and dissolution measured in situ by X-ray reflectivity and atomic force microscopy N .C.S turchio Earth and Environmental Sciences, University of Illinois, Chicago, Illinois 60607 USA

P-Fenter & L.Cheng Environmental Research Division, Argonne National Laboratory, Argonne, Illinois 60439 USA

H .Teng Earth and Environmental Sciences, George Washington University, Washington, D.C. 20054 USA

ABSTRACT. Orthoclase (001) surface topography and interface structure were measured during dissolution by using in situ atomic force microscopy (AFM) and synchrotron X-ray reflectivity at pH 1.1 - 12.9 and T = 25 - 84°C. Terrace roughening at low pH and step motion at high pH were the main phenomena observed, and dissolution rates were measured precisely. Contrasting dissolution mechanisms are inferred for low- and high-pH conditions. These observations clarify differences in alkali feldspar dissolution mechanisms as a function of pH, demonstrate a new in situ method for measuring face-specific dissolution rates on single crystals, and improve the fundamental basis for understanding alkali feldspar weathering processes.

1 INTRODUCTION

2 EXPERIMENTAL METHODS

Feldspar weathering influences global cycling of Si, Al, and alkali and alkaline earth metals; atmospheric CO2 concentration; natural water composition; and soil formation. Weathering occurs naturally when feldspars dissolve in waters having pH values of 5 to 7. Few direct observations of dissolving feldspar surfaces (Jordan et al. 1999) have been made at any pH, however, and there is no clear consensus regarding the associated molecular-scale mechanisms (Blum & Stillings 1995). Available evidence suggests that there may be distinct dissolution mechanisms at low and high pH, based on the pH dependence of dissolution kinetics. Observations by numerous workers indicate that feldspar dissolution is generally incongruent at low pH (Casey et al. 1988, Chou & Wollast, 1984, Hellmann et al. 1996, Nugent et al. 1998) and congruent at high pH (Casey et al. 1988, Hellmann et al. 1996). Significant uncertainty persists, however, regarding the microscopic structure of the feldspar surface during dissolution. We explored the use of in situ AFM and X-ray reflectivity techniques to investigate the dissolution mechanisms, dissolution rates, and structures of dissolving orthoclase-water interfaces (Teng et al. 2001). This work has provided new insight regarding alkali feldspar dissolution mechanisms as a function of pH.

We examined the evolution of dissolving orthoclase (001) cleavage surfaces on a gem-quality crystal [Or94.5Ab4.5, monoclinic, Itrongay, Madagascar (Kimata et al. 1996)l. Surface topography was imaged at room temperature (23 k 2 "C) in 0.1 M HCL, deionized water, and 0.1 M NaOH having pH values (at 25 "C) of 1.1, 5.6, and 12.9. Measurements were made in TappingMode@using an in situ fluid cell AFM (Digital, Nanoscope IIIa MultiMode) with solution flow rates ranging from 3 to 200 mL h-'. Crystals in the fluid cell were imaged briefly once every 8-15 h during dissolution experiments to eliminate AFM tip-induced changes at the surface. Sample preparation for X-ray experiments followed the same procedure as that for AFM experiments. X-ray data were collected in situ from 25 "C to 84 "C in flowing solutions of 0.1 M HCI, 0.1 M oxalic acid (H2C2O4), and 0.1 M NaOH having pH values (at 25 "C) of 1.1, 1.3, and 12.9, respectively. Flow rates in the X-ray experiments varied from 3 mL h-' to 650 mL h-*. Synchrotron X-ray reflectivity measurements were made at the Advanced Photon Source (beamlines 12-ID and 12-BM) using monochromatic X-rays with photon energies ranging from 17.5 to 19.6 keV. Time-resolved measurements of dissolution kinetics and mechanisms were performed in a flow-through TeflonTMsample cell in transmission geometry through -3.5 mm of water 431

and two 0.13 mm KaptonTMwindows. The incident and reflected fluxes were measured with an ion chamber and a NaI scintillation detector, respectively. More details about our reflectivity measurement procedures are given elsewhere (Fenter et al. 2000). Samples were loaded into the cell in de-ionized water (DIW) and sample alignment was performed at room temperature. The cell was then externally heated to the desired temperature, taking advantage of the much slower dissolution rate at pH near 6, as compared with those at pH 1 and 13. To change the pH, we rapidly flushed the cell with 7 ml of solution (about 5 cell volumes) resulting in a brief (-2minute) decrease of the fluid temperature which was monitored continuously by a thermocouple positioned 2 mm above the sample surface. We then continued to flow solution through the cell at a rate of 3 mL h i ’ throughout the experiment using a syringe pump.

3 RESULTS 3.1 Atomic force microscopy

Dissolving orthoclase (001) cleavage surfaces at pH = 1.1 and 12.9 were imaged using AFM. The initial cleavage surfaces exhibited atomically flat terraces separated by steps (Figs la, 2a) haying heights of small integer multiples of 6.3 k 0.3 A, which corresponds to the (001) cell parameter, 6.475 A. At pH 1.1 and flow rate = 3 mL h-’, the surface roughened within 15 h, after which a continuous surface coating on terraces formed through gradual accumulation of secondary material (Fig. lb). This gel-like coating appeared to be soft and viscous, tended to stick to the AFM tip, and could be removed readily

either by scanning the AFM tip in contact mode or by increasing the solution flow rate. The surface exposed after removing a 2-nm-thick coating that formed in 120 h showed extensive development of well-defined nanopores on terraces. Negligible step motion was observed under these conditions. At pH 1.1 and flow rates of > 150 mL h-’, no surface coating developed on terraces. AFM images showed apparently complete removal of dissolved material, leading to the development of a rough surface with nanopores (Figs. lc, Id), resembling that exposed after removal of the surface coating developed under slow-flow conditions. This observation indicates that the formation of these gel-like surface coatings, that were identified in previous ex situ (Casey et al. 1988, Nugent et al. 1998) and in situ (Jordan et al. 1999) studies as “leached layers”, is primarily the result of solution mass transport kinetics. Thus, the transient non-stoichiometric behavior observed in powder dissolution studies at low pH does not represent an intrinsically nonstoichiometric acidic dissolution reaction, but results from growth of a surface coating due to the differential solubilities and/or dissolutiodprecipitation rates of the relatively insoluble reaction products. Unlike calcite, the mechanism controlling orthoclase dissolution at low pH affects terrace sites without preference for steps or other defects. At pH 12.9, step motion defined by the retreat of steps separating existing (001) terraces or the nucleation and spreading of etch pits (Fig. 2b, 2d) was the only apparent dissolution mode. Terraces remained intact and uncoated throughout the experiment (up to 2000 h) regardless of flow rate. Unitcell-high steps, each a bilayer consisting of two tetrahedral sheets, split into separate half-cell steps (Fig. 2c) as dissolution proceeded; this was seen for etch-pit formation as well (Fig. 2d). The step-

Figure 1. AFM images of (A) freshly-cleaved orthoclase (001) surface in deionized water and (B-D) after contact with pH 1.1 solution. The cleavage surface (A) exhibits atomically flat terraces and a monolayer step. Interaction with low-pH solution led to the development of a surface coating (B) at low flow rate (3 mL h-I). This coating washed away when the flow rate was increased to 150 mL h-’ revealing surface nanopores (C) and (D). Images (A), (C), and (D) were taken at the same location and show negligible step motion after 96 h.

432

Figure 2. AFM images of (A) freshly-cleaved orthoclase (001) surface in deionized water and (B-D) after exposure to pH 12.9 solution. The cleavage surface (A) exhibited a screw dislocation, atomically flat terraces, and both mono- and multi-layer steps. Step retreat (B) was the only observable dissolution mode. Split steps (C) revealed the bilayer structure of the unit cell along [OOl]. Coalescence of single- and halfunit-cell-depth etch pits after extensive dissolution (D) resulted in a patchy surface. splitting phenomenon, coupled with the predominance of step motion under these conditions, indicates that under-coordinated sites at steps and etchpit walls control orthoclase dissolution at high pH.

under the two pH conditions. At pH 12.9, the reflectivity decreased by a factor of -5 before returning to the initial value corresponding to a freshly-cleaved surface. The recovery of reflectivity after one monolayer dissolution therefore indicates that the termination of the dissolving orthoclase surface is essentially unchanged from the freshly-cleaved orthoclase-water interface. The substantial decrease in reflectivity before recovery implies that dissolution involves nucleation and growth of unit cell-depth etch pits as well as step motion. Thus, reaction at high pH is limited to un-

3.2 X-ray reflectivity In situ synchrotron X-ray reflectivity measurements (Fig. 3) of orthoclase dissolution were made at 50 "C to 84 "C. Surface dissolution processes can be characterized by time-resolved measurements of X-ray reflectivity at the "anti-Bragg" condition [Q = (47dh)sin(8) = ddool = 0.48 A-', where 0 is the angle of the incident X-ray beam]. This scattering condition maximizes surface sensitivity and results in rapidly decreasing reflectivity for any increase in surface roughness. Measured X-ray reflectivity versus time (Fig. 3) did not decrease monotonically during dissolution, as would be characteristic of random dissolution (e.g., where all exposed tetrahedral sites dissolve at the same rate). Instead, the reflectivity exhibited an oscillatory pattern for both low and high pH, implying that lateral dissolution processes are involved in both pH regimes. The dissolution rate can be estimated from the oscillation period. For 1.1 HCl at 52 "C pH and pH 12.9 NaOH at 50 "C, calculated rates were 4.0 x 10-l' and 1.5 x 10mol KAlSi08 m-2sec*1,respectively. These rates are consistent with the range of dissolution rates reported previously for powder dissolution under steady-state conditions (Blum and Stillings 1995). Thus, x-ray reflectivity provides a new method for measuring precise mineral dissolution kinetics that is statistically averaged over the macroscopic surface area corresponding to the footprint of the Xray beam. This approach also avoids the heterogeneity inherent to powder dissolution studies. The plots of reflectivity versus time (Fig. 3) are distinct

''

Figure 3. : In situ X-ray reflectivity measurements of dissolving orthoclase (001) surfaces. X-ray reflectivity versus time (measured at Q = 0.48 A-') during exposure to (A) pH 12.9 NaOH at 52 "C (triangles), (B) pH 1.1 HCl at 50°C (circles), and (C) pH 1.3 oxalic acid at 50 "C (squares). The X-ray reflectivity R(t) is normalized to the reflectivity of the freshly-cleaved surface R(0). The removal of each monolayer (ML) is noted.

433

areas, a distinguishing characteristic of steps and other under-coordinated sites is the presence of adjacent NBO-bearing T sites where both T1 and T2 sites attached to NBOs are exposed. The predicted weakening of T-O-T linkages at high pH due to the deprotonation of adjacent NBOs at steps is supported by our observations of step motion and step splitting by AFM (Fig. 2c) and the dominance of lateral dissolution processes under these conditions using X-ray reflectivity (Fig. 3) that prove the high reactivity of edge sites under highpH conditions.

der-coordinated sites such as steps (Fig. 2b-d) and involves fully congruent layer-by-layer dissolution, consistent with the AFM images. In contrast, the variation of reflectivity versus time at pH 1.1 exhibits a strongly damped oscillatory pattern. The same pattern was found for all flow rates, as high as 650 mL h-', confirming that the dissolution process at low pH is independent of the presence of surface coatings (Fig. lb) (Casey et al. 1988). The same behavior was found with 0.1 M oxalic acid (pH 1.3) at 50 "C; the 30% slower dissolution rate is attributable solely to differences in pH and temperature. Thus, oxalate has minimal influence on the dissolution process andor kinetics, consistent with previous studies (Blum and Stillings 1995). The overall decrease in x-ray reflectivity accompanying dissolution at low pH (Fig. 3b) implies that the dissolving surface (i.e., its roughness andor termination) is substantially modified from the freshly-cleaved surface under these conditions.

ACKNOWLEDGEMENTS Supported by U. S. DOE, Geosciences Research Program, Office of Basic Energy Sciences, under contract W-3 1-109-ENG-38 to Argonne National Laboratory.

REFERENCES CONCLUSIONS Blum A. E. & L. Stillings 1995. Feldspar dissolution kinetics. Rev. Mineral. 3 1: 29 1-35 1. Casey W. H., H. Westrich & G. Arnold 1988. Surface chemistry of labradorite feldspar reacted with aqueous solution at pH 2,3, and 12. Geochim. Cosmochim. Acta 52: 2795-2807. Chou L. & R. Wollast R. 1984. Study of the weathering of albite at room temperature and pressure with a fluidized bed reactor. Geochim. Cosmochim. Acta 48: 2205-2218. Fenter P., H. Teng, P. Geissbuhler, J. Hanchar, K. Nagy, & N.C. Sturchio 2000. Atomic-scale structure of the orthoclase (001)-water interface measured with highresolution X-ray reflectivity. Geochim. Costnochim. Acta 64: 3663-3673. Hellmann R., C. Eggleston, M. Hochella & D. Crerar 1996. The formation of leached layers on albite surfaces during dissolution under hydrothermal conditions. Geochim. Cosmochim. Acra 54: 1267- 1281. Jordan G., S. Higgins, C. Eggleston, S. Swapp, D. Janney & K. Knauss 1999. Acidic dissolution of plagioclase: In situ observations by hydrothermal atomic force microscopy. Geochim. Cosmochim. Acta 63: 3183-3191. Kimata M., S. Saito, M. Shimizu, I. Iida & M. Tomoaki 1996. Low-temperature crystal structures of orthoclase and sanidine. N. Jb. Miner. Abh. 171: 199-213. Nugent M. A., S. L. Brantley, C. G. Pantano, & P. A. Maurice 1998. The influence of natural mineral coatings on feldspar weathering. Nature 395: 588-591. Teng, H. , P. Fenter, L. Cheng & N.C. Sturchio, 2001. Resolving orthoclase dissolution mechanisms. Geochim. Cosmochim. Acta in press.

These new results demonstrate that orthoclase dissolution is controlled by at least two separate surface reaction mechanisms having distinct reactive sites. The dominant mechanism at low pH is active across the entire surface resulting in etch pit formation and roughening of terrace areas, whereas the dominant mechanism at high pH is active primarily at steps and other defects leaving the intrinsic feldspar-water interface essentially unchanged. The observations can be understood mechanistically in terms of the structure of the orthoclase surface, which consists of bridging oxygen (BO) and nonbridging oxygen (NBO) attached to tetrahedral framework AI and Si cations (T) in interconnected four-membered tetrahedral rings. Of the two distinct types of T sites, only one (Tl) is attached to an NBO at the (001) surface. Protonation of surface oxygens creates surface silanol (Si-OH) and aluminol (AI-OH) sites from NBOs and hydroxyl (T-OH-T) sites from BOs. Previous results of nuclear magnetic resonance and theoretical studies, suggest that feldspar dissolution at low pH is promoted by the higher proton affinity of A1-OH bonds than Si-OH bonds. This mechanism explains the observed reactivity and roughening because these sites are present on terraces of freshlycleaved orthoclase surfaces (Fenter et al. 2000). At high pH, BO linkages between adjacent T sites attached to NBOs are weakened when Si-OH and A1-OH sites are deprotonated. While no two NBOs are found on adjacent T sites within terrace 434

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Disseminated calcite in a global suite of granitic rocks: Correlations with experimental solutes A.F.White, M.S .Schulz, D.V.Vivit & T.D.Bullen U S . Geological Survey, Menlo Park, CA 94025

ABSTRACT: The rapid dissolution of trace amounts of disseminated calcite in crystalline rocks has important implications in global fluxes of Ca, Sr and inorganic carbon. To determine the range of calcite contents in granitoids, a world wide suite of 100 rock samples are analyzed. Detectable CO, is found in all the samples corresponding to a calcite concentration range of 10' to 104ppm. These results confirm the ubiquitous presence of calcite in granitic rocks. No correlation was found between the amount of calcite and major element compositions. Solutes produced from closed-system weathering of the rock samples were dominated by calcite dissolution as evidenced by a 2 to l correlation between Ca and alkalinity. Rocks containing elevated calcite reached thermodynamic saturation, which controlled corresponding solute concentrations. A comparison of s7Sr86Sr ratios indicates that calcite is slightly more radiogenic than plagioclase, suggesting calcite formation during late-stage open-system magmatic cooling. The impact of calcite in natural granitic weathering is dependent both on the amount of calcite initially present and the intensity of weathering which subsequently depletes calcite.

1 INTRODUCTION The release of Ca and the consumption of CO, during the weathering of silicate rocks have a number of important geochemical implications. The consumption of hydrogen ions during hydrolysis is the principal buffering mechanism associated with acid precipitation in crystalline silicate watersheds. Such acidification, in addition to deforestation, has led to the depletion of base cations in some soils to the extent that Ca has become a limiting nutrient (Huntington et al., 2000). Chemical weathering of Ca-silicate minerals is also recognized as controlling long-term climate change by providing a feedback through CO, drawdown. In contrast, weathering of carbonate rocks does not have a corresponding impact because all the CO, consumed during weathering is reintroduced back into the atmosphere by the relatively rapid precipitation of carbonates in the oceans (Berner & Berner 1997). Although the bulk of Ca in granitic rocks resides in plagioclase feldspars, recent studies have shown that for specific watersheds, significant amounts of both dissolved Ca and carbon (DIC) are derived from the weathering of small amounts of calcite (Mast et al. 1990). White et a.1 (1999) experimentally investigated the weathering of several granitic rocks taken from well-characterized watersheds and demonstrated that effluents were initially dominated by calcite dissolution and were gradually superseded

over a two year reaction interval by silicate hydrolysis reactions. The extent to which relatively rapid calcite dissolution dominated Ca and DIC fluxes in watersheds is dependent both on the relative amounts of calcite disseminated in the granitic rocks and the geomorphic age of the watershed regolith. Watersheds exposed to relatively recent glaciation exhibited significant Ca and DIC enrichments while watersheds comprising relatively old unglaciated terrains did not show such enrichments. The frequency at which calcite is present in a relatively large suite granitic rocks, the range in car-

Figure 1. Variation in selective oxide contents of granitic samples.

435

bonate concentrations and the possible correlations with granitoid compositions are presently considered. Coupled with the effects of disseminated calcite on experimentally weathered solutes, the results of this investigation are important in assessing the effects of disseminated calcite during the weathering in silicate rocks, particularly in the context of global cycling and potential climate feedbacks with atmospheric CO,. 2. RESULTS 2.1 Methods A total of 100 granitic samples were selected for the study from the rock collections of the United States Natural History in Washington D.C (n = Sl), the British Natural History Museum, London (n =6) and Wards Scientific Incorporated, Rochester, NY , USA (n = 13). Sample selection criteria included the lack of observable weathering, oxidation or alteration. Samples were selected to reflect a diverse range in chemical compositions based on catalogue descriptions of rock type and mineralogy (Figure 1). Granitoid rock types included granites, syenites, monzonites, diorites and gabbros. Samples were also selected to include all the major landmasses on the earth’s surface. The samples were crushed and split into subsamples subsequently used for bulk chemical analyses by X-ray florescence spectroscopy and for experimental weathering studies. Acid digestiodgas chromatography was used to determine inorganic carbon in the granitoid rocks (White et al., 1999). Ten gram rock samples were placed in 300 ml glass bottles containing deionized water saturated with a 5% CO,/air mixture. The headspace was further flushed with the gas mixture and the bottles sealed, periodically shaken and opened after 75 days. The pH and alkalinity of the aliquots were immediately determined after which the remainder of the fluid sample was filtered through a 0.10 pm acetate filter. Cation concentrations were determined by inductively coupled plasma mass spectrometry (ICPMS).

distribution frequency for calcite in the present suite of granitic rocks is plotted in Figure 2. The mean concentration of calcite is 2530 ppm and the median concentration is 720 ppm. As shown in Figure 1, the granitic rocks sampled in this study represent a significant range in composition. Multivariant analysis was undertaken to determine if a correlation existed between major oxides in the granitoids and calcite concentrations. The results produced regression coefficients (r) of between -0.18 and 0.26 indicating the lack of statistically meaningful correlations. Perhaps most significant is the lack of any correlation (r = 0.10) between calcite and CaO in the granitic rocks, which varied between 0.4 and 12.7 wt.%. This lack of a correlation implies that calcite behaves as a trace phase and that the bulk of the Ca is contained in silicate minerals, predominantly plagioclase. In contrast, essentially all the carbon contained in the granitoid rocks is in the form of CaCO,. Calcite abundances in the granitic rocks, as determined by CO, analyses, were confirmed by selective thin section observations using cold stage cathode luminescence. Disseminated calcite occurs as large interstitial grains (up to 200 pm in diameter) as well as small, disseminated grains along silicate grain boundaries. Calcite also occurs as replacements in seriticized cores of plagioclase that may have been originally enriched in Ca. Finally, calcite occurs as fracture fillings that often cut across silicate grains. 2.3 Experimental solute compositions The significance of disseminated calcite in granitic rocks lies in its role of potentially impacting solute

2.2 Calcite distributions in granitic rocks All the granitic rock samples produced measurable CO, upon acidification, varying over a concentration range of 0.3 to 830 pM. These results, based on a relatively large population of samples, indicate that trace amounts of carbon are ubiquitous in granitic rocks. Based on microprobe and SEM analyses, this carbon corresponds stoichiometrically to calcite. The range of calcite concentrations is 50 to 18,800 ppm, which overlaps previously reported calcite concentrations in granitic rocks associated with watershed weathering (Mast et al., 1990; White al., 1999). The

Figure 2. Frequency distribution of calcite in granitic rocks

436

compositions associated with regolith weathering. The dissolution rate of calcite is 4 to 6 orders of magnitude faster than for that of silicate minerals such as plagioclase. Therefore, the presence of calcite is expected to impact solute compositions in a manner significantly out of proportion to its mass relative to Ca-containing silicates. This hypothesis is confirmed by the solute compositions produced from experimental weathering of the suite of granitic rocks used in this study. Solute Ca concentrations plotted against alkalinities, predominately in the form of bicarbonate, produce a strong linear correlation (Figure 3). Each data point corresponds to a solute composition associated with a specific granitic rock sample collected after two months of reaction. Two potential reactions are capable of producing this correlation. The dissolution of calcite can be written as CaCO, + CO, +H,O

+ Ca2++ 2HC0,'

(1)

and silicate hydrolysis, represented by plagioclase weathering, can be written as C ~ N a ( , - ~ ) A ( ~ + ~ ~+S i(l+x)%O, ~ , - ~ ) O *+ (2+2x)H20 -+ xCa2+ + (1-x)Na + (3-x)SiO2 + (l+x)HCO,- + (l+x)Al(OH), (2) where x is the anorthite component and 1-x is the albite component of plagioclase. Equation l produces

a 2 to 1 correlation between bicarbonate and Ca and as indicated by the dashed line in Figure 3 (slope = 2.0). The slope of the solute data for the granite rock samples is only slightly higher (slope = 2.44). In contrast the slope of the plagioclase hydrolysis reaction is dependent on the molar ratio (l+x)/x. The average Ca/Na ratio in the suite of granitic rocks is less than unity (Figure l), implying that the slope of CdHCO, resulting from silicate dissolution would also be less than 1. These results indicate that the short-term batch dissolution experiments produce solute compositions dominated by calcite dissolution irrespective of the initial calcite concentration nor the granitoid composition. The extent to which the solute Ca and alkalinity concentrations correlate with the respective calcite concentration of specific granite samples was also investigated. Results show a direct correlation at relatively low calcite concentrations in the granitoids but that solute Ca and HCO, are insensitive to the specific calcite content at high concentrations. The extent to which these aqueous species are controlled by calcite solubility was determined using the PHREEEQE chemical speciation code (Parkhurst 1997). The extent of calcite saturation in the solutes is plotted versus the calcite content of the granitic rocks in Figure 4. At low calcite concentrations, the corresponding solutions remain undersaturated (log IAP/K, < 0) and correlate strongly with the calcite contents of the granite. When calcite concentrations exceeded approximately 1000 ppm, the solutions exceeded calcite solubility (log IAP/K, > 0). Experiments employing granites with moderate amounts of calcites (5,000 to 10,000 ppm) produced the greatest supersaturation (log IAP/Ks =1.8) while granites with the highest calcite (< 18,000 ppm) produced solutes with lower supersaturation. This trend mimics that commonly observed for nucleation kinetics in which maximum supersaturation is achieved prior to the onset of precipitation. This trend suggests that secondary calcite is experimentally precipitated under closed system conditions as CO, is decreased and pH increases via Reactions 1 and 2. 2.4 Calcite paragenesis There appears to be a number of sources for calcite in granitic rocks. Some calcites such as those, which occur in vein fillings that crosscut primary phenocrysts are clearly of secondary hydrothermal or meteoric origin. However in many instances, disseminated calcite grains occur within or in direct

Figure 3. Correlation between Ca and alkalinity in solutes produced by weathering of a suite of granitic rocks. The dashed line corresponds to a CdHCO, ratio of 2 based on Equation 1. The solid line is the best fit to the solute data.

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what more radiogenic than the primary magma represented by the plagioclase isotopic ratio. This is consistent with late stage cooling processes in which hydrothermal fluids have mixed with other sources of Sr. CONCLUSIONS

Figure 4. Solute calcite saturation as a function of calcite content of the granitic rocks. Saturation is determined by the ratio of the ionic activity product (IAP) and the saturation constant K,.

contact with pristine primary silicate grains suggesting a contemporaneous origin. The paragenesis of calcite in igneous rocks has not been investigated in detail. Although both Ca and CO, are abundant phases during late stage magmatic cooling, thermodynamic and experimental uncertainties remain as to calcite stability at such PT conditions. One possible method of determining the paragenesis of calcite contained in granite rocks is based on ratios. Since Sr isotopes do not fractionate, minerals containing relatively low radiogenic Rb concentrations such as plagioclase are expected to approximate the 87Sr/86Srof the primary magma. Since calcite, likewise, contains relative little Rb, isotopic similarities with plagioclase would suggest a direct magmatic origin for calcite. Due to the small grain size of the disseminated calcite, physical separation and 87Sr/86Srdetermination is not possible. However the predominance of calcite dissolution (reaction 1) compared to plagioclase dissolution (reaction 2) establishes solute 87Sr/S6Sr ratios as reasonable surrogates for calcite. Sr isotopic ratios are compared for calcite and plagioclase ratios in Table 1. As indicated the calcite 87Sr/86Srratio, although similar to plagioclase, is in all cases slightly higher suggesting that calcite formed from a source some-

A suite of 100 granitic rocks from around the world was analyzed for CO,. Results indicated that calcite is a ubiquitous minor phase in granitic rocks. Petrographic evidence indicates that at least some calcite forms during late stage magmatic cooling, a process supported by 87Srs6Srdata. The weathering rate of calcite is several orders of magnitude faster than silicate minerals. Therefore calcite dissolution is expected to significantly impact solutes derived from crystalline rock weathering if residual amounts persist in the regolith. This is substantiated by experimental weathering in which solute Ca and HCO, is derived predominantly from carbonate dissolution. REFERENCES Bemer, R. A. & E. K. Berner 1997 Silicate weather& and climate. Tectonic Uplift and Climate Change. ed. W. F. Ruddiman. 353-364, New York: Plenum Press. Huntington, T. G., R. P. Hooper, B. T. Aulenbach, R. Cappellato & A. E. Blum 2000 Calcium depletion in forest ecosystems of the southeastern United States. Soil Science Society of America 5.64: 1845-1858. Mast, M. A., J. I. Drever & J. Barron 1990 Chemical weathering in the Loch Vale watershed, Rocky Mountain National Park, Colorado. Water Resources Research 26: 297 1-2978. Parkhurst, D. L. 1997 Geochemical mole-balance modeling with uncertain data. Water Resources Research 33: 19571970. White, A. F., T. D. Bullen, D. V. Davison, M. S. Schulz & D. W. Clow 1999 The role of disseminated calcite in the chemical weathering o f granitoid rocks. Geochimica et Cosmochimica Acta 63: 1939-1999.

Table 1 Comparsion of 87Sr86Sr ratios of plagioclase and calcite in several grantic rocks. Granitic Rock Yosemite, CA USA Luquillo, PR Loch Vale, CO, USA Panola, GA, USA

Plagioclase 0.7065 0.7041 0.7422 0.708 1

Calcite 0.7100 0.7063 0.7489 0.7191

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Aqueous Dissolution Studies of Synthetic and Natural Brannerites Y.Zhang, G.R.Lumpkin, B.S.Thomas, Z.Aly, R.A.Day, K.P.Hart & M.Carter Australian Nuclear Science and Technology Organization, PMB 1,Menai, NSW 2234, Australia

ABSTRACT: The dissolution of synthetic brannerite is incongruent and pH dependent with a minimum in the dissolution rate at near pH 8. Preferential release of U leaves Ti02 on the surface with different morphologies; smooth uniform layers in acidic media and nano-spherule agglomerations in alkaline media. The measured apparent activation energies at pH 5.6 to 9.8 suggest surface reaction-controlled dissolution mechanisms. A natural brannerite sample is amorphous and has shown very little alteration over geological time; under laboratory test conditions the U release is less than an order of magnitude more than in synthetic brannerite.

1 INTRODUCTION

2.2 Electron microscopy Scanning electron microscopy (SEM) was carried out with a JEOL 6400 instrument operated at 15 kV, and fitted with a NORAN Voyager IV X-ray microanalysis system (EDX). Calibrations were carried out using a comprehensive set of standards for quantitative analysis (Vance et al. 1996). Transmission electron microscopy (TEM) was carried out with a JEOL 2000FXII instrument equipped with a Link-ISIS EDX system. It was operated at 200 kV and was calibrated for quantitative thin-film analyses using an extensive set of natural and synthetic reference materials (Lumpkin et al. 1994).

Brannerite has been characterized as a resistate mineral with the general formula U1-xTi2+x06(Ifill et al. 1996). It has attracted our recent attention since it also exists, as a minor phase, in the ceramic formulations designed for immobilisation of surplus plutonium (Jostsons et al. 1999). The present study aims to provide better understanding of the dissolution behaviour of synthetic brannerite and the effect of amorphisation on the dissolution. 2 EXPERIMENTAL

2.3 Methods 2.1 Materials The synthetic brannerite was prepared using the alkoxidehitrate route as reported elsewhere (Vance et al. 2000). The sample contains mainly brannerite with -5-7% of rutile and only trace amounts of U02. A powdered sample (75-150 pm) was prepared and washed with acetone to remove the fines. The specific surface area was measured by a multi-point BET method using argon gas to be 0.082 t- 0.001 m2 g-'.

The natural brannerite sample was obtained from a uranium ore deposit at Cordoba, Spain. A powdered natural sample (75-150 pm) was prepared by the same procedure as that used for the synthetic brannerite. The specific surface area was measured by the BET method to be 0.33 t- 0.05 m2 g-'.

Static dissolution tests were carried out in duplicate at 9OoC using about 0.02 g of sample and 15 mL of pH 4 aqueous solution (0.05 M KHphthalate and 0.025 M HCl) or deionised water in an open atmosphere without leachant replacement. Dynamic dissolution tests were conducted for synthetic brannerite at 20-90°C in solutions with pHs ranging from 2 to 12 and flow rates between 10 and 15 mL d-' in an open atmosphere. The detailed experimental setup including the leachant compositions is described elsewhere (Knauss & Wolery 1986, McGlinn et al. 1995). The same pH solutions were used in the study of both natural and synthetic brannerites and only where specifically noted was deionised water used. The effluent was collected regularly and acidified after pH measurement.

439

Inductively coupled plasma mass spectrometry (ICP-MS) was used to determine the levels of elements in the solutions. The samples after testing were gently rinsed with deionised water and dried for microscopic examination. 3 RESULTS AND DISCUSSION 3.1 Dissolution of synthetic brannerite The U release rates in different pH solutions at 70°C are shown as a function of time in Figure 1. Except at pH 7.9, initially U releases decrease with time and reach relatively stable after 50 days whereas at pH 7.9, U releases remain fairly constant. Compared to U, Ti release rates generally are about 1-2 orders of magnitude lower indicating incongruent dissolution of brannerite, consistent with the formation of Tirich alteration rims observed for natural samples (Lumpkin et al. 1999). 3.1.1 Effect ofpH Rate constants for U releases were calculated using linear regression of cumulative U released vs. time (Luce et al. 1972). The rate constant vs. pH at 70°C (Fig. 2) demonstrates the existence of a minimum near pH 8. For pHs 2-8, the reaction rate varies inversely with pH, with reaction order of -0.27 for hy-

Fig. 3. The plot of rate constants against inverse temperature

drogen ion activity. For pHs 8-12, the reaction rate increases with increasing hydroxyl ion activity, with reaction order of -0.23. Above pH 10, the reaction rate is apparently lower. This could be a result of the fast deactivation of the active surface sites due to the presence of C0j2- or reflect differences in ratecontrolling steps (Thomas, unpubl.). The kinetic rate laws display a fractional reaction order with respect

-

Fig. 1 . U release rates from synthetic brannerite in different pH solutions at 70°C

Fig. 2. Effect of pH on the U release rate from synthetic brannerite at 70°C.

Fig. 4. SEM secondary electron micrographs of the synthetic brannerite samples after testing, a) 185 days in pH 5.6 solution; b) 156 days in pH 1 1.9 solution.

440

Table 1. Compositions of the unaltered and altered natural brannerite sample Wt% (oxide) p205

SiO, Ti02 ThO, U02 y203

=E203 CaO MnO Fe0 PbO

Fig. 5. Backscattered electron image of the natural brannerite. A thin alteration rind (-100 pm) is revealed by darker contrast.

to pH suggesting surface-reaction controlled dissolution mechanisms (Grandstaff 1980). 3.1.2 Effect of temperature The plot of rate constants against inverse temperature (Fig. 3) gives the U release reactions at pH 5.6, 7.9 and 9.8 with apparent activation energies of 67+2, 3924 and 53+1 kJ/mole, respectively, consistent with surface reaction-controlled dissolution mechanisms (Lasaga 1998). 3.2 Surface characterization of synthetic samples SEM examination of the synthetic samples after dissolution tests (Fig. 4) revealed different surface morphologies with smooth uniform TiOz layers in acidic media and nano-spherule agglomerations in alkaline media. 3.3 Characterization of the natural brannerite Backscattered electron images of natural brannerite show that the sample is largely free of alteration, except for a thin (approximately 100 pm) alteration rind revealed by darker contrast in the images (see

Unaltered' < 0.1 < 0.1

37.4 f 0.8 1.1 *0.6

52.0 f 0.9 0.3 f 0.2 0.5 f 0.1 0.4 f 0.2 4.1 f 0.6

0.1 f 0.2 1.O f 0.6

3.2 f 0.5

Altered rim"

4.4 f 0.1 0.4 f 0.2 73.8 f 1.7 3.1 f 1.8 7.7 f 0.7 3.0 f 0.6

10.1

10 million years if the selected value for effective diffusion coefficient (Deff)value is representative.

Figure 4. Transient out-diffusion of marine pore water at time = t years, DeE= 5x10'" m2/s.

451

7 CONCLUSION

5.2 Water-rock interactions Geochemical model calculations, in which there are quite large uncertainties, indicate that the pore waters are in equilibrium with calcite and dolomite. Cation exchange with clays controls the distribution of Na, K, Ca and Mg. Analyses of ammonium, NH4, in pore waters from DA and FB give concentrations between 0.5 and 4 mg/l. NH4’ may be significant in exchange equilibria due to its strong sorption onto clays. Analyses of samples of mudrocks from FB show that total NH4 is 280-640 pg/g, with more of it being bound than exchangeable. NH4 accounts for ~ 3 % of the total exchangeable cations. Data for dissolved and sorbed NH4 indicate that &(NH,I) is between 10 and 103 ml/g. 6 COMPARISONS BETWEEN SITES

Changes with depth of salinities and stable isotope ratios of pore waters in Mesozoic mudrocks contain information about the evolution of these groundwater systems. Variations between them are probably related to different histories of depositional conditions, burial and exhumation, erosion, and the processes of mixing with meteoric water. This ‘palaeohydrogeological’ evolution has influenced the budgets of reactive solutes and the potential for water-rock reactions. Interpretations of solute movements and geochemical equilibria in mudrocks are relevant to weathering and diagenetic studies, as well as being important in assessing mudrocks as hydraulic and chemical barriers for containment of contaminants. ACKNOWLEDGEMENTS

Depth profiles of chloride at the different sites have varying concentration gradients. These might be related to different uplift, erosion and groundwater histories, affecting the duration of exposure to meteoric water influx. They also depend on whether advection or diffusion has controlled water and solute transport. The steepest gradients of chloride versus depth are at FB, followed by M and some of the DA profiles. The lowest gradients are at EBB, EE and some of the DA profiles. The steeper gradient at FB suggests that the Jurassic sequence has been exposed to meteoric water influx for a relatively shorter period. This is possibly due to greater erosion rates in this region than at other sites. The steeper profiles of chloride at DA are on the upthrow side of a fault, and movement on the fault would have resulted in greater erosion of the upthrown block. The lower chloride gradients in mudrocks that are in or near to the fault zone suggest that groundwater advection occurs in the disturbed zone, giving shorter path lengths for diffusive dilution of pore waters in adjacent mudrocks. The low gradient of chloride at EBB suggests that advection may be significant here, because laterallycontinuous interbedded limestones impose an upwards hydraulic gradient through the mudrocks. The salinity increase at HW starts below the chalk aquifer at 90 m depth and initially has a fairly steep gradient but flattens out in the Oxford Clay (Fig. 2). The boundary concentrations for salinities in mudrocks at HW are constrained by the deep aquifers.

The studies on which this paper is based were carried out while the author was at the British Geological Survey, and the contributions of BGS colleagues are acknowledged. Various parts of the studies were funded by UK Nirex, the European Commission and the UK Department of the Environment.

REFERENCES Bath, A.H., Ross, C.A.M., Entwisle, D.C., Cave, M.R., Green, K.A., Reeder, S. & M.B. Fry 1989. Hydrochemistry of pore waters from Lower Lias siltstones and limestones at the Fulbeck site. Safety Studies Research Report NSS/R171. Harwell: UK Nirex. Brightman, M.A., Bath, A.H., Cave, M.R. & W.G. Darling 1985. Pore fluids from the argillaceous rocks of the Harwell region. Report FLPU 85-6. Keyworth: British Geological SUNey. Entwisle, D.C. & S. Reeder 1993. New apparatus for pore fluid extraction from mudrocks for geochemical analysis. In D.A.C. Manning, P.L. Hall & C.R. Hughes (eds). Geochemistry of Clay - Pore Fluid Interactions. 15:365-388. Mineralogical Society Series: 4. London: Chapman & Hail. Metcalfe, R., Reeder, S., Cave, M.R., Green, K.A., Entwisle, D.C. & J.R. Davis 1998. Fault-controlled groundwater flow in mudrocks at Down Ampney, UK: geochemical evidence. In Proc. NEAIEC Workshop on Fluid Flow Through Faults and Fractures in Argillaceous Formations, Berne, June 1996: 369-380. Paris: Nuclear Energy Agency of the OECD. Ross, C.A.M., Bath, A.H., Entwisle, D.C., Cave, M.R., Fry, M.B., Green, K.A. & S. Reeder 1989. Hydrochemistry of porewaters from Jurassic Oxford Clay, Kellaways Beds, Upper Estuarine and Upper Lias formations at the Elstow site, Bedfordshire. Safety Studies Research Report NSS/R172. Harwell: UK Nirex.

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Groundwater in the urban area of Catania (Sicily, Italy). Geochemical features and human-induced alterations M.Battaglia & P.Bonfanti Dipartimento di Scienze geologiche, sezione di Geologia e Geofisica, Universitd degli Studi di Catania, Corso Italia 55, 95129 Catania (Italy)

ABSTRACT: A total of 59 groundwater samples, taken from wells located in the urban area of Catania, have been analysed for major, minor and some trace element concentration. Water samples show low temperatures, near-neutral values of pH, medium - low conductivities and bicarbonate alkaline-earth to chloride-sulphate alkaline-earth composition. The concentration of some minor components and metals is usually very low and often below the detection limits. Using a quick classification and mapping methodology of quality of groundwater resources recently proposed by Civita et al. (1993), a groundwater quality map of this area has been prepared. The map clearly shows an extremely serious environmental picture and large variability of chemical parameters. The results of this research show that the studied area is largely characterised by groundwater unsuitable for potable use, with some reservations for irrigation and industrial uses.

1 INTRODUCTION Urbanisation cannot only cause radical changes in mechanisms of groundwater recharge, but also in their chemical composition and quality for drinkable, agricultural and industrial purposes. As a result, a notable decrease of water resources availability, an increase of water-supply costs, and a growing risk for human health has occurred. The severity of such phenomenon varies as a function of manifold factors: the most important are the local hydrogeological setting and the development level and coverage of sewer system. On the urban area of Catania, the second most important city of Sicily (Italy), there is a human load of about 400.000 inhabitants. Due to increasing population over the past years, the city has chaotically expanded. That implies serious pollution risks for groundwater, because of the partial lack of suitable and functional sewers (Ferrara & Pennisi 1995). The aim of this study was to perform a hydrogeochemical characterisation of groundwater inside the urban area of Catania, in order to determine its suitability for use (mainly the drinkable one) and the factor which determine the resource quality. This methodology has already been applied in several areas of Sicily, in order to verify - in various hydrogeological and hydrochemical settings - the capability and the sensitivity of the methodology to

represent quality.

disparate situations of

groundwater

2 GEOLOGICAL AND HYDROGEOLOGICAL SETTING The Catania metropolitan area is situated on the eastern coast of Sicily, between the Ionian Sea and the southern flank of Mount Etna. The area enjoys a typical Mediterranean climate, with average annual values of rainfall of about 700 mm, the bulk of which occurs between November and February, and temperatures of about 18”. Precipitation is the main source of groundwater recharge (Ferrara 1975). The stratigraphical sequence in the urbanised area of Catania is generally made up of terrains with sedimentary and volcanic origin, representing a time period that extends from Pleistocene to Holocene, locally covered by recent and modem deposits, both marine and continental (Monaco et al., in press). In the southern part of the urban area, at the “Terreforti” hills, a sedimentary series of Pleistocene age outcrops. Grey-bluish silty-marly clays with no clear stratification, of Lower-Middle Pleistocene age (Wezel 1967), sometimes interbedded with quartziferous sandy silts (Francaviglia 1940) are the oldest term of the succession. A Middle Pleistocene (Kieffer 1971) sequence of coastal sands and coarser deposits, consisting of fluvial - deltaic gravels and pebbles, follows upward.

453

It is unconformably overlapped by terraced deposits of both marine and coastal alluvial origin, distributed at various elevations and characterised by sands, conglomerates and silty clays (Monaco et al., in press). In the northern part of Catania, the effusive materials from the Etna volcano represents the most frequently cropping out rocks. These consist of thick basaltic lava flows that cyclically invaded the urban area in both pre-historical an historical times, filling the deep valleys entrenched in Pleistocene substratum. South of the Terreforti Hills lies the Simeto River coastal plain, a lowland characterised by recent and present alluvium deposits grading to coastal deposits. 3 MATERIALS AND METHODS During winter 1999, a total of 59 groundwater samples were collected from different sites, comprising principally wells located in the urban area of Catania. Sample locations are shown in Figure 1. Before sampling and in-situ measurements, tap water was allowed to run for several minutes. Water and air temperature, pH, and conductivity were determined directly in the field with portable instruments. Alcalinity was also determined the same day of collection by titration with HCl O.lN, using methylorange as colorimetric indicator. Water samples were collected and stored in 250 ml high-density polyethylene bottles with screw caps for the laboratory analyses. Water samples were analysed in the laboratory as follows: Na , K by atomif: emission spectrophotometry (AES), Mg2+, Ca by atomic sorption spectroscopy (AAS) on unfiltered samples; +

Figure 1 - Simplified geological sketch map and location of sampling sites

454

Cl- were determined by titration with A8N03; SO4 by visible spectrophotometry with Ba2 ; NO3 by ultraviolet spectrophotometry ( U V S ) with brucina; trace elements by inductively coupled plasma optical emission spectroscopy (ICP-OES) on filtered (0.45 pm Millipore membrane filter) acidified (ultra-pure HNO3) samples. 4 RESULTS AND DISCUSSION 4.1 Groundwatergeochemistry outlines All the analytical results are given in Table 1 in statistical form. Water samples show low temperatures, nearneutral values of pH, medium - low conductivities and bicarbonate alkaline-earth to chloride-sulphate alkaline-earthcomposition. The chemical composition of the analysed waters is certainly conditioned by human activities as the intense urbanisation and the agricultural practices, the last ones particularly developed in the southern sector of the city. This account for the elevated concentrations of ions C1, SO4, NO3 (the only nitrogenous compound Table 1. Chemical composition of the analysed groundwater (in ppm). T in "C, Cond. in pS/cm. Detection limits are indicated in parentheses. The number of measures is referred to the only samples with concentrations higher than detection limits. N. of Mean Median samples T 59 17.6 18.4 59 7.44 7.36 PH Cond. 1224 1206 59 TDS 916.4 873.8 59 59 -1.91 -1.82 LOg(PC0z) Na 130.5 125.2 59 K 16.5 18.1 59 Ca 59 112.2 84.6 59 46.3 45.5 Mg CI 59 161.6 145.4 HC03 424.2 402.7 59 SO4 171.7 154.5 59 NO3 59 60.9 61.9 NH4 (0.01) 10 2.97 1.41 NO2 (0.01) 14 0.42 0.21 As (0.03) 10 0.044 0.04 Zn (0.02) 36 0.265 0.0965 0.09 0.109 3 Pb (0.02) Ni (0.02) 0.021 0.02 3 6 Ba (0.02) 0.074 0.0885 Fe (0.01) 34 0.09 0.047 58 B (0.02) 0.664 0.663 Mn (0.02) 0.104 0.0885 12 36 v (0.01) 0.038 0.036 3 c u (0.02) 0.319 0.041 . , i

Min

Max

StdVar

9.2 6.72 769 464.3 -3.41 56 2.3 22.8 19.7 78 195.3 32 2.7 0.14 0.02 0.04 0.021 0.028 0.02 0.022 0.012 0.089 0.021 0.011 0.024

25.4 3.7 8.8 0.42 1988 269 1559.3 257 -1.14 0.45 281.8 42.53 63.7 10.9 309.2 70.66 129.9 18.9 443.2 59.26 1031.2 118.1 429 87.37 215.9 40.38 15.6 4.68 1.23 0.44 0.06 0.007 2.77 0.486 0.132 0.055 0.024 0.002 0.102 0.03 0.595 0.127 1.284 0.363 0.225 0.063 0.076 0.017 0.893 0.497

values of which were found in areas without sewerage network or with unsuitable sewers. 4.2 The groundwater quality map

Figure 2 - Langelier-Ludwig diagram for the analysed groundwaters.

constantly and severely present) and B. Boron shows a weak positive correlation with nitrate and, for the southern zone of the City, also with chloride. As regards the trace elements, Fe, V, Zn and - in small measure - also Mn, are very diffused (sometimes in elevated concentrations). The presence of such elements is certainly connected both to water-rock interaction (Giammanco et al. 1996) and to human - induced alterations. Besides it is evident that V is almost lacking in the waters circulating in the sedimentary basement, whereas the presence of Fe and Mn, surely linked to the clayey component of the acquifer it is frequent (Karro 1999). The concentration of the remaining minor components and metals is usually very low and often below the detection limits. Cr, Cd and COhave been detected in any sample. The only exception are the nitrogen cycle compounds NH4 and NOl, the highest

Description Verygood Acceptable Poor Class A B C c1 c2

Following the quick classification and mapping methodology proposed by Civita et al. (1993) and shown in Table 2, a groundwater quality map of this area was created. Civita et al. (1993) selected a small number of parameters (see table 2) the values of witch are subdivided into three intervals, defined on the basis of the European law. The described intervals correspond to progressively worse quality classes, to which a use assessment is given. In order to obtain a classification of water quality, first indicates the class of parameters of group 1, then the one of the group 2. (i.e., if all values of groups 1 and 2 belong to class "B", it is a B1B2 type water: but even if only one of parameters of the first group belongs in the "C" class interval, the water is classified as ClB2). The map was prepared using a GIS: first we assigned to each sampling site a numeric code (10, 20, 30 for the Al,Bl,Cl classes, and 1,2,3, for A2,B2,C2 classes). Then, a zoning of the study area by using a common interpolation method (kriging in this case) was performed. The usual procedure through the overlapping of isopleth maps of each parameter was disregarded because it could create empty areas or non-existent classes. The map could be used to evaluate the quality and to find out the sectors suitable for exploitation. It is clear that most of the urban area of Catania is distinguished by groundwaters unsuited for potable supply, but suitable for limited industrial and irrigation uses ("C 1'I class). This characterisation is mainly determined by high values of nitrate and - to some extent - by total hardness and chloride. The described circumstance might be attributed to the superimpositionof high anthropogenic loads (nitrate from sewer networks) to hydrogeological factors

Parameters group I (chemical-physical) E1.Cond. SO4 C1 NO3 @/cm (mgll) (mg/l) (mg/l) < 1000 < 50 < 50 < 10 1000 - 2000 50 - 250 50 - 200 10 - 50 > 2000 > 250 >200 >50

2 (undesirable mbstances) Th Fe Mn NHq ("F) (mdl) (mgJ1) (mgll) 15-30 A 0,2 > 0,05 > 0,5 Description Drinking water without treatment; suitable for almost all industrial and irrigation use. Water drinking without treatment; some restriction for industrial and irrigation use. Water unsuitable for drinking without treatment and with restrictions for other uses, Must be subjected to specific treatment. Must be subjected to simple or extensive treatment.

Class

455

ACKNOWLEDGEMENTS The authors would like to thank D. Avola, A.C. Sorge and 0. Musumeci for his friendly support in the fieldwork; Dr. Silvia Rizzo and Dr. Nino Brancato for help in laboratory. This research was supported by a University of Catania grant.

REFERENCES Civita, M., Dal Pra, A., Francani, V., Giuliano, G., Oliviero, G., Pellegrini, M. & Zavatti, A. 1993. Proposta di classificazione sintetica e mappatura della qualita di base delle aque sotterranee. Quaderni di tecnica e protezione ambientale 49: 107-109. Bologna: Pitagora editrice. Ferrara V. 1975. ldrogeologia del versante orientale dell’Etna. Atti del 3” Convegno Internazionale sulle Acque Sotterranee, Palermo. Ferrara, V., 1990. Carta della vulnerabilita all’inquinamento dell’acquifero vulcanico dell’Etna. C.N.R. - G.N.D.C.I. Ferrara, V., 1998. Carta della wlnerabilita all’inquinamento dell’acquifero alluvionale della Piana di Catania (Sicilia NE).C.N.R. -G.N.D.C.I. Ferrara V., & Pennisi A. 1995. Lo sviluppo urbano nell’area metropolitana di Catania ad i conseguenti problemi di protezione delle acque sotterranee. Atti del 2” coiivegno nazionale slrlla protezione e gestione delle acqrre sotterranee: metodologie, tecniche e obiettivi: 1.193-1.198 Francaviglia, A. 1940. Osservazioni geologiche sulle colline delle terreforti regione etnea. Giornale di Geologia 14: 5581. Giammanco, S., Valenza, M., Pignato, S. and Giammanco, G. 1996. Mg, Mn Fe and V concentrations in the ground waters of Mount Etna (Sicily). Wat. Res. (30) 2: 378-386. Karro, E., 1999. Chemical composition of groundwater in the Espoo area, Southern Finland. In J. Chilton (ed.), Groiiiidwater in the urban environment: Selected CiQ Profiles: 165-170.Rotterdam: Balkema. Kieffer, G. 1971. Depsts et niveaux marins et fluviatiles de la region de Catane Sicile, leurs correlations avec certains episodes d’activite tectonique ou volcanique. Mediterranee, 5-61 591-626. Monaco, C., Catalano, S., De Guidi, G., Gresta, S., Langer, H. & Tortorici, L. in press. The geological map of the urban area of Catania (Eastern Sicily): morphotectonic and seismotectonic implications.Boll. Soc. Geol. It. (in press) Wezel, F.C. 1967. I terreni quaternari del substrato dell’Etna. Atti Acc. Gioenia Sc. Nat. Catania, Suppl. Sc. Geol. 27 1281.

Figure 3 - Groundwater quality map

(high hardness due to the interaction between water, gaseous phases and rocks, and - secondarily chloride from sea water mixing in coastal areas). Moreover, further sub - areas have been defined from the distribution of parameters of group 2: in this case the distinguishing factor is essentially the ammonium, strictly linked to the introduction into the aquifer of untreated urban sewage. Iron seems instead to be responsible of a rough subdivision in two areas: the western one is characterised by A2 groundwaters, while in the eastern one a B2 classification predominates. Manganese shows concentrations even included in the interval of the “A”class (< 0.02 ppm). 5 CONCLUSIONS

This paper represents a contribution to the problem of groundwater quality worsening assessment in urban areas. Groundwater plays a fundamental role in the potable supply of Catania area, being a potentially valuable but underused resource. This study evidenced that the lack of suitable and functional sewers engenders severe derogation of quality in large areas so it seems reasonable to take urgent measures for protecting groundwater resources. The spread of sewage contamination over the area is suggestive of a result of a multi-point source input. The correlation between groundwater quality maps and vulnerability maps already available for the whole Etnean area and for the Catania Plain (Ferrara 1990, 1998) could allow, to brief term, to construct maps useful not only for forecast and prevention of the pollution, but also for planning exploitation.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Study on water quality in the area of Wadi Shueib, Jordan Valley, Jordan KBecker, W.Ali & H.Hoetz1 Department of Applied Geology, Karlsruhe University, Germany

ABSTRACT: The study deals with water quality investigations on an area in the eastern escarpment of the Jordan Valley west of Amman, Jordan. Structurally the area is affected by the Jordan Rift Graben System. Three hydrogeological units exist in the study area of Wadi Shueib: a-The Lower Cretaceous Aquifer Complex, b- The Upper Cretaceous Aquifer Complex and c- The Upper Tertiary/ Quaternary Jordan Valley Alluvium Aquifer Complex. The springs discharge forms the base flow of Wadi Shueib. 78 water samples were taken from springs, reservoir and wells of the shallow aquifer in Wadi Shueib area were analyzed in year 2000. They were collected mainly in the middle to the lower part of the Wadi. The preliminary results show that electrical conductivity values range from 2000 to more than 3500 pS/cm. The Chloride values range between 200 and 750 mg/l, the Sulphate values range between 90 and 700 mg/l and the Nitrate values between 160 and 180 mg/l. Results of isotope analysis show that the values of 6 0 vary between -4,68 and 6,25. Two hydrochemical groups can be distinguished in the study area. The values of the second group are higher, indicating a longer flow path within limestone aquifer matrix. 1 INTRODUCTION

3 AQUIFERS

Water resources are limited in the Middle East region. Studies on water demand, water quantity, and quality are therefore very essential. The results presented in this paper are part of a multinational Research Project in the area of the Jordan Rift Valley. The Wadi Shueib is located (Fig. 1) in the eastern escarpment of the Jordan Rift Graben System between Amman the capital of Jordan in the east and the Jordan River in the West. This study concentrates on the hydrogeological patterns and the surface and ground water quality.

The stratigraphy in the study area shows outcropping strata of Cretaceous or younger sediments. Three aq-

2 HYDROLOGY

The precipitation in the study area ranges from 570 mm per year in the upper catchment by Salt in the highland to a value of 176 mm per year in Shuna in the lower catchment area. The mean precipitation over the wadi catchment is 430 mm (average values 1985 - 1999). Figure 2 shows the annual for the Same period for the stations Salt, Wadi Shueib and Shuna.

Figure 1 . Location of Wadi Schueib on the eastern Escarpment of the Jordan Valley.

457

The increase of the water level from 1980 to 1993 is due to leakage from the Shueib dam and King Abdullah Channel, which are only few meters apart from this well. The decrease in the last 5 years sums up to 17 meters.

4 WATER QUALITY

Figure 2. Annual rainfall (1985 - 1999) for the stations Salt, Wadi Shueib and Shuna (data from Ministry of Water and Irrigation, Amman).

uifer groups exist in the study area (Subah 1998, Ta'any 1992, Hirzalla 1974): a- The Kurnub Group of Lower Cretaceous age (Lower aquifer) is composed of sandstone. b- The Ajlun and Belqa Groups of Upper Cretaceous to Lower Tertiary age (Upper aquifer) consist of alternations of limestones, dolomites and mark with increasing chert and chalk content in the stratigraphically higher sections c- Finally the clastic sediments and evaporites of the Jordan Valley Group of Upper Tertiary age to recent (alluvial aquifer). The diagram in Figure 3 shows the changes in groundwater level in one observation well (AB 1340). This well monitors the shallow aquifer in Wadi Shueib area near the village of Shuna. (Groundwater level observations data from the Ministry of Water and Irrigation /Amman). The longterm groundwater level indicates seasonal variations. The level is increasing in winter months due to infiltration of precipitation and decreasing in the hot summer months due to high pumping rates and evaporation.

78 water samples from springs, shallow aquifer and from Shueib reservoir were collected in year 2000 and analyzed in the Laboratory of the Department of Geology in AmmadJordan. Two types of water can be identified in the shallow aquifer (Fig. 4): 1. Alkaline earth water with high alkaline component (Mg, Ca, Cl) predominately chlorid. 2. Alkaline water suphate/chloride

(SO4,

Cl)

predominately

This water chemistry results fiom mixing processes due to fieshhalt water interaction with the rock formations. An increased nitrate content shows the effect of the agricultural activities in the area (Lenz 1999). Analysis of isotopes in water samples (I8 6 0 and Deuterium) were carried out at the Umweltforschungzentrum Leipzig Halle in Germany. All water samples are located between the global meteoric waterline and the eastern Mediterranean waterline. Two groups can be distinguished (Fig. 5): Group 1 with two springs is located at the upper part of Wadi Shueib south of Salt. The water origi-

Figure 4. The results of the hydrochemical analysis of the water samples from wells in the shallow aquifer in the Shueib reservoir and Shouna area.

Figure 3. Groundwater level changes at Shuna observation well AB1340 (data from MWI, Amman).

458

Figure 5 . Group 1 and 2 (18 6 0 and Deuterium compositions) in water samples in Wadi Shueib area.

Lenz; S. 1999. Hydrological Investigations along Wadi a1 Kafrein and the Kafrein Reservoir Jordan.- Master Thesis University of Karlsruhe.; Karlsruhe. MC Donald SIR and Partners 1965. East Bank Water Resources, 6 Vols; VOL2: East Ghor Side Wadis-Central Water Authority; Amman. Rimawi, 0. 1985. Hydrogeochemistry and Isotope Hydrology of the Ground- and Surface Water in North Jordan (NorthNortheast of Mafiaq, Dhuleil-Hallabat, Azraq-BasinDissertation and der technischen Universitat Munchen, 240p.; Munchen. Subah, A. 1998. Environmental Isotope Study of the Artificial Recharge to the Groundwater in Jordan. Case of Jordan Valley.-unpubl. Report Water resources Studies Dept.; Ministry of Water and Irrigation MWI.; Amman. Ta‘any; R’A. 1992. Hydrological and Hydrochemical Study Of the Major Springs in Wadi Shueib Catchment Area.-M.Sc. Thesis, p 300 -Yarmouk University; Jordan.

nates from rainfall in the highlands. The chloride values with 57 and 109 mg/l are lower than those of group. The 6 0 values of group 2 represent the wadi water and the wells of the shallow aquifer in the lower part of the Wadi. These values are higher than those of group 1. This is due to interaction of meteoric water and limestone which usually causes an increase in the l 8 6 0 value of the water (Subah 1998, Rimawi 1985, Drever 1982). The arid climate and evaporation may increase the l 8 6 0 value. The chloride concentration of group 2 is higher than that of group 1, ranging between 138mg/l for the wadi water up to 664 mg/l for the shallow aquifer. ACKNOWLEDGEMENT This paper is part of the activities of the GIJP Joint Research Program for the Sustainable Utilization of Aquifer Systems” sponsored by the German Federal Ministry of Education, Science, Research and Technology (BMBF). The authors are grateful for administrative and financial support from the BMBF (Dr. J. Heidborn) and the Forschungszentrum Karlsruhe (Dr.W. Robel, Dr. H. J. Metzger, Mrs. S. Proboscht). Special thanks to our Jordanian colleagues (Prof. E. Salameh , Dr. H. El-Nasser) and their staff for their help in field and for data providing. Sincere thanks to Dr. S. Geyer and Prof. P. Moller for isotopes analysis. Thanks to all colleagues in the joint research program for their continuos support.

REFERENCES A1 Kuisi, M. 1997. Effects of Irrigation Water with Special Regards to Biocides on Soils and Groundwater in the Jordan Valley Area/ Jordan-Inaugural. Dissertation for PhD Universitat Munster.; Munster. Drever; J.I. 1982. The Geochemistry of Natural Waters.Prentice- Hall INC., 388 pp.; New York. Hirzalla, B. 1974. Wadi Jurei‘a and Wadi Shueib Groundwater Evaluation.-Paper Groundwater Section Water Resources Division-Natural Resources Authority NRA.; Amman.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Chemical evolution of ground waters in W-Iceland (Snzfellsnes) E.Bedbur, M.Petersen & H.Biallas Institute of Geoscience, University of Kiel, Gerrnany U.Wollsch1ager Institute of Environmental Physics, University of Heidelberg, Germany

S .Schmidt Federal Institute for Geoscience and Natural Resources, Hannover, Germany

ABSTRACT: Ground water from Snzfellnes Peninsula (W-Iceland) has its origin rain water and melt water from glaciers. Chemical analysis of waters from springs provided a set of data which was processed with statistical methods as factor and cluster analysis. Through factor analysis two main factors were identified and could be explained by marine influence and silicate weathering respectively. Cluster analysis allowed to distinguish between waters from balsalts, hyaloclastites, melt water and rain water, marine intrusions, cold carbon dioxidehhermal springs and near coast waters. Forward and inverse geochemical modeling with the computer code PHREEQC was carried out. Three flow paths were modeled, a basalt, a hyaloclastite and a thermal spring. Inverse modeling allowed to estimate the proportion of rainwater and seawater in the input. For springs influenced by volcanic gases like CO2, H2S and HC1, this input could be calculated.

1 INTRODUCTION

allowed the calculation of mass transfers of waterrock interactions within the aquifer, taking into account the analytical error.

Ground water in the western part of the Snzfellsnes Peninsula (W-Iceland) has its origin in rain water and melt water from the glaciers. The main aquifers are formed by Tertiary and Quaternary basaltic lavas which show high hydraulic conductivities (kf) of 10°-10-3 m/s. Less permeable aquifers are built by Quarternary subglacially formed hyaloclastites (kf = 10-2-10"6m/s) (Sigurhon 1990). In July and August 199'7 water samples were taken from the different aquifers and from thermal and cold carbon dioxide springs. The parameters pH, redox potential (EH), electric conductivity, temperature, HC03- and 0; content were measured in the field. The contents of Cl-, SO:-, N03-, F-, Br-, Na', K', Mg2', Ca2+,A13+,Sitot,Mntot,Fetot,Sr2', and B,,, were determined in the laboratory. Because of short flow paths in silicic rocks at low temperartures and with almost no vegetation, only little water-rock interactions could be expected. For most samples the total dissolved solids (TDS) was therefore less than 100 mg/l. Only the waters from the thermal or CO; springs gave TDS up to grams per liter. This set of data for all sample points was analyzed with statistical methods as factor and cluster analysis. Forward and inverse geochemical modeling with the computer code PHREEQC (Parkhurst 1995)

2 GEOSTATISTICS Factor and cluster analysis was carried out using the whole data set for each sampling point (Petersen, unpubl.; Biallas, unpubl.). The geostatistical analysis was performed for the western and the eastern part of Snzfellsnes separately. 2.1 Factor analysis The number of factors was determined with the Kaiser criterion, i.e. only factors with Eigen-values higher than 1 were extracted. After a varimax rotation the factor analysis yields two main factors. The result is similar for the eastern and the western peninsula. Factor 1 shows high loading (>0.6) of the following parameters: el. conductivity, Cl-, SO:-, Na', K', (Mg"), Sr2'. This factor indicates the marine influence on the groundwater chemistry. Factor 2 has high loading (> 0.6) in the following parameters: el. conductivity, HCO;, Mg2+, Ca2+, Sitot, Sr2+, reflecting the process of silicate weathering (Tables 1,2).

461

than ten times higher than the solubility of the crystalline basalt (Gislason 1985, Gislason & Arnorsson 1993). The higher solubility results in a higher proton transfer rate which leads to an increase in pH of the solution in a closed aquifer system. The higher solubility should lead to a higher TDS in the water samples taken from hyaloclastites which was not observed in the present study. In E-Snzefellsnes the situation appears to be more complicated. Clusters are more differentiated and volcanic influence can be shown. The age of the different lavas and their topographic exposition are also important for the chemical evolution of the respective waters.

Table I. Loading of factors, W-Snzefellsnes (Petersen, unpubl.) Parameter temperature electric conductivity PH HCO;

c1so:-

NO; Na' K' Mg2+ Ca2+ ~ 1 3 +

SiklL Sr2

Factor 1 0.24 0.78 0.19 0.33 0.90 0.92 0.41 0.86 0.79 0.70 0.57 -0.09 0.18 0.70

Factor 2 0.08 0.56 0.20 0.9 1 0.17 0.20 0.43 0.46 0.52 0.63 0.73 -0.02 0.65 0.60

3 THERMODYNAMICAL MODELING

2.2 Cluster analysis

Cluster analysis allows the differentiation of the aquifer material in W-Snzefellsnes by building two clusters with water samples taken from basaltic and from hyaloclastitic rocks. Three more clusters group samples with low TDS (melt and rain water), samples with high TDS (marine intrusions and cold carbon dioxide springs) and samples taken near the coast or near CO,-rich springs respectively. Samples from hyaloclastitic aquifers have a higher pH (8.2-9.3) than samples from the basaltic aquifers (pH 6.4-8.3). These relatively high pH values reflect their hydrochemical evolution in a closed system (Gislason & Eugster 1987a). Basalt and hyaloclastite have the same geochemical composition, but hyaloclastite is rich in basaltic glass in contrast to crystalline basalt (Gislason & Eugster 1987a). This results in a higher solubility of the hyaloclastite which is more Table 2. Loading of factors, E-Snzefellsnes (Biallas, unpubl.) Parameter temperature electric conductivity PH HCO;

c1s0:-

NO; Na' K+ Mg" Ca2+ ~ 1 3 '

Sit,, Sr2+

Factor 1 0.31 0.68 0.07 0.27 0.89 0.89 0.25 0.81 0.59 0.18 0.34 0.07 0.45 0.5 1

Factor 2 0.23 0.73 0.10 0.94 0.20 0.28 0.07 0.56 0.53 0.83 0.92 -0.38 0.60 0.66

Saturation indices (SI) for the different samples were calculated with the computer code PHREEQC (Parkurst 1995). In addition, inverse modeling of selected flow paths in basaltic and hyaloclastitic aquifers with PHREEQC was carried out (Wollschlager, unpubl.; Schmidt, unpubl.). The results of the modeling are consistent with the results of the cluster analysis. The mineral phases used in the inverse modeling are chosen from literature (Bistry 1986, Jakobsson 1972, Gislason & Eugster 1987b, Gislason et al. 1993, Nesbitt & Young 1984) (Table 3). Mass transfer was calculated by using the secondary phases only, which had shown supersaturation in the forward modeling. Depending on the flow path either a melt water or a rain water was used as the initial solution. There is a good correlation between the marine influence on the water and the topographic height (marine born aerosols). This was taken into account by giving the model either a concentrated rain water and/or the possibility to mix raidmelt water with sea water to produce a starting water for the inverse model. The modeling of waters from the eastern part of Snzefellsnes gives a different picture (Schmidt, unpubl.). The volcanic influence to some of the samples is reflected clearly in the geochemical models Table 3. Mineral phases used for the inverse modeling Mineral Forsterite Diopside Albite Anorthite Adularia Illite Kaolinite Ca-Montmorillonite Gibbsite Calcite Laumontite

462

Composition Mg,SiO, CaMgSi,O, NaA1Si30, CaAl,Si,O, KAlSi30, &.6M~0.2SA12.3Si3.So~O~oH

Al2Si,OS(OH), c%.16A12.33si3.67010(0H)2

Al(OH)3 CaCO, CaAl,Si,O,,

4H20

Figure 1. Mass transfers along a flow path in a basaltic aquifer. by a calculated CO, partial pressure significantly higher than 1O-I.j bar. 3.1 Flow path in a basaltic aquifer A sample (G8) representing water from the basalt cluster was inversely modeled with an input water containing 62% melt water, 38% rain water (5x concentrated) and 0.02% sea water (Fig. 1). In the model the primary minerals forsterite, diopside, albite, anorthite, adularia, and the secondary phases kaolinite, gibbsite, illite and Ca-montmorillonite were used. During the underground passage 0.02 mmol/l forsterite, 0.09 mmol/l albite and 0.04 mmol/l anorthite were dissolved, and 0.001 mmol/l illite and 0.07 mmol/l Ca-montmorillonite were precipitated (Fig. 1).

Figure 3. Mass transfer for a thermal spring The selected sample is a typical water from a hyaloclastitic aquifer. It has a low bicarbonate content of 15 mg/l. This is also a typical bicarbonate content of basaltic aquifers, but the pH in the hyaloclastitic aquifer is much higher (pH 8.9). The partial pressure of CO, for this sample was calculated to 10-4.sbar, which indicates that the hyaloclastitic aquifer system is a closed system. The saturation index for calcite is close to saturation equilibrium. The observed relatively low bicarbonate content is assumed to be due to the precipitation of carbonate. In this model 0.02 mmoVl forsterite, 0.05 mmol/l albite and 0.68 mmoVl anorthite dissolve and 0.02 mmol/l illite, 0.48 mmol/l gibbsite, 0.38 mmol/l Ca-montmorillonite, 0.56 mmol/l calcite precipitate (Fig. 2). 3.3 Thermal springs, cold carbon dioxide springs

3.2 Flow path in a hyaloclastitic aquifer A sample (G117) from a spring representing water

from the hyaloclastite cluster was inversely modeled with an input water containing 83% rain water, and 17% rain water (5x concentrated) (Fig. 2).

Modelling of the different thermal springs and cold carbon dioxide springs in the eastern part of Snaefellsnes gives very large rates of solution and precipitation due to volcanic influence on the waterrock interaction by volcanic gases and/or increased temperatures. For a cold carbondioxide spring the input of CO, was calculated and determined to 91.1 mmol/l. For a thermal spring the input of the gases H,S and HC1 were calculated to 1.1 mmol/l and 2.5 mmol/l repectively. The model of a warm carbon dioxide spring gives an input of 1.1 mmol HCI, 0.4 mmol/l H,S and 47.7 mmol/l CO, (Fig. 3). 4 CONCLUSIONS

Figure 2. Mass transfers along a flow path in a hyaloclastitic aquifer.

Geostatistical methods like factor and cluster analysis are means to differentiate the main factors which determine the water composition of the waters from SnEfellsnes Peninsula (W-Iceland) and were able to separate clusters of samples. Hydrochemical inverse modeling allowed quantification of the mass transfers within the aquifer. Mass transfer is higher in the

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hyaloclastitic aquifer than in a crystalline basaltic aquifer but is highest in the cold carbon dioxide springs and thermal springs due to the influence of temperature and volcanic gases. Dissolution of anorthite in the hyaloclastite is three times higher than in crystalline basalt. The precipitation of Ca-montmorilloniteis 6 times higher. It is shown that the hyaloclastitic ground water system is most possibly a closed system, where carbonate is precipitating. Although the mass transfer is high the TDS is low, because of high dissolution and high precipitation of minerals in the hyaloclastites. H' is consumed for precipitation of secondary mineral phases. This leads to high pH-values, although the TDS is similar to samples from basaltic aquifers. REFERENCES Biallas, H. 1998. Statistische Untersuchungen mariner, vulkanischer und petrographischer Einflusse auf die Grundwasserbeschaffenheitder ostlichen Halbinsel Snzefellsnes, West-Island. Dip1.-Arb. Univ. Kiel (unpubl.). Bistry, T. 1986. Naturlicher und anthropogener Stoffeintrag in das Grundwasser der vulkanischen Ozeaninsel La Palma. Ber. -Rep. Geol,Pal.Inst. Kiel 85: 1-172. Gislason S.R. 1985. Meteoric water-basalt interactions. A field and laboratory study. Ph.D.Thesis. John Hopkins Univ., 238p. Gislason S.R. & S. Am6rsson 1993. Dissolution of primary basaltic minerals in natural waters: saturation state and kinetics. Chem. Geology 105:117-135. Gislason S.R. & H.P. Eugster 1987a. Meteoric waterbasalt-interactions.I: A laboratory study. Geochim. Cosmochim. Acta 5 112827-2840. Gislason S.R. & H.P. Eugster 1987b. Meteoric waterbasalt interactions 11: A field study in N.E. Iceland. Geochim. Cosmochim. Acta 5 1:2841-2855. Gislason, S.R., D.R. Veblen & K.J.T. Livi 1993. Experimental meteoric water-basalt interactions: Characterization and interpretation of alteration products. Geochim. Cosmochim. Acta 57: 14591471. Jakobsson, S.P. 1972. Chemistry and distribution pattern of recent basaltic rocks in Iceland. Lithos 5:365-386. Nesbitt, H.W. & G.M. Young 1984. Prediction of some weathering trends of plutonic and volcanic rocks based on thermodynamic and kinetic considerations. Geochim. Cosmochim. Acta 48: 1523-1534. Parkhurst, D.L. 1995. User's guide to PHREEQC - a computer program for speciation, reaction-path, advective-transport, and inverse geochemical calculations. Water-Resources Invest. Rep. 954227. 464

Petersen, M., 1998. Regionalstatistische Untersuchungen der Grundwasserbeschaffenheit auf der westlichen Halbinsel Snaefellsnes(W-Island). Dip1.Arb. Univ. Kiel (unpubl.). Schmidt, S. 1998. Hydrochemische Klassifizierung und thennodynamische Modellierung von Grundw2ssern der ostlichen Halbinsel Snsefellsnes, West-Island. Dip1.-Arb. Univ. Kiel (unpubl.). Sipasson, F. 1990. Iceland. In United Nations (eds.), Groundwater in Eastern and Northern Europe. Nat. Resourc. / Wat. Ser. 24:123-137. Wollschlager, U., 1998. Thermodynamische Modellierung regionaler Einflusse auf die Grundwasserbeschaffenheit des Snzefellsjokull-Gebietes(WIsland). Dip1.-Arb. Univ. Kiel (unpubl.).

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Hydrogeochernistry in the Flumendosa river basin (Sardinia, Italy) R .Caboi, A .Cristini, M .Collu ,F.Podda & L .Rundeddu Dipartimento di Scienze della Terra, Universita di Cagliari, Italy

ABSTRACT: The chemical quality of shallow groundwater in the Flumendosa basin was investigated in order to highlight a possible contamination from mineralised areas and abandoned mines, and to test suitable procedures for creating hydrogeochemical risk maps. The principal component analysis and geostatistical approach show a close association between lithologies and hydrogeochemistry in the area; at the same time, the anomalous concentrationsascribable to the mineralisations in the area are pointed out.

1 INTRODUCTION

2 GEOLOGICAL AND HYDROLOGICAL FEATURES

The Flumendosa river basin, located in the southeastern part of Sardinia, is the most important water resource in the region. In the past decades, several artificial reservoirs were built on the river to supply water for drinking, irrigation, and industrial purposes to approximately 50% of the island’s population, including that of the Cagliari metropolitan area (approx. 500,000 inhabitants). The Flumendosa basin has a surface area of approximately 750 km2. It is characterised by sectors with a limited anthropic impact (these areas will be included in the Gennargentu National Park), and territories that have been affected by mining activity now abandoned, mainly for sulphide ore, as at Gadoni (Fig. l). The presence of galleries, tailings, and mine effluents exposes the water resources to a risk of contamination owing to metal leaching and mobilisation (Bertorino et al. 1987). The situation is particularly critical in the reservoirs of the middle course of the Flumendosa and Mulargia rivers, where also urban wastewaters are discharged, and in the lower course where the overexploitation of the coastal aquifer for the irrigation determines a salinization process (Ardau & Barbieri 1994). This study is a part of the research program that will involve surface, lake, and stream waters with seasonal sampling and analyses of both dissolved and suspended metal contents. The purposes of the present study on shallow groundwaters, are to assess the water quality by the main hydrogeochemical features, to highlight the toxic metal pollution from mineralised areas, and to test the procedures for hydrogeochemicalmapping.

Figure 1 shows a schematic map of the geology in the Flumendosa basin. The geology of the study area is largely dominated by rocks of the Palaeozoic metamorphic-sedimentarycomplex, characterised by very low permeability, with limited and superficial water circulation. This complex is composed of quartzose sandstones interbedded with argilloschists (Cambrian-Ordovician), mostly black schists and limestones (Silurian-Devonian). In this sector there are subordinated outcrops of granitic and volcanic rocks (Permian-Carboniferous), and dolomites and limestones (Jurassic) (Barca et al. 1996). The rocks in the southern part of the basin are highly fractured, with macro and micro fractures easily permeable to groundwater flow. The Jurassic formations, Miocene sands, and Holocene alluvial sediments at the mouth of the Flumendosa river show high permeability, and host important aquifers. In the area the mean annual rainfall is around 600-700 mm. 3 METHODS

3.1 Sampling and laboratory analyses A first sampling campaign was carried out in JulyAugust 1999: a total of 37 samples (34 springs and 3 wells) were collected. Due to the drought over the past years, low flow values (loo m depth) have high concentrations of SiO,, Na' and F- due to dissolution of silicate minerals. (3) Factors controlling the chemical compositions of surface waters and groundwaters result from the following associations (F 1) CP-K+-NO,-, (F2) HCO, -Ca2+-S12'-Mg2+,(F3) F--Na+-SiO, and (F4) SO:-. This indicates that the inflow of external contaminants, the dissolution of carbonate minerals, the weathering of silicate minerals and the oxidation of sulfide minerals control the chemical characteristics of natural waters in the Okchon zone. (4) Hydrogeochemical modeling and activity diagrams suggest that dissolution of silicate and carbonate minerals is possible and weathering products are kaolinite and illite. REFERENCES Ball, J.W. & D.K., Nordstrom 1991. User's manual for WATEQ4F with revised thermodynamic data base and test cases for calculating speciation of minor, trace and redox elements in natural waters. US. Geol. Sum. Open File Rep., 91-183. Chon, H.T., Cho, C.H., Kim, K.W. & H.S. Moon 1996. The occurrence and dispersion of potentially toxic element in areas covered with black shales and slates in Korea. Applied Geochemistry 11: 69-76. Chon, H.T.,Lee, H.K.,Lee, J.U., Lee, D.H., Ryu, D.W. & S.Y. Oh 1997. A study on the variation of the surface and groundwater flow system related to the tunnel excavation in DONGHAE mine area (a)-Hydroeochemical consideration. J. Korean Society of Groundwater Environ. 4: 27-40. Fetter, C.W. 1994. Applied Hydrogeology. Macmillan College Pub. Co., New York. Helgeson, H.C. 1969. Thermodynamics of gydrothennal systems at elevated temperatures and pressures. Amer. J. Sci. 267: 729-804.

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New geochemical data of the high PCO, waters of Primorye (Far East Russia) 0.V.Chudaev Far East Geological Institute, Vladivostok, Russia

V.A .Chudaeva Pacific Institute of Geography, Vladivostok,Russia

K .Sugimori Toho University, School of Medicine, Tokyo, Japan

K .Nagao ,B .Takano ,M .Matsuo & A .Kuno Tokyo University, Tokyo, Japan

M .Kusakabe Okayama University, Missasa, Japan

ABSTRACT: The noble gas and chemical composition of CO2 - reach mineral waters from various structures of the Sikhote-Alin were studied. The chemical composition of CO2 - reach mineral waters depends on the host rock composition, CO2 partial pressure and residence time. 3He/4He (R/Ra) is shown to be within the (2.3-4.7) range, reflecting the presence of a considerable mantle He portion in the bicarbonate waters described. Ar and Ne were found to be of atmospheric origin. The U3He ratio (4.7-12 *107) and 6I3C values ranging between -4.19 o/oo and -8.19 o/oo indicate a magmatic origin for the carbon. Thus, the high PC02 mineral waters of the Primorye region have origins in the following factors: an exogene-water component and an endogene-carbon and helium gas component.

1 INTRODUCTION In Prymorye the most widespread mineral waters are high PC02 cold waters. Their characteristic features are a high content of CO2 (up to 98%) in the gaseous phase, and a great partial pressure, which may reach 1.6 atm. As a rule, these waters are localized in the major fault zones, which divide this territory into its major tectonic block-terranes. The data on chemical composition of waters (including REE) were published (Shand et al. 1995, Chudaeva et al. 1999), but data on gas phase is poorly studied. The origins of gases in various geological structures are still debatable. He, COz, and other gases of the Earth’s crust in active volcanic zones are considered to be mostly of mantle origin. In the areas like Primorye, where volcanic activity terminated in the recent past, the origin of gases is rather difficult to establish. Often it is determined by using isotopic measurement; helium, in this case, is the most informative. This paper reports the results of the first study of noble gas composition of high PC02 waters of Primorye. As the object of research the authors used several major groups of CO2 -reach waters localized in various geological structures of the Sikhote-Alin mountains. The samples were collected from west to east, perpendicular (from oldest to youngest) to the meridianal geological structures of the Sikhote-Alin. The following groups of CO2 -

reach waters were studied: Shmakovka (Medveji & Avdeevskii wells), Shetukhinskya (Bolshoi kluch well), Samarka (Sodovy well), Chuguevka (Luzki well), Leninskoe (Narzanyi well) and Gornovodnoe (Fig. 1). The Shmakovka group is located between the ancient metamorphic Khanka massif to the west and

Figure 1. Location of high PCOz waters in Primorye Wells: 1,2-Medvejii & Avdeevskii, 3-Bolshoi kluch 4-Sodovyi, 5Narzanyi, 6-Luzki, 7-Gornovodnoe

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the main geological structures of the Sikhote-Alin in the east. The last volcanic activity was during the Pliocene. The Shetukhinskya wells are located in the Western Sikhote-Alin Mountains. In this region Upper Permian-Triassic tuffs, volcanic rocks, sandstones and granites are predominant. The Samarka wells are located in the central SikhoteAlin Mountains and host rocks are mainly sedimentary. The Leninskoe wells are situated close to the main fault of the Primorye-Central SikhoteAlin fault system. This territory is located on a few tectonic blocks and consists of Triassic-Jurassic sedimentary rocks and Lower Cretaceous volcanic rocks. The Chuguevkay group of waters is located in a volcano-depression, which is filled by Lower Cretaceous terrigenous rocks and Paleogene lavas. The Gornovodnoe wells are situated in the Eastern Sikhote-Alin volcanic belt where Upper Cretaceous andesites and basalts are widespread. It is important to note that Miocene-Pliocene volcanic activity was typical of these studied areas. 2 RESULTS AND DISCUSSION 2.1 Chemical composition of waters The high P C 0 2 waters of Primorye are cold, with temperatures ranging between 5.8-12.7 'C. Total mineralization of waters ranges from 200 to 3000 mg/l. The main anion is HC03* with a maximum concentration of c 2000 mg/l. Ca2+is the dominant cation, but in some springs (Na+K)+ can play an important role in the element budget. pH varies from 3.8 to 6.04 and is supported by a high partial pressure of C02. Hydrogen and oxygen isotopes ratios of C02-reach waters are close to the world meteoric line and similar to river waters of Primorye indicating local meteoric recharge (Shand et a1 1995). All waters are supersaturated with clay minerals, albite, K-feldspar and quartz. At the same time they are undersaturated with calcite. A comparison of the previous chemical data of these waters (Shand et al. 1995, Chudaeva et al. 1999) and our data from the 1999 field season allows us to conclude that there are no significant differences in the concentrations of most chemical elements. Some differences, however, were found for Cu, Ga, and Ge. Their concentrations were higher in the samples taken in 1999. The maximum contents of REEs are found for the Bolshoi kluch (Shetykhinskaya group) and Gornovodnoe wells. In Bolshoi Much water the light REE (La/Lu=17) are predominant and they are associated with maximum concentration of Be, Al, Zn, As, and Sb. In the Cornovodnoe wells, La/Lu = 0.6 and the maximum heavy REE correspond to the

maximum of Mn, MO, and Bi. Distributions of microelements are irregular in Medvejii spring (Shmakovka group) and we found rather high Sc (>2 pg/l), V (0.5 pg/l), Cu (7.9 pg/l), Sr (1.1 mg/l) and U (2.6 pg/l). In Avdeevskii spring, of the same group, the concentrations of CO and Rb are relatively high. In Bolshoi kluch spring the contents of Se, Y, and Ba are high compared to other wells of high PC02 waters. The highest concentrations of Li (>990 pg/l), Ba (553 pg/l), V (1.2 pg/l), Cr (16.4 pg/l), Cu (>10 pg/l), Ga(21.5 pg/l), Ge (5.6 pg/l), Br (81.7 pg/l), Sr (2.8 pg/l), and relatively high contents of MO and Cs are found in the Niznii Luzkii well. The Gornovodnoe well yielded the highest concentrations of Mn (3.2 pg/l), MO (2.6 pg/l) and Cs (7.8 pg/l). The highest concentrations of Mn (3.2 pg/l), MO (2.6 pg/l) and Cs (7.8 pg/1) were obtained from the Gornovodnoe well. In Narsanyi spring (Leninskoe group) concentrations of Rb, Se, Sr and Cs are very high. Additional measurements included In (7-14ng/l), Hg (the maximum of 15.6-72.4ng/l, found in Avdeevskii spring), T1 (0.5-69.6ndl maximum in Avdeevskii spring), and T1 (up to 214.9ng/l) in the Shetukhinskaya group of waters. These distinguishing characteristics of the chemical composition of high PC02 waters in Primorye, as we concluded earlier (Chudaeva et al. 1999), depend upon the composition of the host rocks, the partial pressure of CO2 and the residence time. Monitoring of this type of waters shows that they are rather stable in their chemical composition. 2.2 Gas composition The main gas component in high PC02 waters is C02, which may be as high as 98% of the total of all the gases (Bogatkov 1962, Shand et al, 1995, Chudaeva et al. 1999). The calculated partial pressure of CO2 for the Shmakovka group ranges from 0.63 to 1.6 atm; for the Shetukhinskaya and Gornovodnoe groups of waters this value is close to 1.8 atm. The concentrations of noble gases are given in Table 1. R/Ra= 3He?He(sample)/ 3He?He(ail)ranges from 2 to 4.66, and shows no relation to age or composition of the crust in the studied areas. It is well known that the R/Ra for MORB is 8. Thus, we can formulate a conclusion about the participation of mantle He. Using the equation: %Hem,,,1,=12.5 *He?He(measured)/3He?He(,ir), (Pinneker et al. 1999, we can roughly estimate the mantle percentage of He. In our samples this value ranges from 40 to 60%. For example in the Baikal rift (Tunka depression) 100% of He comes from mantle (Polyak et al. 1992, Pinneker et al. 1995). Comparison of these data with the helium isotopic ratio in Circum-

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magmatic chambers. As we mentioned above, magmatic lava of the Miocene-Pliocene ages is widespread in the studied areas. To understand the origin of volatile fluxes the U3He ratio is often used (Tolstikhin 1986, Prasolov & Tolstikhin 1987, Marthy & Jambon 1987, Polyak et al. 1992, Kharaka et al. 1999). These authors suggest different values of U3He from 107 to 10" for the mantle. Tolstikhin and Prasolov estimated C/3He as 2*107for the mantle. This value finds support in fumaroles and hot springs in Eastern Kamchatka and Iceland (Prasolov & Tolstikhin 1987). In instances of atmospheric influence, carbonates and organic matter have a high CI3He ratio of up to 10l2 (Pol ak et al. 1992). Kharaka et al. (1999) proposed a C/ He value for the mantle of 10". Marthy & Jambon (1987) provided data showing that volcanic glasses of MORI3 and inclusions have a U3He ratio close to 2* 109; they proposed this value for the mantle. We can use the dissolution rates of CO2 and He to explain these observed values of mantle CI3He. Dissolution of CO:! in water is higher than He. During degassing of magma COz, helium and other volatile gases rise via fractures and faults to the upper part of the earth's crust, and part of the CO2 could be dissolved in the interstitial waters where, finally, some of it could be lost in springs. In Table 1 ratio of U3He varies from 4.7 to 12 * 107, allowing us to propose mantle sources of CO2 in high PC02 waters of Primorye. This example of the mantle origin of CO2 in spring waters is not unique. We have described the Malki high PC02 waters in Kamchatka, which have magmatic CO2 (Chudaev et al. 2000). In Italy PC02 waters were discovered to

Pacific volcanic arcs shows that in the Kamchatka Kurile-Honshu-Ryukyu system R/Ra varies from 5.7-6.5 and mantle He reaches 80% (Poreda & Craig 1989). In the China platform, which is close to Western Sikhote-Alin, R/Ra is a very low 0.37-0.48 (Poreda & Craig 1989). The R/Ra for the studied area of Primorye occupies an intermediate position between a stable geological structure -platform and an active volcanic arc. The ratio of isotopes Ar and Ne is close to air and they are of atmospheric origin (Table 1). If we wish to understand the evolution of high PC02 waters in Primorye, we must determine the origin of the CO2 itself. Earlier we proposed that CO;! is magmatic (Shand et al. 1995). New data (Table 1) supports this idea. Take into consideration that g3C varies from - 4.19 o/oo to - 8.19 o/oo, which is too heavy to be of biogenic origin and too light to be from marine carbonate. As with 3He/4He, the values 613C are not dependent on the composition of the surrounding rocks. CO2 gas from the Kelua volcano (Hawaii) during the period of 1960-1985 had a 613C value of around - 3.4 o/oo. In kimberlite 613C is close to -5"/oo. 613C of magmatic origin has the value -8°/oo (Hoefs 1997). Our data on 613C are close to those for magmatic carbon, but this is indefinite because in the natural water's system fast fractionation for carbon was found (Wigley et al. 1978). We can get the same value of 613C in Primorye waters by using different proportions of carbonate and organic matter, but in this case the variations of 613C were more essential, because the studied areas, sometimes, have no carbonate in the crust. In our opinion, part of the carbon in CO2 gas is being released from long lived Miocene-Pliocene

Y

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have deep CO2 gas (Caboi et al. 1993). n a r a k a et al. (1999) proposed a significant role for mantle co2 in the springs Of the San-lh’ldreas fault in California* 3 CONCLUSIONS

1. Monitoring of the high P C 0 2 waters shows in their composition’ Some distinguishing characteristics were found for c u , Ga and Ge. Their concentrations were higher in the sam les from 1999. 2. He~He(,,,,~,) / 3HePHe(ai,)range from 2 to 4.66,

r

and are not related to the age and composition of the crust in the studied areas. In high PC02 waters we conclude the participation of mantle He is up to 60%. Ar and Ne have atmospheric origin. 3. The U3He ratio and 6I3C allows proposing a magmatic (mantle) origin for part of the carbon in CO2 gas. 4. Thus, the high PC02 mineral waters of the Primorye region were formed due to the following factors: the exogene-water component, and endogene-carbon and helium gas.

Marthy, B. & A. Jambon 1987. U3He in volatile fluxes from the solid Earth: implication for carbon geodynamics. 1987. Earth and Planetary Science Letters 83: 16-26. Pinneker, E.V.. Pissarskiv, B.I. & S.E Pavlova 1995. Helium isotopic data for the ground waters in the Baikal rift zone. Isotopes Environ. Health Stud. 31: 97-106. Malasia. Polyak, B.G., Prasolov, E.M., Tolstikhin, I.N., Kozlovzeva, S.B.. Kononov. V.I. & M.D. Khvtorskii 1992. Helium isotopes in fluids of the Baikal rift zbne. Imestia of Russian Acad. Sci. (geology) 10:18-33 (in Russian) Poreda, R. & H. Graig 1989. Helium isotope ratio in circum pacific volcanic arcs. Nature 338: 473 - 478. Prasolov, E.M. & I.N. Tolstikhin 1987. Juvenile gases-He, CO2, CH,: their relations and input to the fluids of earth’s crust. 1987. Geochemistry 10:1406-1414. Shand, P., Edmunds, W.M., Chudaeva, V.A., Lutsenko, T.N., Chudaev, O.V. & A.N. Chelnokov 1995. High PCOz cold springs of the Primorye Region. Proceeding of the 8fi international Symposium on Water-Rock Interaction: 393396. Rotterdam: Balkema. Tolstikhin, I.N. 1986. Geochemistry of helium, argon and rare gases. Nauka: Leningrad. Wigley, T.M., Plumme,r L.N. & F.J. Pearson 1978. Mass transfer and carbon isotope evolution in natural water systems. Geochimica et Cosmochimica Acta 42: 11171139.

ACKNOWLEDGEMENTS We acknowledge financial support from Russian Foundation for Basic Research (Project 98-05-5377).

REFERENCES Bogatkov, N.M. 1962. Mineral springs of Priamuria. Special hydrogeology of Siberia and Far East: 48-52. Irkutsk (in Russian). Chudaev, O.V., Chudaeva, V.A., Sugimori, K., Nagao, K., Takano, B., Matsuo, M., Kuno, A. & M. Kusakabe 2000. New data on the gas and waters composition in high PCOz mineral waters of Primorye. Proceedings of the conference on the fundamental water’s problem in 3rdmillennium: 280284.Tomsk (in Russian). Chudaeva, V.A., Chudaev, O.V., Chelnokov, A.N., Edmunds, W.M. & P. Shand 1999. Mineral waters of Primorye (chemical aspect). Dalnauka. Vladivostok. 160 p. (in Russian). Chudaev, O.V., Chudaeva, V.A., Shand, P. & W.M. Edmunds 1998. Geochemistry and origin of the two groups of mineral waters in South Kamchatka. Proceedings of 33rd conference of SITH: 30-33. Japan. Caboi, R., Cidu, R., Fanfani, L., Zuddas, P. & A.R. Zanzari 1993. Geochemistry of the high PC02 waters in Logudoro, Sardinia, Italy. Applied Geochemistry. 8:153-160. Hoefs, J. 1997. Stable Isotope Geochemistry. Springer: Berlin. Kharaka, Y., Thirdsen, J. & W. Evans 1999. Crustal fluids: CO2 of mantle and crustal origins in the San Andreas fault system, California. Geochernistry of Earth’s Surface: 515518.Rotterdam: Balkema.

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Origin of fluorine within the Afyon-Isparta volcanic district, SW Turkey: is fluorrnica the key? H.Coban, $.Caran & M.GOrmii~ Department of Geology Engineering, University of Siileyman Demirel, Isparta, Turkey

ABSTRACT: Fluorine is a well known element effecting drinking-waters from the Golciik potassic volcanics in the Isparta region (SW Turkey). The focus, here, is on the flourine data obtained from potassic, ultrapotassic and pyroclastic rocks, F-carrying mineral phases (mica, amphibole, apatite, titanite, glassy groundmass) and springs within the Afyon-Isparta volcanic district. Overall data indicate that high fluorine is a prominent aspect of this volcanic region and the main source of fluorine is fluormica (predominantly fluorphlogopite). The results show a strong correlation between K and F, and indicate that other accessory minerals (apatite, amphibole, titanite) provide less fluorine than micas. High fluorine in groundwater discharging from the volcanics is attributed to abundance of micas.

1 INTRODUCTION Due to the effects of high flourine contents in drinking water from the Isparta city (SW Turkey) to the human health, particularly on the serious dental and medical problems, the majority of previous literature has been associated with its origin and geochemistry of waters (Ayhan 1983, Ozgur et al. 1990, Pekdeger et al. 1992). High flourine was dealt with the Golciik volcanism in Isparta. However, its exact origin has not been fully understood , and no studies have been carried out to investigate the role of mineral-chemistry. The study covers the Neogene volcanics of the region between south of Afyon and Isparta known as Isparta Angle (Fig. 1). In the volcanic district, to determine the water-rock interaction with respect to the F-element, and to find out its genesis; three kinds of material were analyzed, volcanic rock samples (potassic / ultrapotassic / pyroclastics), F-bearing mineral phases and springs. F- analysis of rock samples was carried out in ACME analytical laboratories in Canada by specific ion-electrode method. F- analysis of constituent F- carrying minerals of the volcanic rocks were obtained from polished thin sections of the analyzed rocks in CYGNUS Consultant Company in Canada by means of an automated Jeol JXA-89OOL electron microprobe using ZAF correction; accelerating voltage 15 kV, beam current 20 nA; 120 s peak 477

count, and 60 s for each background count, 5 um beam diameter; F line analysed K-alpha; 3-sigma lower detection limit 0.08 wt% F; standard McGill fluorite containing 48.67 wt% F. F-contents of springs were determined by Scott Sanchis colourmetric method using Zr-Alizarine Red solution (TS-266) in Agriculture Center, Isparta, Turkey.

2 VOLCANISM WITHIN THE GEOLOGICAL FRAMEWORK Geological features of the Isparta Angle in the Lake District of Turkey have resulted from the highly complex convergence of the African and the European plates (Keller 1983). Neogene volcanic products were erupted on a wide region covering Paleozoic metamorphic rocks, Mesozoic carbonates, ophiolites and Tertiary sediments between Afyon and Isparta (Fig. 1). They include potassic (trachytic / trachyandesitic) to ultrapotassic lavas and pyroclastics showing post-collisional intraplate alkaline character related to extensional neo-tectonic regime (Koqyigit 1984, Bilgin et al. 1990, Yagmurlu et al. 1997). The pyroclastics outcrop over wide areas in the region. The potassic and ultrapotassic lavas are seen in the form of cones, plugs and dykes protruding through the Tertiary sediment cover. Three types of ultrapotassic lavas were defined from the region as RPT (Roman Type), lamprophyres and lamproites

3 FLOURINE GEOCHEMISTRY Results on the fluorine are divided into two sections. Firstly, looking at the volcanic rock samples and focusing on F-carrying mineral chemistry, and secondly presenting spring-water analysis. 3.1 . Volcnriicsarid Minerzrl Phnses

Figure 1. Generalizied geological map of the study area (modified from Yagmurlu 1997). 1, Plio-Quaternary sedilnents; 2, Neogene volcanics; 3 , Mesozoic ophiolitcs; 4, Tertiary sediiiients; 5 , Mesozoic carbonates; 6, Paleozoic metamorphics; 7, locations & analysed rock samples; 8, spring-water saniplcs; 9, boundary; 10, thrust; I, Golcuk; 11, Isparta; 111, Bucak; IV, Senirkent; V, Suhut. (Coban et al. 2000). Recent K-Ar data (Besang et al. 1977, Lefevre et al. 1983, Kazanci 1995) obtained on mica minerals from several of these volcanics and tuffs, give ages ranging from 8.6 2 0.2 to 1.3 f. 0.13 Ma (Upper Miocene to Pliocene).

F values in ultrapotassic rocks are from 0.1 to 1.39 wt% (lamproites 0.78-1.23; lamprophyre 1.06-1.39; RPT 0.1-0.72 wt% (Table 1). They range between 0.2 and 0.4 wt% in potassic rocks, and 0.05-0.18 wt% in pyroclastics. So, relative enrichment of the ultrapotassic/potassic and depletion of the pyroclastic rocks in F is obvious. The data display strong correlation between K20, P205 and F. The following explains the similar results and examples from various localities of the world. Foley et al. (1986) recorded that the ultrapotassic rocks are a compositionally heterogeneous group of rocks in which volatile species (H20, CO;?, F, C1, S02) are more abundant than in less alkaline rocks. He indicated that the fluorine is the most abundant in lamproitic rocks and RPT rocks contain the lowest amounts of fluorine amongst the ultrapotassic rocks. Similarly, the effect of fluorine is greater in ultrapotassic rocks than in other mafic rocks because fluorine correlates positively with K content (Aoki et al. 1981). Some examples of fluorine contents in ultrapotassic rocks are 2000 to 5400 ppm for West Kimberley, Australia (Jaques et al. 1986), 5900-7600 ppm for the Leucite Hills, USA (Kuehner et al. 1981). According to data (Table 2), the main F- source minerals in the rocks are mica, apatite, amphibole and titanite. The F- values in phlogopite range from 2.24 - 5.25 wt%; biotite 1.23 - 2.47 wt%; apatite 1.98 - 3.45 wt%; amphibole 1.23 - 1.57 wt%; titanite 0.68 -1.67 wt%. No fluorine was determined in glassy groundmass. The percentage of the micas implies that they are the main rock forming minerals, whilst the others are the accessory phases. Phlogopite is characteristic mica mineral in ultrapotassic rocks while biotite is in potassics. Data also indicate that the deficiency of F in some ultrapotassics such as RPT is due to relatively low amounts of phlogopite in the rocks, and F- contents in phlogopite from ultrapotassics are fairly high. So, presence of fluormicas signify a real enrichment of fluorine in volcanics. Aoki & Kanisawa (1979) and Aoki et al. (1981) reported the fluorine contents of phlogopite, amphibole and apatite in various types of basalts from continental and oceanic regions, and they emp-

478

Table 2. Average F contents of F- carrying miiierals withiii the volcanics.

Table 1. F, K 2 0 & P 2 0 5contents (wt %) of volcanics

1 2 3 4 5 6 7 8 9 10 11

0.33 0.40 0.20 0.28 0.37 0.16 0.18 0.05 1.39 1.06 1.14

5.36 5.13 4.91 4.63 5.19 4.63 5.28 5.14 6.83 7.06 6.45

0.79 0.30 0.80 0.24 0.30 0.24 0.34 0.25 2.81 2.23 1.55

12 15 7 11 12

25 32 30

12 13 14 15

1.10 0.78 1.23 1.16

6.51 6.82 7.01 6.42

1.35 0.65 1.18

22 17 25

16 17 18 19 20 21

0.68 0.10 0.59 0.13 0.72 0.08

1.30 9.60 0.59 9.04 0.31 9.62 0.93 6.51 0.24 8.88 1.14 5.82 0.16

23 15 5 13 5 20

t rach . $, .g 5 G

2.47pj 3.11,Jj 1.57pj 3.17iIj lid 1.54(,, na 1.37,lj 3.45Clj nd 1.76(,, na 1.43,,, 3.15,,, lid I1 1.23(,, 4.63,,, na 2.45(1, nd 4.12,,, na iia 1.98(1! lid I11 lid 4.87,,, nd 2.33(1j lid IV 1.12,2, 2.24,,, nd 2.45(,, 1.67(,, nd 2.84(1, nd 3.1 O.68,,, V lid 2.14,;j na 2.440j 1.39(lj na 5.25(3, 1.23(,, 2.32,,, 1.44il, I. 7-15 3-8 Cl->SO;- for anions. These relations are considered representative of the waters that evolve mainly by CO:! solution and water-rock interaction. A discussion of the water-rock interaction processes is beyond the scope of this preliminary analysis of the data set, and has been addressed in other volcanic regions (Gislasson & Eugster 1987a, 1987b, Gislasson & Arn6rsson 1993). When the waters composition is modified by seasalts spraying, the chemistry is expressed in equivalent units as Cl->HC03->SO;' for anions. In some islands it was possible to relate the more frequent

Figure 1. Concentration range for each major species in sohtion. The maximum, minimum and the median are plotted.

Figure 3. Relationship between chloride and sodium, interpreted as a result of sea-salts spraying.

3 DISCUSSION

482

able to assess these differences due to lithologic control. The influence of the aquifer lithology can also be shown by plotting the Mg2+content as a function of Ca2' (Fig. 6 ) . The highest concentrations of these species occur on the basaltic aquifers comparing to more evolved volcanic rocks. The linear relation is more well-defined for the basic rocks, while in the evolved volcanics is more scattered.

4 CONCLUSIONS Figure 4. Plot of electrical conductivity and bicarbonate content in groundwater.

The preliminary analysis of a data set from groundwater chemistry of cold spring discharges from perched-water bodies at the Azores archipelago enables to define two major evolution trends: (1) waters evolving by CO2 solution and silicate weathering and (2) waters which composition is modified by sea-salts spraying. Waters are all low mineralized, which is expected in a short time of residence medium. Slightly differences can be assess between discharges from basaltic rocks, generally with a lower silica content and higher calcium and magnesium content comparing to aquifers made up of more evolved volcanic rocks.

wind directions to chloride content (Cruz et al. 1992), which shows the importance of seawater spraying over the islands. Springs discharging from more evolved volcanic rocks, as the central volcanoes of trachytic nature, present generally an higher content in Si02 and lower content in Ca2', comparing to more basic volcanic rocks occurring in central volcanoes and fissural zones of basaltic s.1. nature. In the Figure 5 the relation between SiOz and Ca2' is shown which en-

REFERENCES

Figure 5. Relationship between silica and calcium content in perched-water bodies.

Figure 6. Relationship between calcium and magnesium content in groundwater.

Carvalho, M.R. ,1999. Estudo hidrogeoldgico do maciqo vulcrinico de Agua de PadFogo (SLio Miguel-Aqores). PhD, University of Lisbon, Lisbon. Coutinho, R.M. 1990. Estudo hidrogeoldgico do maciqo das Sete Cidades. MSc, University of Lisbon, Lisbon Coutinho, R.M. 2000. Elementos para a monitorizaqGo sismovulcdnica da ilha do Faial (ACores): caracterizaqLio hidrogeoldgica e avaliaqLio de anomalias de Rn associadas a fendmenos de desgaselficaqLio. PhD, University of Azores, Ponta Delgada. Cruz, J.V. 1992. Hidrogeologia da ilha de Santa Maria. MSc, University of Lisbon, Lisbon. Cruz, J.V. 1997. Estudo hidrogeoldgico da ilha do Pico, Aqores, Portugal. PhD, University of Azores, Ponta Delgada. Cruz, J.V., Silva, M.O. & M.R. Carvalho 1992. Hidrogeoquimica das Bguas subterriineas da ilha de Santa Maria (Aqores). Geolis 6 : 121-135. Cruz, J.V. & R. Coutinho 1998. Breve nota sobre a importibcia das aguas subterriineas no arquipdago dos Aqores. Aqoreana 8: 59 1-594. Cruz, J.V., Coutinho, R.M., Carvalho, M.R., Oskarsson, N. & S.R. Gislason 1999. Chemistry of waters fiom Furnas volcano, S%oMiguel, Azores: fluxes of volcanic carbon dioxide and leached material. J. Volcanol. Geotherm. Res. 92: 151-167. Cruz, J.V. & M.O. Silva 2000. Groundwater salinization in Pico island (Azores, Portugal): origin and mechanisms. Environmental Geology 39: 1 18 1- 1 189. Custodio, E. 1975. Hydrogeologia de las rocas volchicas. Proc. 3rd UNESCO-ESA-IHA Symp. on Groundwater, 2369.

483

Custodio, E. 1978. Geohidrologia de terrenos e islas volcrinicas. Publication 128, Madrid:Centro de Estudios Hidrograficos and Instituto de Hidrologia. Freeze, R.A. & J.A. Cherry 1979. Ground water. New Jersey: Prentice-Hall. Gislasson, S.R. & H.P. Eugster 1987a. Meteoric water-basalt interactions. I: a laboratory study. Geochim. Cosmochim. Acta 5 1: 2827-2840. Gislasson, S.R. & H.P. Eugster 1987b. Meteoric water-basalt interactions. 11: a field study in N.E. Iceland Geochim. Cosmochim. Acta 51: 2841-2855. Gislasson, S.R. & S. Amorsson 1993. Dissolution of primary basaltic minerals in natural waters: saturation state and kinetics. Chemical Geology 105: 117-135. Langmuir, D. 1997. Aqueous environmental geochemistry. New Jersey: Prentice-Hall Lobo, M.A. 1993. Contribuiqzo para o estudo fisico-quimico e microbiologico da agua para consumo humano do arquipelago dos Aqores. PhD Thesis, Universidade dos Aqores, Angra do Heroism0 Macdonald, G.A., Abbott, A.T. & F.L. Peterson 1983. Volcanoes in the sea. The geology of Hawaii. Honolulu: Univ. Hawaii Press Paradela, P.L. 1980. Hidrogeologia geral das ilhas. Comun. Serv. Geol. Portugal 66: 24 1-256 Peterson, F.L. 1972. Water development on tropic volcanic islands type example: Hawaii. Ground Water 10: 18-23 Peterson, F.L. 1993. Hydrogeology of volcanic oceanic islands. In Y Sakura Y (ed) Selected papers on environmental hydrogeology, Selected Papers 4: 163-171. Hannover: Heise.

484

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Groundwater circulation at Mt. Etna: evidences from 3H contents

l 8 0 , 2H and

W.D' Alessandro & C .Federico Istituto Nazionale di Geofisica e Vulcanologia, Sezione di Palermo, via U. La Malfa 153, 90146 Palermo, Italy

A .Aiuppa, M .Long0 & F.Parello Dipartimento CFTA, Universitd di Palermo, via Archirafi 36, 90123 Palermo, Italy

P.Allard & P.Jean-Baptiste Laboratoire de Sciences du Climat et de 1 'Environnement, CEA-CNRS, GiflYvette, France

ABSTRACT: Groundwaters from Mt Etna and the local meteoric recharge were analyzed for their 8D and 6l80 and tritium contents. 8D and 8l80 values of groundwaters define the local meteoric water line (8D=88180+18).Rainwaters show a wide range for both 8D (-84 to -12 060)and 6l80 (-12.8 to -3.0 %) with the more negative values measured in colder periods and at higher altitudes. Isotopic data codirm that the &fferences in chemical composition existing between the SW and the E flanks of the volcano are mostly related to their peculiar hydrology. Waters collected in the eastern hydrogeological basin have mean recharge altitude sipficantly lower (410 m) than those of the northern (850 m) and south-western (1000 m) ones. Moreover the higher tritium contents indicate that the eastern basin has generally shorter and higher-grdent circuits. The longer residence times of the south-western basin allow the groundwaters to reach very high magmatic CO2and He contents. 1 INTRODUCTION

Mt. Etna, among the most active volcanoes in the world, hosts one of the greatest groundwater systems of the island of Sicily. A good knowledge of the Etnean aquifer's hydrology is very important for at least two main reasons: i) the quantification of the water resources for their correct management, ii) the estimation of mass and energy budgets of the volcanic system. Isotope hydrology represents a useful tool to evaluate hydrologcal circuits, recharge areas and transit times of groundwaters. In order to obtain a whole circulation model on Mt Etna, a detailed survey on both groundwater and rainwater isotope gemhemistry is thus requested. In this work, we present a data set of a1'0, 8D and tritium measurements on rainwaters monthly collected at variable elevation on Mt Etna. Moreover, several groundwater samples have been analyzed for their isotope composition in order to define the local meteoric water line. Finally, temporal variations on isotope composition of groundwaters have been observed on 14 selected water samples. 2 METHODS Rainwater samples were monthly collected by a network of 10 rain gauges located at various altitudes along the flanks of Mt. Etna (Fig. 1). Rain gauges were made of a PE funnel of 30-cm &ameter and a

PE container of about 30 liters, protected from direct sunlight. To prevent evaporation 250 ml of paraffin oil were added to the PE container. The D/H and l80/l6O ratios of water samples were determined by mass spectrometry following routine methods. The results are reported in 8 units 0 vs. V-SMOW, with a respective precision of If: 0.5 %O on 6D and k0.05 (LSCE) and *0.1 (INGV, CFTA) 060on 6l80 values. The tritium content of water samples was also determined by mass spectrometry at LSCE, following a routine procedure (Jean-Baptiste et al. 1992), after previous degassing of the samples and subsequent measurement of the accumulating amount of tritiogenic 3He. 3 DESCRIPTION OF STUDY AREA The Mt. Etna, located in eastern Sicily, is about 3350 m a.s.1. high and covers an area of about 1200 km2. It is an alkaline strato-volcano, which has grown on a thick continental crust made up of carbonate and terrigenous deposits of MesozoicPleistocene age in an area of intense geodynamic activity (Barberi et al. 1974). The composition of its products ranges from alkali-basalts to trachites although most lavas have a hawaiitic composition (Tanguy et al. 1997). Hydrogeological studies (Ogniben 1966, Ferrara 1991) have demonstrated a huge groundwater circulation system. Due to the high permeability of the vol485

portant wet air masses come from the Ionian Sea, thus maximum of rainfall is concentrated on the eastern flank, as the volcano itself induces the condensation. Data collected over thirty years (1965-1994) display the lowest values (400-700 mm) on the lower SW flank up to values of 1000-1200 mm at 600-700 m elevation on the E flank, with an average for the entire area of about 800 mm (Regione Siciliana 1998). At higher elevation (> 2000 m), snow represents a significant part of the precipitations for the greatest part of the year. 4 RESULT AND DISCUSSION

Figure 1 . Location of sampling sites and isohyet map. Squares indicate the rain gauges and crosses the groundwater sampling sites ofTable 2. Gray bold lines indicate the limits of the main hydrogeologic basins.

canic rocks, effective infiltration is about 75% of the total meteoric recharge (0.88 km3/a), while run-off accounts only for less than 5% ( O p b e n 1966). The volcanic aquifers of Etna are underlain by impermeable sediments (Miocene flysch to the NW and Pleistocene clays to the SE) that form a plane gently dipping to the SE (Ogniben 1966). On the basis of geophysical and geological studies, Ferrara (199 1) subdivided the whole area in three main hydrogeological basins (Fig. l), tributaries of the Alcantara (N) and Simeto (SW) rivers and the Ionian Sea (E). The volcano displays peculiar climatic conditions with respect to the Mediterranean climate of the surrounding areas, due to the high altitude and the geographic position. In fact, an upward variation of climatic conditions, from subtropical to cold, through temperate, is observed. Precipitations are strongly influenced by elevation and exposure. The most im-

1 . 1 Relationshp between 6D and 6"O Figure 2 shows the 6D and 6l80 values measured on 49 groundwater and 47 rainwater samples collected all around the Mt. Etna area. All samples plot between the global meteoric water line and the East Mediterranean meteoric water line. Groundwaters display a fair positive relationship plottin on a straight line defined by the equation 6D=865 80+18 (1?=0.95). This linear relation representing the local meteoric water line (LMWL), confirms the results of previous authors (Anza et al. 1989, Allard et al. 1997). Rainwaters e h b i t a wider range for both 6D (-84 to -12 %) and 6l80 (-12.8 to -3.0 "60) with the more negative values measured in colder periods and at fugheraltitudes. Deviations from the LMWL are due to evaporation processes in the hot season or to different origin of Table 2. Variability of 6l*O (Ym) and tritium (TU) values of the Etnean groundwaters. Sampling points of the SW basin are in italics. Sample N. Ilice 12 Guardia 10 s 21 11 Bongiardo 9 Pontefmo 12 Petrulli 4 S.Giacomo 13 12 S.Paolo V m 11 13 Romito Valcorrente 13 Currune 12 AcquaGrassa 12 Cherubino 12

Table 1. Description ofthe rain gauges and measured 6l80 (%) and tritium (TU) values Rain alt. sea mm . 6"O TU gauge m min max vwm N . min max ARO 350 12 563 -9.2 -2.0 -6.5 5 0.0 7.5 CAT 190 3 602 -9.8 -1.1 -6.2 4 2.9 6.0 FON 2 0.2 905 -9.3 -1.7 -5.3 5 3.1 6.5 MAL 950 26 510 -12.9 -0.6 -8.0 4 3.5 8.7 NIC 670 11 831 -10.7 -2.1 -7.3 5 2.6 7.0 PRO 1820 15 874 -11.2 -3.1 -9.2 4 5.1 8.9 SLN 1740 17 890 -9.8 -2.9 -8.7 6 2.7 7.8 TDF 2940 15 549 -11.2 -5.1 -9.9 2 5.6 13.7 VER 570 17 579 -11.2 -0.6 -7.2 4 4.0 8.8 ZAF 730 8 1453 -10.0 +1.9 -7.6 7 0.0 6.5 Sea= distance from the sea in km; mm = yearly precipitation height; vwm = volume weighted mean

min -7.8 -6.9 -7.2 -6.6 -8.2 -7.5 -7.6 -7.2 -6.8 -9.1 -7.7 -8.9 -8.6 -8.9

max -7.2 -6.3 -6.7 -6.0 -7.7 -7.6 -7.0 -6.7 -6.5 -8.7 -7.3 -8.5 -8.3 -8.4

mean

U

-7.44 -6.72 -6.94 -6.15 -7.95 -7.55 -7.29 -6.95 -6.70 -8.90 -7.50 -8.75 -8.51 -8.72

0.14 0.18 0.15 0.22 0.18 0.05 0.17 0.16 0.10 0.13 0.11 0.09 0.10 0.14

N . min. max. 1 14.9 8

6.9 14.2

9 4

12.0 16.8 15.0 19.7

1

6.5

10 11

0.3 2.7 1.8 4.0

Table 3, 6l80values (%) and mean recharge altitude (MRA) of the Etnean groundwaters Area Meas. Patemo area 43 Rest SW basin 39 zLiEema area 9 RestE basin 28 N basin 14

486

rnin -9.0 -9.1 -7.8 -7.8 -8.7

rnax -6.2 -6.1 -5.8 -5.4 -6.5

mean -8.0 -8.2 -6.7 -6.6 -7.8

U

MRA(m)

0.68 93W180 0.73 10W200 0.61 45M160 0.56 41W150 0.60 85W160

Figure 2. 8"O vs. 8D diagram of natural waters collected in the Etnean area.Crosses refer to rainwater samples while circles to groundwater samples. The global meteoric water line (GMWL) is shown in gray and the East Mediterranean meteoric water line (EMMWL) in white.

the air masses that discharge their rain on the northern flank of the volcano. This latter process is more marked for the rain gauges VER and MAL that probably collect a higher proportion of rainwaters derived from the less evaporated western Mediterranean (Thyrrenian sea, 30-35 km away). Aiuppa et al. (2001) attributed to the same process the hgher content of marine-derived chloride in the rainwaters collected in the gauges of the northern flank (VER, MAL, FON) with respect to the southern one (NIC, CAT, ARO). 4.2 Temporal variations of 6180 values

The annual range of 6l80 values of rainwaters, collected in the Etnean area over the period Oct. 1997Sep. 1998, is comprised between 6.1 (TDF) and 12.3 (MAL) 8 units. Although it is worth to note that samples collected in the period of greatest precipitation (Oct. to Jan.) display a much narrower range (24 8 units). Figure 3 also highlights the different origin of the rainwaters collected in the northern flank of the volcano with respect to the southern ones. The Qfference is especially evident in the spring period. The samples collected in the S gauges display, in fact, minimum values in February followed by a sharp increase 6l80 values (- 8 6 units) in March when the N gauges display their minimum values.

Figure 4. Sampling altitude vs. wvmean 6'*0 values (Oct. 1997 - Sep. 1998) of rainwaters collected in the Etnean area. White symbols gauges below 1000 m altitude and black symbols above.

The other 4 gauges, not shown in Figure 3, follow an intermediatetemporal pattern. In contrast, groundwaters display very little variations in time (0.3 to 0.7 8 units) indicating an isotopically well-mixed groundwater reservoir. 4.3 Variationsof 6"0 values with altitude

Figure 4 shows an inverse relationship between the yearly volume-weighted mean of 6l80 for each rain gauge versus sampling altitude. Rain gauges located at altitudes less than 1000 m define a line whose slope (-2.7060h) strongly differs from that of high altitude gauges (-l.O%/km). This difference is probably due to the fact that our rain gauges failed to entirely collect the snow, whch represents a considerable part of precipitations at higher altitude. The missing proportion of snow was quantified in 35% at TDF, 20% at PRO and 10% at TDF and its presumably very negative isotopic composition (not measured) could account for the different isotopic gradient. The data points of the low altitude gauges display a good alignement in Figure 4 (8=0.99) with no sig nificant differencedue to geographic distribution. Table 3 displays the mean recharge altitudes (MRA) of the Etnean groundwaters. Waters collected in the eastern hydrogeologic basin have mean recharge altitudes (MRA) that are significantly lower (410 m) than those of the northern (850 m) and south-western (1000 m) sectors. This result agrees with the fact that rainfall maxima on the eastern flank of Etna occur at much lower altitudes than on the remaining sectors of the volcano. 4.4 Tritium

Tritium content was measured both in rainwater and groundwater samples. Values of the former range fiom 0 to 13.7 TU with the highest values measured at higher altitude and on the northem flank of the volcano. Groundwaters display values from 0.1 to 19.7 TU. Samples collected in the SW hydrogeologic basin have values always lower than 4 TU while the remaining groundwaters always hgher than 6 TU.

Figure 3 . Temporal variation of 6l80 values measured in rainwater samples collected from 6 gauges in the Etnean area fiom Oct. 1997 to Sep. 1998. White symbols gauges of the northern flank of the volcano and black symbols from the southern flank.

407

40

0

1°10

0%

8

REFERENCES

MgmU 16 Figure 5. Magnesium content vs. tritium values measured in groundwater samples collected in the Etnean area. Mean values are used for groundwaters collected several times. Diamonds represent samples collected in the southwestern hydrogeological basin. The following samples are not reported in Table 2: Acquarossa (2.6 TU), Aqua Difesa (2.3), Gulli (l.O), Faro Pennisi (6.1), S. Leonardello (8.5), Zummo (14.6). 0

groundwater of this area is eventually limited by the exsolution of a free gas phase.

8

Figure 5 indicates that there is a fair inverse relationship between tritium and magnesium content. Magnesium has been proved to be a good index of the degree of WRI in the Etnean aquifer (Brusca et al. 2001). High T-low Mg values refer to groundwaters with shallow and rapid circulation (less then 10 years). On the contrary the groundwaters of the SW basin with low T and high Mg contents have transit times greater than 50 years allowing a more intense WRI. Two groundwaters (Petrulli and Ponteferro) display intermediate Mg contents and TU values higher than present-day precipitations. These waters have probably a longer circuit (about 20 years) with respect to the rest of the groundwaters of the E basin, in fact, they display also more negative values of PO. 5 CONCLUDING REMARKS

Significant hfferences between the groundwaters of the eastern and the southwestern hydrogeologmil basins of Mt. Etna have been reported. The water chemistry of the SW basin, despite to the rather uniform hawaiitic rock composition of the Etnean aquifers, indicates a higher degree of WRI (Aiuppa et al. 2000, Brusca et al. 2001). These waters display also the highest magmatic CO2 and He contents (Allard et al. 1997, D’Alessandro et al. 1999). T h s study, presenting new data on the isotope composition of rainand ground-waters of the Etnean area, confirms the hypothesis that these differences are principally due to the peculiar hydrologic conditions of the two areas. The eastern basin, in fact, together with a higher meteoric input, &splays generally shorter and higher-gdent circuits. This leads to shorter residence times confirmed by the higher tritium contents. The longer residence times of the southwestern basin, especially the Paterno area., allow the groundwaters to dissolve greater amounts of magmatic CO2 rising through faults. This in turn lowers considerably the pH values enhancing WRI processes. Carbon dioxide pressure built up in some

Aiuppa, A., Allard, P., D’Alessandro, W., Michel, A., Parel10, F., Treuil, M. & M. Valenza, 2000. Mobility and fluxes of major, minor and trace metals during basalt weathering at Mt. Etna volcano (Sicily). Geochim. Cosmochim. Acta 64: 1827-1841. Aiuppa, A., Bonfanti, P., Brusca, L., D’Alessandro, W., Federico, C. & F. Parello, 2001. QuantifLing the environmental impact of volcanic emissions: Insight fiom the chemistry ofrainwater in the Mt. Etna area. (Sicily). Appl. Geockm. (in press). Allard, P., Jean-Baptiste, P., D’Alessandro, W., Parello, F., Parisi, B. & C. Flehoc, 1997. Mantle-derived helium and carbon in groundwaters and gases of Mount Etna, Italy. Earth Planet. Sci. Lett. 148: 501-516. Anzl S., Dongad, G., Giammanco, S., Gottini, V., Hauser, S . & M. Valenza, 1989. Geochimica dei fluidi dell’Etna. Miner. Petrogr. Acta 32 : 23 1-251. Barberi, F., Civetta, L., Gasparini, P., Innocenti, F., Scandone, R. & L. Villari, 1974. Evolution of a section of the Africa-Europe plate boundary: paleomagnetic and volcanological evidence fiom Sicily. Earth Planet. Sci. Lett. 22: 123132. Brusca, L., Aiuppa, A., D’Alessandro, W., Parello, F., Ailard, P. & A. Michel., 2001. Geochemical mapping of magmatic gas-water-rock interactions in the aquifer of Mount Etna volcano. J. Volcanol. Geotherm. Res. (in press). D’Alessandro, W., Inguaggiato, S., Federico, C. & F. Parello, 1999. Chemical composition of dissolved gases in groundwaters from Mt.Etna, Eastern Sicily. Proc. .5Ih Internt. Symp. on Geochem. of the Earth’s Surface, Reykjavik, Iceland, 16-20 Aug. 1999, 491-494. Ferrara, V. 1991. Modificazioni indotte dallo sfruttamento delle acque sotterranee sull’equilibrio idrodinamico e idrochimiCO dell’acquifero vulcanico dell’Etna. Mem. Soc. Geol. It. 47: 619-630. Jean-Baptiste, P., Mantisi, F., Dapoigny, A. & M. Stieve nard, 1992. Design and performance of a mass spectrometer for measuring helium isotopes in natural waters and fbr low-level tritium determination by k e ingrowth method. Appl. Radiat. Isotop. 43: 881-891. Ogniben, L. 1966. Lineamenti idrogeologici dell’Etna. Rivista Mineraria Siciliana 100-102: 151-174. Regione Siciliana 1998. Climatologia della Sicilia. Palermo, Tipografia Priulla. Tanguy, J.C., Condomines, M. & G. Kieffer, 1997. Evolution ofthe Mount Etna magma: Constrains on the present Ming system and eruptive mechanism. J. Volcanol. Geotherm. Res. 75: 221-250.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2

Groundwater geochernistry in the Broken Hill region, Australia P.de Caritat Cooperative Research Centrefor Landscape Evolution and Mineral Exploration (CRC LEME), cl, Australian Geological Survey Organisation, GPO Box 378, Canberra ACT 2601, Australia

N .Lavitt CRC LEME-Present address: Centrefor Water and Waste Technology,School of Civil and Environmental Engineering, University of New South Wales, Sydney NSW 2052, Australia

D .Kirste CRC LEME, cl- Australian Geological Survey Organisation, GPO Box 378, Canberra ACT 2601, Australia

ABSTRACT: Groundwater geochemistry is being evaluated as a potential tool for mineral exploration on the western, sediment-covered margins of the Broken Hill Domain. 74 water samples collected from sedimentary and bedrock aquifers show that the groundwaters are dominantly brackish, of Na-CI-SO4 and Na-CI types, of near-neutral pH, and mildly oxidizing to strongly reducing. The waters are mostly saturated with respect to calcite, and have high and variable bicarbonate alkalinity values. This suggests that a source of protons exists within this system to drive alkalinity generating reactions. An excess of SO4 above what can be explained by gypsum dissolution suggests that oxidation of sulfides, and possible hydrolysis of metal cations, may provide the additional SO4 and protons. This is postulated to then lead to the dissolution of silicates and carbonates and the release of bicarbonate. 1 INTRODUCTION The Broken Hill region in southern Australia is host to a world-class Ag-Pb-Zn ore body and numerous smaller Pb-Zn, Cu-Au and other mineralizations. Extensive mineral exploration of the outcropping basement over the last 100 years has had limited success in discovering other major economic deposits. Consequently, exploration focus is now shifting to those parts of the Broken HIII and Olary Domains in the Curnamona Province (New South Wales and South Australia) (Fig. 1) that are covered by up to 200 m of weathered Cainozoic sediments of the southern Callabonna Sub-basin. In this context, the development of tools for exploration under cover is crucial to the economy of the region and Australia. Here, we report on preliminary results of application of groundwater geochemistry and interpretation of water-rock interaction to explore in areas of thick sedimentary cover. We show that groundwater has the potential to acquire and preserve geochemical signatures from interaction with buried ore deposits.

2.2 Sedimentary cover The Cainozoic Callabonna Sub-basin is delimited to the east by the Banier Ranges, to the south by the

2 GEOLOGICAL SETTING 2.1 Basement The southern Curnamona Province is a highly prospective basement terrain with numerous highPartiCular1Yfor stratiform to quality target stratabound, multi-commodity (Pb, Zn, Ag, c u , CO,W, MO)mineralization (Leyh & Conor 2000).

Figure 1. Simplified geology of the Curnamona Province (dashed outline) and location of study area (frame). Modified after Leyh & Conor (2000).

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Olary Ranges, and to the west by the Flinders Ranges. It opens to the north toward the main Lake Eyre Basin. A basement high, the Benagerie Ridge, extends northwards from the Olary Ranges, dividing the Callabonna Sub-basin into two elongate troughs. Where it overlies the Curnamona basement, the basin fill consists of the Paleocene-Eocene Eyre Formation (sandstone, carbonaceous clastics and conglomerate), Miocene Namba Formation (grey, green and white clay, fine-grained sand and carbonate, with minor conglomerate) and undifferentiated Quaternary sands (red and yellow-brown sand and sandy clay, with gypsum and carbonate palaeosols); various indurated horizons have been recognized (silcretes, ferricretes) (Callen et al. 1995). 2.3 Surface regolith Regional regolith-landform maps have been produced for the Curnamona (Gibson 1996) and the Broken Hill (Gibson & Wilford 1996) map sheets. These maps show that the study area is covered mainly by alluvial and aeolian sediments, with minor in-situ regolith. Gypsum and carbonates are common minerals in the regolith.

Figure 2. Location of Callabonna groundwater samples (subdivided into 4 groups). Heavy and light solid lines are major and secondary roads, respectively. Dashed line is the outline of the Broken Hill Domain. BH = Broken Hill.

2.4 Hydrogeology

set of electrodes. Bicarbonate alkalinity and Fe2+, S2-, N03- and NH3 concentrations were also measured in the field. In the laboratory, the samples were subjected to a comprehensive analytical protocol, including pH and EC, major anions and cations, trace elements (including Au) and stable isotopes (Caritat et al. 2000).

Recharge zones within the study area comprise the Olary Ranges to the south and the Barrier Ranges to the east. The basin is up to 200 m deep under the Mundi Mundi Plain. The main aquifers are likely to include a series of palaeochannels located in the central part of the study area (Dobrzinski 1997), coarser units from the Eyre Formation (especially basal pebbly levels), underlying saprolite (where permeable) and fractured bedrock. Toward the centre and northeast of the study area, deep Jurassic aquifers of the Great Artesian Basin (GAB) are encountered in bores. Groundwater flow is assumed to be to the north from the Olary Ranges and to the northwest from the Barrier Ranges. Deep GAB aquifers have a potential south-southeast flow direction (Waterhouse & Beal 1978). Some bores in the study area have a subartesian potential head. The Benagerie Ridge will influence the subsurface flow paths. The discharge zone is Lake Frome (Fig. l), a normally dry, salt lake.

4 RESULTS The groundwaters are circum-neutral (pH from 6.4 to 8.3), strongly reducing to oxidizing (Eh from -383 to +219 mV), and brackish (total dissolved solids from 2030 to 25,400 mgk). Waters are mainly of the Na-CI-S04 (25 of 74 waters), Na-Cl (20) and Na-Mg-Cl-SOd (10) types, and have temperatures between 20 and 31°C (Figs 3A,B). The waters are subdivided into 4 groups (Gpl to Gp4), on the basis of their hydrogeological setting (see Fig. 2): 0 Gpl: shallow groundwaters recharged from the Olary Ranges, southern part of the study area; 0 Gp2: palaeochannel groundwaters, central part; 0 Gp3: deeper groundwaters influenced by mixing with GAB water, northern part, and; 0 Gp4: shallow groundwaters recharged from the Barrier Ranges, eastern part. Gp2 waters are the most saline. Gpl and Gp4 waters have the lowest temperatures (medians of 22.7 and 23.9"C), whilst Gp2 and Gp3 waters are warmer (26.9"C and 28.3"C). Gpl and Gp4 waters

3 METHODS Groundwater samples were collected from 74 boreholes (Fig. 2) following extended pumping to ensure that water representative of aquifer conditions was being retrieved. Temperature, electric conductivity (EC), pH, Eh (redox potential), and dissolved oxygen were monitored and measured in the field with a 490

Figure 3. Piper diagram (A), Eh vs Temperature (B) and calcite vs gypsum saturation (C) diagrams.

Figure 4. HC03 vs SiOz (A), SO4 vs Ca (B) and fluorite saturation vs Ca (C) diagrams.

are more oxidizing (Eh = +42 and +26 mV) than Gp2 and Gp3 waters (-201 and -155 mV).

strongly reducing conditions of Gp2 and Gp3 waters suggest greater residence times, consistent with their location further away from the ranges. The high salinity and the ion make-up of these waters (mainly Na, C1 and SO4) may be attributed to the availability of soluble salts such as gypsum and halite in the soil profiles of this semi-arid landscape. In order to understand the water-rock interaction processes based on major ion concentrations, it is

5 DISCUSSION The salinity, temperature and redox state of Gpl and Gp4 waters are consistent with low residence time groundwaters typical of recharge zones. The

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pertinent to understand the carbonate system dynamics, in particular the saturation state of carbonate minerals and the alkalinity of waters. The circum-neutral pH values suggest that the system is effectively buffered. Most water samples are saturated with respect to calcite (Fig. 3C), a common regolith mineral here. Weakly acidic rainwater promotes dissolution of these widely available minerals, leading to rapid attainment of calcite saturation during recharge. Were this to be the only source of bicarbonate, buffering would lead to very similar alkalinity values throughout the system. However, bicarbonate alkalinity values vary widely (2-20 mmol/L overall) (Fig. 4A), particularly for Gpl and Gp4, and independently of calcite saturation. Whereas it is not unusual for waters to show such wide variation in alkalinity, the consistent state of calcite saturation (Fig. 3C) shows that saturation must have been achieved by the additional release of bicarbonate, in response to increased proton activity. Dissolution of carbonates and silicates in the regolith generates alkalinity and consumes protons. However, a source for these protons other than dissolved CO2 needs to be identified. A clue to this is revealed when examining the relationship between Ca and SO4 (Fig. 4B). Gypsum, common in the soil profiles, provides a ready supply of SO4 to the system. However, the amount of SO4 in the system can clearly not be explained by gypsum dissolution alone, even when considering potential Ca depletion by cation exchange reactions (which NdCI ratios suggest are not significant). Further, given that carbonate dissolution would readily supply Ca to the system, an additional major source of SO4 is required. The mineralized Olary and Broken Hill Domains clearly contain significant ore deposits, many of which comprise sulfide-rich ores. It is suggested that oxidizing waters (eg, Gpl and Gp4) readily interact with sulfide minerals, releasing SO4 and protons to the system. This has the potential to be exploited as a regional h ydrogeochemical vector to mineralization. Further, hydrolysis of co-released metal ions, such as Fe2+,to oxy-hydroxides may further add to the proton activity. The cumulative effect is to promote bicarbonate-generating reactions such as silicate and carbonate dissolution. Silicate dissolution reactions are clearly substantiated for the Gpl waters (Fig. 4A), for which a close correlation exists between dissolved Si02 and HC03. Likewise, the lower Si02 values for Gp4 waters may be explained by a wider availability of carbonate minerals, which are more reactive than silicates. The scatter in data values provides an insight into the dynamics of these water-rock interaction systems. Clearly, fractured rock systems provide no uniform groupings of values, a fact that undoubtedly

reflects the variation in the chemical history of water and the apparent short residence time of these systems, relative to Gp2 samples. This wide variability is exemplified by the fluorite saturation vs Ca relationship (Fig. 4C). Fluorite is a common gangue mineral in mineralized terrains. Figure 4C clearly shows the wide scatter of fluorite saturation values.

6 CONCLUSIONS Major element chemistry of groundwaters in the Broken Hill region suggests that water-rock interaction processes take place in the regolith (eg, dissolution of carbonates, gypsum, halite) and in the (potenti a11y mineralized) fractured basemen t (eg, ox i dation of sulfides, precipitation of oxy-hydroxides, dissolution of fluorite). Using groundwater geochemistry to delineate zones of significant sulfide oxidation may be useful for mineral exploration in sedimen t-covered terrains. ACKNOWLEDGMENTS This research was financially supported by the Australian Government’s CRC Program, and the NSW Department of Mineral Resources. We thank our colleagues for discussions and particularly David Gray, Steve Hill, Ian Hutcheon and John Wilford for their reviews of this manuscript. PdC and DK publish with permission of the CEO of AGSO. REFERENCES Callen, R.A., N.F. Alley & D.R. Greenwood 1995. Lake Eyre Basin. In J.F. Drexel & W.V. Preiss (eds), The Geology of Soutk Australia, Vol 2 Tke Pkanerozoic. Mines and Energy South Australia, Bulletin 54: 188-194. Caritat, P. de, M.F. Killick, N.Lavitt, K.P. Tan & E. Tonui 2000. 3D conceptual modelling to aid mineral exploration in the southern Callabonna Sub-basin. MESA Journal (Quarterly Earth Resources Journal of Priinary Industries and Resources South Australia) 19: 46-47. Dobrzinski, I. 1997. Beverley and Honeymoon uranium projects. MESA Journal (Quarterly Eartk Resources Jouriial of Priirzary Industries arid Resources South Australia) 5: 9-1 1. Gibson, D.L. & J.R. Wilford 1996. Broken Hill Regolith Landforms (1500 000 map scale). Cooperative Research Centre f o r Landscape Evolution arid Mineral Exploration, (CRC LEME), PertWCanberra. Gibson, D.L. 1996. Curnamona Province Regolith Landforms (1500 000 map scale). Cooperative Research Centre f o r Landscape Evolution and Mineral Exploration, (CRC LEME), PerthICanberra. Leyh, W.R. & C.H.H. Conor 2000. Stratigraphically controlled metallogenic zonation associated with the regional redox boundary of the Willyama Supergroup - Economic implications for the southern Curnamona Province. MESA Journal (Quarterly Eartk Resources Jouriial of Priniary Iiidustries and Resources Soutk Australia) 16: 39-47. Waterhouse, J.D. & J.C. Beal 1978. An assessment of the hydrogeology of the southern Frome Embayment. Mineral Resource Review, Soutk Australia 149: 9-21.

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The mineralised springs of the Marche and Abruzzi foredeep, central Italy: hydrochemical and tectonic features G .Desiderio & S .Rusi Dipartimento di Scienze della Terra, Universitd degli Studi “G.D’Annunzio”, Chieti, Italy

T.Nanni Dipartimento di Scienze della Terra, Universith degli Studi di Ferrara, Italy

P.Vivalda Dipartimento di Fisica e Ingegneria dei Materiali e del Territorio, Universitd degli Studi di Ancona, Italy

ABSTRACT: Some analogies between the Marche and Abruzzi mineralised waters from the Adriatic foredeep are observed, especially when considering the water chemical facies and the structural setting of the discharge zones. Particularly the saline waters are in relation with the presence of compressive tectonic structures, while the sulphureous waters are connected with the Messinian evaporitic deposits, with the tectonics and with the hydrogeological characters of the Plio-Pleistocene sequence. 1 INTRODUCTION In many areas of the Marche and Abruzzi foredeep, saline and sulphureous waters discharge from the Messinian and Plio-Pleistocene deposits and from the Quaternary deposits of the alluvial plains. The springs are structurally controlled, as they are very numerous in zones characterised by buried thrusts. In this study, the mineralised springs of the Marche region, already studied by the same authors (Nanni & Vivalda, 1999a; 1999b), and the Abruzzi mineralised springs, are compared. By extending the analysis to the Abruzzi sector of the Adriatic foredeep, our intent is to characterise the springs of this region and to find possible similarities in both the water chemistry and the structural setting of the discharge areas.

2 LITHOSTRUCTURAL SETTING The stratigraphic sequence of the Central Adriatic foredeep (Fig. 1) consists of: pre-Miocene deposits found in wells for hydrocarbon exploration and outcropping in the Apennines; Miocene sequence, in the Marche comprised of pelagic and hemipelagic marly-calcareous and marly sediments and by arenaceous, marly-arenaceous and marly-clayey terrigenous deposits of the Bisciaro, Schlier, preevaporitic deposits, Gessoso-Solfifera and Argille a colombacci formations. In the Abruzzi region the sequence is constituted of limestones, marly limestones and calcareous marls and of the Marnoso Arenacea Formation in the northern area and of the

Gessoso-Solfifera Formation south of the Pescara river. In addition there is the thick Pliocene and PlioPleistocene sequence, formed by marly-clayey and clayey-marly deposits with interbedded arenaceous and conglomeratic bodies which are more numerous in the southern Marche and in the Abruzzi region, that ends with the transitional arenaceous and conglomeratic arenaceous deposits. The whole sequence is closed by the alluvial terraced deposits which form the alluvial plains from the Apennine to the Adriatic coast. In the Abruzzi region, the Allochtonous with Pliocene slabs (Colata gravitativa) is of great importance and it is comprised of clays with interbedded marls, limestones with arenaceous turbidites, calcarenites, gypsum and evaporitic limestones. The thickness varies from 10 to 1000m. The foredeep, that in outcropping is formed by the Plio-Pleistocene deposits, is characterised by a structural setting typical of the Adriatic area, with folds and Apennine and antiapennine faults, generally not evident on the surface. The preorogenic Pliocene deposits, outcropping in the western part of the foredeep and in the coastal area of the Marche, are characterised by folds bordered by thrusts and interrupted by Apennine and antiapennine faults. Under the Plio-Pleistocene cover, as the geophysical studies prove (Ori et al., 1991), the pre-orogenic Pliocene has a similar setting, with a buried structure which sometimes reaches the base of the Plieistocene. Therefore, the Marche and Abruzzi foredeep is characterised by faults and thrusts always buried by the Plio-Pleistocene cover and

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Table 1. Analytical data of the mineralised waters of the Marche (1-8) and Abruzzi (9-12). 1: Isola del Piano; 2: Valzangona; 3: Vallone; 4: Moie; 5: Polverigi; 6: Aspio; 7: Tolentino; 8: Offida; 9: Vomano valley; 10: Cenerone d’Atri; 11: Pescara valley; 12: Sangro valley.

Figure 1. Geolithological scheme of the Marche and Abruzzi foredeep. Meso-Cenozoic (l), Miocene (2), Plio-Pleistocene (3) deposits; (4) Alluvial terraced deposits, Pleistocene-Holocene; (5) Thrusts in the Meso-Cenozoic and Miocene deposits; (6) Blind thrusts in the adriatic foredeep; (7) Faults in the Meso-Cenozoic and Miocene deposits; (8) Mineralised waters analysed; (9) Mineralised waters; (10) Geological sections.

outcropping only in some areas (for example Polverigi, Port0 San Giorgio in the Marche and from Vomano-Tordino rivers to Pescara river in the Abruzzi region). The Plio- Pleistocene sequence is of variable thickness proceeding from the northern

Marche to the Abruzzi region (Fig 2). In the area north of Cingoli-Mt Conero structure, the thickness of the cover is reduced (1000-2000 m). In this area, along the Adriatic coast between Pesaro and Ancona, there is a remarkable tectonic rising that

Figure 2. Longitudinal geological sections through Marche and Abruzzi foredeep. Pre-Miocene (l), Miocene (2), Pliocene (3) deposits; (4) Allochtonous with Pliocene slabs (Colata gravitativa) ; (5) Thrusts and faults; (6) Mineralised waters; (7) Borehole. See Fig1 for location of the sections

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causes the formation of the coastal ridges and the outcrop of the pre-orogenic Pliocene deposits. The ridges are bordered, toward the sea, by thrusts and are affected by an Apennine and antiapennine tectonics. In the area south of Cingoli-Mt. Conero structure toward Porto San Giorgio-Mt.Ascensione area, the Plio-Pleistocene cover has a thickness up to 5000 m. In the coastal zone, south of Mt. Conero, there is a rising that affects only the sediments of the basal Pleistocene, but, in the Porto San Giorgio structure, the pre-orogenic Pliocene sediments outcrop. In this case too the coastal ridge is bordered by a thrust mapped in the sea by the geophysical prospectings. In the adriatic region between Tronto and Moro-Sangro rivers, the cover is very thick (up to 6000 m). In the area between Moro and Trigno rivers the cover is remarkably reduced with a thickness of about 2000-3000 m, due to the rising of the limestone Apulian Platform (Casnedi et al., 1982).

3 THE MINERALISED WATERS The springs of the Adriatic central foredeep (Table l), are saline and sulphureous springs, generally emerging in fractured zones, in structural heights and in the hilly area east of the Apennine limestone ridge. Mineralised waters are frequent also in wells of the alluvial plain deposits whose substratum is certainly fractured owing to tectonic lines evident on the surface. In the discharge area, the saline springs form mud volcanoes while the sulphureous springs are often along streams whose waters are colored grey. The mineralised springs have a constant discharge, up to 3-4 l/min. The temperature follows the seasonal regime; the pH varies from 6 to 9; the Eh normally has negative values and the electrical conductivity varies fiom values lower than 1 mS to values higher than 900 mS. From the Piper diagram (Fig.3) we can observe the two groups of mineralised waters, saline and sulphureous, in the Marche, while the mineralised waters of the Abruzzi region are more diluted and fall in the central area of the diagram, with the exception of some waters very similar in chemistry to the Marche mineralised waters. All the waters processes through the clayey membrane. Some mineralised springs of the Marche region have a salinity even higher than marine water, while the Abruzzi waters are normally less diluted than seawater. The saline waters, with a chloride-sodium facies, normally emerge from the Plio-Pleistocene deposits, they are very numerous and with high salinity in the Marche region, but some cases are also present in

Figure 3. Piper diagram of the mineralised waters of Abruzzi region (I), Northern Marche (2), Southern Marche (3).

the Abruzzi region. In the Marche, saline springs are generally present in the Pliocene ridges where the lower Pliocene and Messinian deposits may outcrop. The discharges are often in the eastern front of the structures (for example Moie and Tolentino springs), but in the coastal ridges they are generally in the western fronts (for example Polverigi and Aspio springs) and in relation either with extensive faults or with the crossing between Apennine and antiapennine faults. Among the Abruzzi saline springs, the Vomano valley springs, emerging from Pleistocene sediments or from the alluvial plains, are situated in an area characterised by fronts of thrusts buried under the Plio-Pleistocene cover. Sometimes they may permit the outcropping of the Pliocene and Messinian deposits. The whole area is characterised by longitudinal and transversal faults. Therefore, the Vomano valley springs are similar to the saline springs of the

Figure 4. C1 vs Na diagrams. In the main diagram the mineralised waters of Marche and Abruzzi are represented. The plots fall close to the theoretical line of evaporationdilution of the seawater. Waters of Marche (I), Abruzzi (2).

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Marche. We believe that also for the Abruzzi springs the Pliocene compressive tectonics (Fig 5 DD’) permitted storage and squeezing up of the deep waters in the fronts of the thrusts; the high percentage of NH4 too, may prove deep circuits of waters which are not in contact with the oxygenated waters of the hydrogeological cycle. However, the Abruzzi waters have a lower salinity and this is due to the considerable presence of the arenaceous bodies which, by carrying fresh waters, produce their mixing with the deep waters which become diluted. The sulphureous springs, with a sulphaticcalcic facies, normally derive their chemistry from the leaching of evaporitic sediments. In the Marche region, they emerge first of all from the Messinian deposits (Gessoso-Solfifera Formation) and secondly from Plio-Pleistocene deposits. A large number of sulphureous springs emerges in the northern Marche, where the Plio-Pleistocene cover is reduced and the Messinian deposits often outcrop with a high thickness, or they are near the surface. As regards to the Abruzzi region, in the areas of Pescara and Sangro valleys, the cover has a small thickness (Fig. 5 EE’) due to the rising of the limestone Apulian platform. Consequently the evaporitic levels are always at a distance of about 100-200 m from the surface, so they have an important role in water mineralization. Therefore in the Pescara and Sangro valleys, there are springs emerging from the Plio-Pleistocene cover with high enrichments in sulphates and bicarbonates, very similar to the sulphureous waters of the northern Marche. However, in these waters, as in the Marche sulphureous waters emerging from Plio-Pleistocene deposits there are also enrichments in chloride and sodium. This phenomenon is due to mixing with the Plio-Pleistocene saline waters, as in the Marche waters. The recharge of the aquifers which supply the

sulphureous springs of the Marche and Abruzzi regions belongs to meteoric waters circulating in the arenaceous bodies present in the pre and postorogenic deposits.

4 CONCLUSION In the Marche and Abruzzi regions the presence of saline waters is in relation to a compressive tectonics, that permits storage and squeezing up of the deep waters in the front of the thrusts. In the two regions the sulphureous waters, connected with the leaching of evaporitc sediments, generally emerge where the Messinian deposits outcrop or are near the surface.

REFERENCES Casnedi, R. Crescenti, U & M. Tonna 1982. Evoluzione dell’Avanfossa adriatica meridionale nel Plio-Pleistocene, sulla base di dati di sottosuolo. Mem. Soc. Geol. It. 24: 243260. Nanni, T. & P. Vivalda 1999a. Le acque salate dell’Avanfossa marchigiana: origine, chimismo e caratteri strutturali delle zone di emergenza Boll. Soc. Geol. It. 1 18: 191-2 15. Nanni, T. & P. Vivalda 1999b. Le acque solfuree della regione marchigiana. Boll. Soc. Geol. It. 118: 585-599. Ori, G.G. Serafini, G. Visentin, C. Ricci Lucchi, F. Casnedi, R. Colalongo, M.L. & S. Mosna 1991. The PliocenePleistocene adriatic foredeep (Marche and Abruzzo, Italy): An integrated approach to surface and subsurface geology. E.A.P.G. ConJ Adriatic foredeep trip guidebook, May 26th -30”, Florence, Italy.

Figure 5.Geological cross-section. Meso-Cenozoic (l), Lower Pliocene (2), Middle Pliocene (3), Upper Pliocene (4), Pleistocene (5) deposits; (6) Allochotonous with Pliocene slabs; (7) Thrust and fault; (8) Mineralised waters; (9) Borehole. See Fig. 1 for location of the sections.

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Waterhock interaction in a karstified limestone sequence, south Galala, Gulf of Suez, Egypt A .A .El-Fiky Department of Environmental Sciences, Faculty of Science, Alexandria Univ., Egypt

M.N.Shaaban & M.A.Rashed Department of Geology, Faculty of Science, Alexandria Univ., Egypt

ABSTRACT: The present work discusses a numerical hydrogeochemical model for spring waters obtained from a carbonate-dominated karstified upper Cretaceous/lower Tertiary sequence, South Galala, Gulf of Suez, Egypt. Water flow takes place in an open system from recharge areas, through the infiltration of dilute precipitation within the karstified sequence, to discharge areas on the slopes of the southern Galala plateau at the vicinity of the upper Cretaceous carbonates. The proposed model, which is based upon mass balance and ion speciation concepts of water samples coupled with detailed fabric studies for the circumjacent rocks, has the advantage of elucidating the natural complexities of the waterhock system at the study area. Quantitative description of the dissolutiodprecipitation process within the studied sequence requires the consideration of the following factors: 1) dissolutiodprecipitation reactions and the saturation states of the recorded mineral phases, 2) gas exchange, 3 ) cation exchange, and 4) meteoric-marine water mixing. Cation exchange, mixing with marine water, and precipitation of the recorded calcite cement are all responsible for the chemical fingerprint, and the undersaturation state of calcite, of the collected spring waters.

1 INTRODUCTION There is a consensus, among sedimentary petrographers and hydrogeologists, that waterhock interaction has a major, if not the major, impact on the chemistry of groundwater and the circumjacent rocks. It is evident that geochemical modeling is essential if we are to cope with a precise evaluation of the progressive stages in the chemical evolution of the waterhock system. Several authors have provided some valuable insights into the waterhock interaction processes in rock sequences of different ages all over the globe (e.g. Plummer et al. 1990; Smalley et al. 1994; Bachu 1997; Sherman et al. 1999). James & Choquette (1984) described limestones that exhibit extensive dissolution features, due to the action of flushing meteoric water, as karstified. In the south Galala area at the Eastern Desert of Egypt, the upper CretaceousAower Tertiary sequence shows some evidence of karstification. This sequence consists of a complex mixture of lithologies although carbonates are the most dominated rocks (Kuss & Malchus 1989). Most conceptual waterhock modeling endeavors are qualitative. In an attempt to gather some practical sense, the present work provides a numerical modeling of modern meteoric water flushing of the upper Cretaceous/lower Tertiary

carbonate-dominated karstified sequence of the St. Paul and St. Anthony monasteries, south Galala area, Eastern Desert, Egypt (Fig. 1). This has the advantage of deciphering the natural complexities surrounding the waterhock system therein.

2 GEOLOGIC SETTING The St. Paul and St. Anthony blocks are subjected to intensive block faulting during pre-rift period. These blocks are separated and terminated by major fault zones (Fig. 2) which were active during different times of deposition (Bandel & Kuss 1987). However, fault systems possess the same vertical extension from the basement complex till the CampaniadMaastrichtian and they lie underneath the thick highly jointed lower Tertiary sequence. Topographically, the upper Cretaceous/lower Tertiary sequence at the St. Paul-St. Anthony area appears as a dissected plateau. The CenomanidTuronian mixed siliciclastic-carbonate sediments constitute the base of the upper Cretaceous strata. The Senonian time has witnessed dramatic variations in the style of sedimentation towards more calcareous facies. The oldest Tertiary strata reported therein belong to the middle Paleocene and they are made up of sandy limestones and marls. These middle Paleocene strata, in turn, 497

4

Comprehensive water sample sets were collected from eight different karstic springs at St. Paul and St. Anthony monasteries in June 1997. Temperature, pH, specific conductance, and alkalinity were measured in the field. Water samples taken for chemical analyses of major ions were filtered in the field using a 0.45 pm membrane filters. Meanwhile, samples chosen for cation determination were acidified in the field to pH < 2 using nitric acid. All water samples were refrigerated till laboratory analyses were conducted according to the standard methods (Eaton et al. 1995). The chemical analysis of major cations and anions, silica, phosphorus, and flouride were performed at the Faculty of Science, Alexandria University. Meanwhile, aluminium was analyzed using an atomic absorption spectrophotometer, model Varian- 10' at the National Institute of Oceanography and Fisheries, Alexandria, Egypt. The obtained data were subjected to geocheniical modeliiig techniques. Saturation indices of spring waters with respect to the different mineral phases were computed using WATEQFP. Geochemical Models that describe dissolutiodprecipitation reactions, mixing of meteoric water with marine water, and cation exchange were evaluated using mass balance equations in NETPATH (Plummer et al. 1994). The chemical analysis of sea water (Drever 1982) and that of dilute precipitation (reported in Awad et al. 1996) were used as reference in the discussed model. Rock samples were collected from the whole upper CretaceousAower Tertiary sequence. Petrographic work was carried out using a polarized microscope and scanning electron microscopy in order to detect the spatial distribution of the different petrographic features and mineral phases. Mineralogical identification were further checked using X-ray diffraction patterns and Alizarin-Red S tests for all examined thin sections.

Figurc I . 1,ocation m a p o f t h c studied arca

underlie unconformably siliceous limestones and chalky limestones of lower Eocene. The whole section is capped by nummulitic limestones of the middle Eocene. 3 HYDROGEOLOGIC SETTING

The Eastern Desert is located within the arid belt of Egypt, where rainfall represents the main source of groundwater recharge. The rainfall ranges between 10 and 25 mm/year and occurs mostly during the period from October to March. Karstic springs are distributed on the slopes of the southern Galala Plateau at the monasteries of St. Paul and St. Anthony. The recharge to these springs occurs through the infiltration of dilute precipitation via a complicated fracture patterns dissecting the carbonate plateau, These springs issue from the fractured limestone, chalk, phosphatized limestone, and mar1 of upper Cretaceous at elevations of +345m at St. Paul monastery and +410 m at St. Anthony monastery. The discharge of these springs ranges froin less than 5 m3/day at St. Paul spring to 100 m3/day at St. Anthony spring.

5

ROCK FABRICS

The fractured carbonate dominated upper CretaceousAower Tertiary sequence possesses a more or less hydraulic continuity. The flow of dilute precipitation took place in an open system from recharge areas, via fractures, to discharge springs that issue in the vicinity of the upper Cretaceous strata. Petrographic and mineralogic investigations reveal that low-Mg calcite, dolomite, quartz, francolite, hydroxyapatite, and clay minerals (montmorillonite, mixed layer smectitehllite, and kaolinite) are the main recorded mineral phases within the studied rocks. Most of the examined carbonates are biomicrites with preserved whole fossils and fossil fragments of nummulitids,

Figure 2. Geologic cross section of'the studied area

498 El-Fiky, A.A., M.N. Shaaban d; M.A. Rashed

MATERIALS AND METHODS

alveolinids, echinoderms, mollusks, bryozoans and benthic and planktonic foraminifers. True chalks display abundant coccoliths. Calcite cements are represented by calcite filling fractures and equant granular mosaics filling some intraskeletal voids. Chemical (diagenetic) alteration of fossils exhibits generic dependence. Partial dolomitization for some fossil fragments and matrix is observed but no pervasive dolomites are recorded. Silica may exist as either chert nodules or thin layers in sharp contact with the original fabric elements or even as partial silicification of fossils and matrix. There is no interrelationship between silica and dolomite fabrics. Partial phosphatization of both matrix and mollusca particles is observed in an upper Cretaceous phosphatized limestone bed where discharge may occur at the vicinity of this bed at the St. Anthony monastery. Porosity is represented by dissolution vugs of different shapes and sizes among all the studied rocks. Moreover, intraskeletal pores are frequent especially in lower Tertiary rocks. Some channels and fractures, sometimes filled with calcite mosaics, cut through the original rock fabrics. Intercrystalline microporosity is frequent in most of the studied rocks although it is hardly dominated over macroporosity. Some micropores are partly filled with some clay minerals.

6 RESULTS AND DISCUSSION Chemical analysis data of spring waters indicate that all water samples are slightly alkaline with pH values that range between 7 and 8. Water salinity shows wide variations throughout the study area ranging from 1152 to 9344 mg f'. The chemical composition of spring waters is dominated by Na(Mg)-C1-and SO4 ions (Fig 3). NdC1 ratio varies from 1.82 to 0.55 approaching that of sea water indicating possible mixing between recharge water (Na/CI = 1.33) and marine water (NdCl= 0.55). The calculated saturation indices of the studied spring waters indicate that most of the studied water samples are undersaturated with respect to calcite, dolomite, hydroxyapatite, and chalcedony while they are supersaturated with respect to clay minerals (Table 1). The chemical evolution of spring waters (of Na-(Mg)-C1-and SO4 composition) from dilute precipitation (of Ca-Mg (Na)-S04-C1-(HC03) composition) was evaluated by mass balance chemical modeling using NETPATH computer code. The chemical models were generated for eight springs in order to evaluate waterhock interactions and mixing with marine water. These models were examined on the light of the chemistry of different water samples, the mineral phases involved, and mixing.

Table 1. Saturation indices of some selected mineral phases in the studied water samples.

No.

Dolomite

-1.06 -1.27 -1.38 -0.19 -0.07 -0.89

-1.32 -1.95 -1.74 0.03 -0.93 -0.94 -1.41 -0.67

-1.02

-0.62

Hydroxyapatite -7.77 -7.90 -9.92 -3.48 -1.80

Chalced-

-0.249

-7.50 -6.88 -6.07

-0.092

1 Illite

According to our proposed model the enhancement of calcite dissolution just beneath the soil zone is related to the abundance of CO2 following the equation: CO2 + H20 + CaC03 = Ca2' + 2HCOy This results in the initiation of vugs and channels withi,] thc uppet parts of the vadose zone. It is evident that as these pores were found in response to water flow permeability was also enhanced. The preferential dissolution of certain fossils is related, in addition to original mineralogy and microstructure, to the evolution of the saturation state of the interacting water. However, with progressive dissolution of calcite the infiltrating water may evolve towards calcite saturation via an open system where carbonate particles react with water in contact with atmospheric gas of fixed Pcoz. The low recharge in the study area would favour rapid saturation with calcite close to the surface. As supersaturation states are maintained calcite precipitates (0.63-8.55 mmol calcitekg water) ending up to groundwaters undersaturated with respect to calcite. Dolomite might dissolve (0.195.69 mmol dolomitelkg water) as calcite

Figure 3. Trilinear diagram for the studied spring watet-s.

499 El-Fib, A.A., M.N. Shaaban & M.A. Rashed

Calcite

REFERENCES

precipitates, a process accompanied by degassing of C02. The infiltrating water would also dissolve fluoroapatite from the phosphatic limestone bed of CampaniadMaastrichtian age resulting in release of fluoride into solution followed by precipitation of hydroxyapatite (0.044-0.067 mmol/kg water). The dominance of Na in spring waters could be explained through exchange of aqueous Ca for adsorbed Na on clay minerals. It is evident, however, that chloride is a typical conservative tracer in groundwater flow. The noticeable variations among chloride contents in dilute precipitation and marine water make it effective in mixing calculations. According to the proposed model, the fraction of marine water which is mixed with dilute precipitation ranges between 0.01 and 0.16, suggesting a relatively higher degree of mixing. This is expressed in the increase of spring water salinities as well ay higher SO4 contents.

Awad, M.A., M.S. Hamza, S.M. Atwa & M.K. Sallouma 1996. Isotopic and hydrogeochemical evaluation of groundwater at Qusier-Safaga area, eastern desert, Egypt. Environmental Geochemistty and Health, 18: 47-54. Bachu, S. 1997. Flow of formation waters, aquifer characteristics, and their relation to hydrocarbon accumulations, Northern Alberta Basin. Amer. Assoc. Petrol. Geol. Bull., 81: 712-733. Bandel, K. & J. Kuss 1987. Depositional environment of the pre-rift sediments of the Galala heights (Gulf of Sues, Egypt. Berlin. Geowiss. Abk. (A) 78: 1-48. Drever, J.I. 1982. The geochemistry of natural waters. Prentice Hall, Englewood Cliffs, New Jersey. Eaton, A.D., L.S. Clesceri & A.E. Greenberg 1995. Standard mc'hods f o r the camination of water an.! wastewdor. 19"' Edition. James, N.P. & P.W. Choquette 1984. Limestones-The meteoric diagenetic environment. Geoscience Canada, 1 1: 16 1-194. Kuss, J. & N. Malchus 1989. Facies and composite biostratigraphy of late Cretaceous strata from northeast Egypt. In: J. Wiedmann (ed.): Cretaceous of the western Tethys. Proceed. 3"'Inter. Cret. Symp., Tubingen, 879-91 0. Plummer, L.N., J.F. Busby, R.W. Lee & B.B. Hanshaw 1990. Geochemical modeling of the Madison aquifer in parts of Montana, Wyoming, and South Dakota. Water Resources Research, 26: 198 1-20 14. Plummer, L.N., E.C. Prestemon & D.L. Parkhurst 1994. An interactive code (NETPATH) for modeling net geochemical reactions along a flow path version 2.0. US Geol. Surv. WRIR 94-4 169, 130 pp. Sherman, C.E., C.H. Fletcher & K.H. Rubin 1999. Marine and meteoric diagenesis of Pleistocene carbonates from a nearshore submarine terrace, Oahu, Hawaii. Jour. Sed. Res., 69: 1083-1097. Smally, P.C., P.K. Bishop, J.A.D. Dickson & D. Emery 1994. Water-rock interaction during meteoric flushing of a limestone: implications for porosity development in karstified petroleum reservoirs. Jour. Sed. Res., 64: 180-189.

7 CONCLUSIONS The carbonate dominated upper Cretaceous/lower Tertiary sequence at St. Paul-St. Anthony area, Eastern Desert, Egypt displays some evidence of karstfication. A numerical hydrogeochemical modeling, based upon mass balance and ion specition concepts of water samples coupled with rock fabric studies of the circumjacent rocks, has been used in deciphering the nature of waterhock interaction within the sequence. The chemical evolution of spring waters (of Na-(Mg)-C1-and SO4 composition) from dilute precipitation (of Ca-Mg(Na)-S04-Cl-(HC03) composition) passes through sequential stages depending on the dissolutiodprecipitation reactions and minerals saturation states, gas exchange, cation exchange and meteoric-marine water mixing. The predicted scenario of waterhock interaction in the study sequence suggests that most of the observed vug and channel porosities were initiated within a dissolution subzone at the upper parts of the vadose zone. However, water might evolve towards calcite supersaturation below this dissolution subzone, under the influence of low recharge, where carbonate particles reacted with water in contact with atmospheric gas of fixed Pc02. As calcite supersaturation was maintained calcite precipitates (0.63-8.55 mmol calcite/kg water) ending with groundwaters undersaturated with calcite. The cation exchange phenomenon of aqueous Ca for Na adsorbed on clay minerals is responsible for the higher sodium contents of the studied spring waters. The proposed model suggests also higher degrees of marine-meteoric water mixing resulting in higher C1 and SO4 contents.

500

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Hydrogeochemical characteristics of Hummar aquifer in Amman-Zarqa basin, Jordan A .R.EL-Naqa & K.M.Ibrahim Institute of Lands, Water and Environment, Hashemite University, Zarqa, Jordan

ABSTRACT: This paper highlights the geochemical characteristics of Hummar Aquifer (A4) which is a confined karstified aquifer. Fourthy-four samples were collected and analyzed for their chemical characteristics. Three distinct hydrogeochemical groups of water were recognized; the first water group is CaMg HC03 type representing the dissolution of dolomitic limestone from aquifer matrix. The second water group is Ca-Mg-HC03-SO4 representing mixing water between A4 and B2/A7 aquifers. The third water group is the Na-C1 waters that results from mixing of the groundwater with partially evaporated and back infiltrated irrigation water and due to excessive pumpage which rendered the aquifer more susceptible to slinization. The thermodynamic equilibria were used in the computation of aqueous speciation, precipitation andor dissolution of selected minerals in the aquifer matrix. The thermodynamic modeling represented by the saturation indices indicate that the dolomite and calcite minerals are precipitated and the groundwater dissolves gypsum from the aquifer. 1 INTRODUCTION The Amman- Zarqa Basin is one of Jordan's biggest basins, which consists of two main aquifers; the upper unconfined aquifer known as Amman-Wadi Sir (B2/A7) and the lower confined aquifer known as the Hummar (A4) (Table 1). The basin has been subjected to several investigations including Awad (1997). This work will be dedicated to study the hydrogeochemical characteristics of groundwaters extracted from Hummar aquifer and their interaction with rock matrix. The study area is about 650 km2, located in central Jordan (Fig. l), where almost half of the populations of Jordan live in the big cities of Amman, Zarqa and Ruseifa. The area is highly developed and therefore suffers from a continuous and increasing demand for water to meet the industrial, irrigation and domestic requirements. The topographic relief in the study area is moderate varying from 550 m along Zarqa River, to 900 m in the western part. The area falls within a semi-arid climatic zone and forms part of the Eastern Highlands of Jordan. The climate is generally characterized by dry hot summer and moderate winter. The temperature can rise in summer to 35" C and in winter drops to few degrees above zero. The rainfall distribution over the basin varies from less than 100 mm in the eastern part to more than 500 mrn in the western part of the basin as

shown in Figure 2. The average rainfall is about 160 mm in dry years and in wet years varies between 200-300 mm.

2 AQUIFER CHARACTERISTICS The B2/A7 aquifer system is one of the most and extensive limestone aquifer outcropping in the high rainfall areas where most of the recharge occurs. The A4 aquifer is a carbonate aquifer that consists

Figure 1. Location map of the study area.

501

Table 1. Simplified hydrogeological scheme of the Upper Cretaceous aquifers in the Amman - Zarqa Basin. Geological Rock Unit Muaqqar Amman-Wadi Sir (B2/47) Shueib(A5-6) _ Hummar (A4) Fuheis(A3) Na’ur (A 1-2) Kurnub

Hydrological Classification Aquiglude Aquifer __ __

. Aquiclude

ChaLk:marl ___ _-_- --50-150 Limestone - silicified limestone 100-250 _ _

__

_ _ Mar1

Figure 2. Isohyetal map of the Amman-Zarqa Basin.

-

_ 33-60

Permeability

8

(ds)

--”

--------

__

- --

10” - 3 ~ 1 0 ~ --

-_

-- -

- -

-

40-5o 30-50 30-150 50-300

Limestone Sandstone

mainly of about 45 m thick limestone and dolomitic limestone. The aquifer is characterized by karstification, which extends at depth, giving the formation high porosity and permeability affinities. It crops out in the northern and western part of the Amman - Zarqa basin forming a narrow band. In the study area the Hummar Formation occurs as a confined aquifer, which is sandwiched between two soft rock units of Shueib Formation (A5/6) above and Fuheis Formation (A3) beneath which act as an aquiclude and aquitard respectively. Structurally, the study area is controlled by the Amman-Zarqa flexure, which influences the groundwater flow. The general flow pattern is from west to east then to the north with steep gradients at some localities due to increased abstraction. The aquifer parameters and thickness vary according to local structure. Natural leak from the upper aquifer is not likely because the potentiometric head in A4 aquifer is higher than the water table of the upper aquifer. The domestic water supply is based on the groundwater extracted fkom the B2/A7 and A4 aquifers which together have a predicted safe yield of 87.5 MCM/y (Humphreys 1984). The actual yield of the basin is about 137 MCM/y. The current

-

_- -

Mar1 Dolomitic limestone

-

AquiferAguiclude Aquifer Aquifer

Saturated Thickness (m)

Lithology

_ _ _ _ 7.5 _ 8~10~ _

---------

5.3~10~ 4.48~10”

number of operating wells in the whole basin is about 662 wells. The area is highly developed and need much water to meet the industrial, agricultural and domestic purposes. 3 CHEMICAL ANALYSES Water samples from 40 production wells penetrating the A4 aquifer were collected. The temperature, conductivity and pH were measured in the field. Major cations (Ca, Mg, Na and K) and major anions C1, SO4, HCO3, CO3, NO3 were analyzed according to standard analytical techniques. The water types present in the study area were determined based on their chemical characteristics as shown in Table 2. Chemical analyses were plotted on trilinear plot as shown in Figure 3. Three distinct groups were deduced from the trilinear plot. The dominant water type is Ca-MgHCO3 representing the A4 dolomitic limestone of water found in deep wells. The second group representing the chemistry of B2/A7 and mixing of A4 with B2/A7 waters which is classified as CaHCO3 water. The third group represents the Sukhneh wells and the type of water is Na-C1. It is believed that continued re-irrigation from the

Figure 3. Trilinear plot of chemical data.

502

Table 2. Water chemistry data of the selected wells (concentration in mg/l). WellNo. AL1637 AL1638 AL1639 AL1640 AL1641 AL1643 AL1645 AL1711 AL1746 ALI821 AL1823 AL1825 AL1826 AL1827 AL1832 AL1836 AL1838 AL1842

E.C. pSkm 632 600 610 590 658 4000 898 878 800 523 520 800 595 630 510 679 1161 800

TDS

pH

Ca

Mg

Na

I(

CI

HC03

SO4

NO3

579.7 456.4 463.6 508.1 556.7 2601.1 556.7 617.4 608.2 414.4 392.2 581.6 498.9 498.9 405.3 515.5 822.5 533.6

7.63 7.29 7.51 7.61 7.60 7.1 7.32 7.31 7.9 7.52 7.45 7.59 7.58 7.48 7.75 7.26 7.24 7.20

156.8 52.2 42.8 44.8 77.8 215.8 49.6 91.8 84.2 50.8 48.8 68.0 58.6 57.6 46.0 60.8 124.4 49.8

31.1 29.4 34.2 31.1 20.4 107.1 29.3 19.9 26.4 27.9 26.8 30.0 30.0 37.8 29.2 34.2 27.9 30.0

23.7 29.9 34.5 33.8 29.7 503.7 78.0 47.4 51.8 15.2 10.8 46.0 19.3 20.9 0.0 27.1 66.5 40.9

3.1 2.7 2.0 4.7 2.7 7.8 6.6 5.5 3.9 1.6 1.2 7.8 3.1 3.1 0.0 2.3 10.2 4.3

48.6 43.3 51.8 49.0 63.9 867.3 141.6 96.6 113.4 29.5 20.6 69.2 39.1 42.2 26.3 51.8 142.7 52.9

255.0 241.6 266.6 266.0 237.3 313.5 184.2 280.0 240.9 256.2 251.3 323.9 277.6 299.5 261.1 316.0 323.9 283.7

28.8 24.0 18.2 309.6 33.1 523.7 49.9 25.0 69.6 24.0 13.0 18.7 15.8 36.0 8.60 15.4 34.1 24.0

32.6 33.3 12.6 99.0 43.2 62.2 17.5 51.2 18.0 9.20 19.7 18.0 10.4 1.80 7.30 7.90 93.7 48.0

al. 1976). By calculating the saturation index (SI) it is possible to identify minerals that regulate the chemistry of the investigated groundwater. On the basis of mineralogical composition of the aquifer matrix, the SI has been calculated with respect to the selected minerals such as calcite, dolomite and gypsum. In terms of thermodynamic considerations (Stumm & Morgan 198l), most of these waters are supersaturated with respect to calcite and dolomite as shown in Table 3. The saturation indices indicate that precipitation of calcite decreases the pH and removes carbonate from the groundwater, causing thereby further dissolution of dolomite of some parts of the aquifer. The calculated pC02 value of the groundwater wells exceeds the pC02 of the air which allows the water to dissolve dolomite in the aquifer (Drever 1988). At pC02 equals to 10” atm

groundwater aquifer causing a build up of salinity. Due to high temperature the irrigation water evaporating from the soil and precipitates salts above and below the soil surface. The salts are flushed and washed out from the soil causing infiltration of solutes into the subsurface thereby increasing the salinity of groundwater. The soil analyses indicates that NaCl and Na2C03 are the major salts in the soil (Nitsch 1990). 4 MINERAL EQUILIBRIA The groundwater chemistry was tested to find out the thermodynamic conditions of dissolution andor precipitation of mineral phases with the help of hydrochemical models using WATEQ4 (Plummer et

Table 3. Saturation indices of selected minerals within the A4 aquifer. Well No.

PC02

AL1637 AL1638 AL1639 AL1640 AL1641 ALl643 ALl645 AL1711 ALl746 AL1821 ALl823 ALl825 ALl826 ALl827 ALl832 ALl836 ALl838 ALl842

5.70E-03 1.21E-02 7.88E-03 5.64E-03 8.3SE-03 2.1OE-03 8.40E-03 1.29E-02 1.81E-02 7.43E-03 8.60E-03 7.85E-03 9.4 1E-03 9.40E-03 4.43E-03 1.71E-02 1.72E-02 1.65E-02

Total alkalinity (meq/l) 4.18 3.96 4.37 3.89 3.02 5.20 3.02 4.59 3.95 4.20 4.12 5.3 4.91 4.9 1 4.28 4.65 5.30 5.18

SI calcite

SI dolomite

SI gypsum

CdMg

0.743 -0.073 0.087 0.372 -0.220 0.180 -0.22 1 0.214 -0.1 10 0.161 0.076 0.430 0.219 0.220 0.359 -0.140 0.303 0.054

1.178 -0.03 1 0.429 0.517 -0.3 16 0.420 -0.3 16 0.1 18 -0.370 0.916 0.247 0.860 0.609 0.610 0.873 -0.140 0.309 0.2 10

- 1.902

5.04 1.07 7.53 2.29 1.02 1.22 1.69 2.78 1.93 1.10 1.10 1.37 9.16 0.92 9.47 1 .oo 2.69 1 .OS

503

-2.326 -2.530 -2.036 -2.590 -0.820 -2.059 -2.1 16 -1.720 -2.320 -2.592 -2.360 -2.137 -2.14 -2.805 -2.36 - 1.920 -2.480

dedolomitization may occur (Hounslow 1995). The high Ca/Mg ratio is usually accomplished by the dissolution of gypsum. 5 CONCLUSIONS The chemical characteristics of the natural waters coming from A4 aquifer which consists mainly of dolomitic limestone indicate that the aquifer is regulated by the dissolution of carbonate minerals within the aquifer matrix. Three groups of water were identified. The water-rock interaction conditions indicated that the groundwater was supersaturated with respect to calcite and dolomite minerals constituting the aquifer matrix. Dedolomitization may also occur. REFERENCES Awad, M. 1997. Environmental study of the Amman-Zarqa basin Jordan. Environmental Geology. 33:54-60. Drever, J.F. 1988. The Geochemistry of Natural Waters. Prentice Hall Inc., 388 pp., New York. Hounslow, A. 1995. Water Quality Data: Analysis and interpretation, Lewis Publisher, 396 pp. Humphreys, H. 1984. Monitoring and evaluation of the Amman-Zarqa aquifers, vol. 1 Amman Water & Sewerage Authority, Jordan. Lloyd, J.W. & J.A. Heathcote 1985. Natural Inorganic Hydrochemistry in Relation to Groundwater. Clarendon Press, 296 pp., Oxford. Nitsch, M. 1990. Soil Salinization in the Wadi Dhuleil and Wadi Arja Irrigation Project BGR FILE I0 306, BGR archive No. 106,4 15 pp. Hanover. Plummer, L.N., Jones, B.F. & A.H. Trusdell 1976. WATEQ4F a Fortran VI version of WATEQ, a computer program for calculating chemical equilibrium of natural waters, US. Geol. Sum., Water Resources Investigation, 76: 13-61. Stumm, W. & J. Morgan 1981. Aquatic Chemistry, New York, John Wiley and Sons, 780 pp.

504

Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Geochemical characterization of groundwaters from the Hyblean aquifers, South-Eastern Sicily R .Favara Istituto Nazionale di Geojisica e Vulcanologia- Sez. di Palermo, 90146 Palermo

F.Grassa & M .Valenza Dipartimento CFTA Universita di Palermo, 90123 Palermo-Italy

ABSTRACT: Groundwaters hosted within unconfined aquifers were collected from sixty-five cold and slightly thermal springs and wells. The studied aquifers are mainly constituted of a thick sequence of carbonates, Meso-Cenozoic in age. Basic volcanic intercalations, occurring during several episodes, are also present. Except for the samples located near the coast where seawater contribution plays an important role, the chemical composition of studied groundwaters is controlled by water-rock interaction. The proposed geochemical mass balance model suggests that Ca+Mg/HC03 groundwaters evolve toward a Na-HC03 type due to interaction with volcanic rocks, thus forming secondary minerals. pC02 values up to 0.1 atm suggest a contribution of a C02-gas source. The 613C~~1c values range between -15 to -8 %O vs PDB, indicating mixed organicinorganic carbon sources.

1 INTRODUCTION

The Hyblean Plateau (Fig. 1) is a carbonate platform (CP) with substantial intercalations of volcanic deposits (VD) that lies in the SE part of Sicily. Volcanism occurred intermittently from the Late Triassic through the early Pleistocene, in an environment characterized by dramatic eustatic and isostatic sea-level fluctuations and active tectonism. Eruption and deposition therefore occurred both under water and above the sea level, leading to the formation of distinctly different facies. The Hyblean Plateau occupies a territory about of 4500 km2 and is subdivided into four great hydrogeological basins. Because of their high permeability values and large extension, CP and VD units represent the most important local aquifers. Despite low-medium average annual precipitation ( - 6 5 0 mdyear), the volume of the infiltrating water in these unconfined aquifers can be estimated about 4 . 5 ~ 1 0 ~m3/year (10Sm3/km2).In the marginal parts of the basins, Miocene-Quaternary sediments made of marls, evaporites, biocalcarenites, and sandstones constitute only small coastal aquifers. This paper is focused on the main geochemical processes affecting the chemistry Of major constituents of groundwaters. Specific arguments are dis-

cussed about the dissolution of carbonate rocks, weathering of silicate minerals, cation exchange, carbon dioxide sources and mass balance modelling.

505

Figure Location map of the studied area. Depositis (m).2= ~ ~ n e - ~sdments, a t e3=&bonate ~ Platform (CP). Dashed line= watershed

2 GROUNDWATERS CHEMISTRY Water samples from 65 widespread springs and wells located within the Hyblean Plateau were collected. The samples show a large variability in the measured physico-chemical parameters. Temperature values range between 8 and 28OC, pH ranges from 6.7-to 9.3, while EC values range from 228 and 1800 pS/cm except for some samples which reach up to 15000. The stable isotopic composition of groundwater samples indicate that all the samples are of meteoric origin (Favara et al. in prep.). The 6D/6'*0 linear relationship is characterized by a slope less than 8, suggesting an evaporation process during infiltration, favoured by semi-arid climatic conditions. On the bases of the major chemical constituents, four aquifer types are recognized. (a) Coastal aquifers are characterized by high EC values, Na and C1 contents suggesting that their water chemistry is mainly controlled by marine contribution (up to 20%). (b) Waters circulating within the CP aquifers are dominated by Ca+Mg and HC03 ions, due to watercarbonate rock interaction. They are characterized by low salinity values < 750 mg 1-' (c) The chemistry of the water discharged from basal aquifers hosted within the VD rocks highlights a Na/HC03 feature. (d) The occurrence of CdS04 water type is linked to dissolution of gypsum rocks that are only in a restricted area.

3 GEOCHEMICAL PROCESSES 3.1 Coastal aquifer As can be seen from Figure 2 where Na content is plotted versus NdC1 ratio, the samples display two different behaviours. Groundwaters from coastal, gyspum and CP aquifers lie close to sea water NdCl ratio while waters discharged from VD aquifers show higher NdCl ratios. The latter group indicates that some interaction processes between these groundwaters and Na-bearing mineralogical phases occur. The former group suggests that the major source for C1 and Na in the sample with low Na contents is probably marine airborne sea water A direct marine contribute, should be take into account for the other samples containing Na contents higher than 1 meq 1-'. As previously inferred from isotope data, evaporation processes occurring during infiltration could increase the salinity in the groundwaters.

Figure 2. NdCl vs Na (in log scale).

3.2 CP acquifers Groundwaters hosted within the Meso-Cenozoic Carbonate aquifer are characterized by very low salinity values and may be considered the typical recharge waters. Ca and Mg are the dominant cations, while HC03 is the dominant anion, their sum being more than 75% of Total Dissolved Solid. Groundwater samples circulating within CP aquifers, show Ca/Mg ratios ranging between 0.68 and 18. In general Ca/Mg ratios in groundwaters in equilibrium with carbonate minerals are representative of the bulk rocks in which they flow. The partitioning of Ca and Mg in the local groundwaters could be usefkl to estimate the Ca/Mg ratio in the bulk solid. Wide range of Ca/Mg ratios observed in CP groundwaters suggests a great variability in the chemical composition of Meso-Cenozoic Carbonate sequence. Thermodynamic calculations by using of PHREEQC (Parkhurst, 1995) computer program have allowed to obtain the saturation indices of specific mineral forming host rocks and carbon dioxide partial pressure (pC0~).The saturation state with respect to calcite has been calculated from the logarithm of the ratio between the product of Ca and CO3 ion activities and the solubility constant of calcite (Ksp) as follows: Almost all the samples are close to saturation state with respect to calcite, but range from undersaturated to saturated with respect to dolomite. This indicates that the Ca content in groundwaters is controlled by calcite dissolution and seems to exclude the occurrence of weathering processes of Ca-

506

silicate minerals. Water samples having Ca/Mg ratio less than 1 are saturated both with respect to calcite and to dolomite. Starting from calcite and dolomite dissolution reactions and by using simple mathematic substitutions, we can compute CaMg ratios as solubility function of calcite (K,) and dolomite (b) constant by using of the relationship: Ca/Mg=K2J& Computed ratios between 15°C and 30°C by using PHREEQC (Parkhurst, 1995) program indicate an equilibrium temperature in the range 20-25°C. Despite the good agreement between calculated and measured ratios in the water samples, the occurrence of solid solutions e.g. magnesian calcite and Ca-rich dolomite, which cause changes in the solubility with respect to the pure phases, should be token into account. In the natural environment Ca may be substituted in the calcite lattice of the pure mineralogical phase by other metal ions like Mg, Sr, Fe to form a mixed solid phase. One of the most common substitutions is Mg which can be present to 20 mol% to form magnesian calcite. Magnesian calcite dissolution reactions considering calcite and dolomite as end-members can be written as follows: Cal-~Ca0,5Mgo,5),C03+ CO2 + H20= (1 - d 2 ) Ca2++ (3) d 2 Mg2++ 2HC03where x assumes values commonly ranging between 0 and 0.30. The diadochy of these elements in relevant amount can modify the solubility with respect to pure phase. As pointed out from Busenberg and Plummer (1989) the excess-mixing Gibbs free energy of formation of a calcite-dolomite solid solution can be calculated by applying the following expression:

derstand of equilibrium conditions between groundwaters and host rocks. 3.3 D groundwaters Groundwaters discharged from VD aquifers are mainly located in northern portion of the studied area. From the recharge zone to the discharge area, the groundwater evolution along the flow path from both physical and chemical parameters can be easily identified. As a consequence of longer residence time and water-rock interaction processes within fractured fresh and altered rocks, initial pH values close to neutrality (7.2-7.8) become alkaline up to 9.0-9.25, and TDS values increase progressively. From a chemical viewpoint, Na and HC03 are the dominant dissolved species and aqueous silica content become relevant (up to 50 mg I-'). VD groundwaters are supersaturated with respect to quartz, while are saturated and undersaturated with respect to chalcedony and amorphous silica respectively. The most common chemical reactions to explain the excess in Na with respect to Ca+Mg are (1) dissolutiodweathering of Na-bearing silicates like albite, to form kaolinite; (2) ion exchange reaction between bivalent dissolved ions and Na-clays that leads to a release of Na into the solution and uptake of Mg into the solid, (Appelo & Postma, 1993). As can be pointed out from mineralogical phase stability diagram considering Ndproton activity ratio and dissolved silica activity (Fig. 3) groundwaters from VD aquifer fall mainly in the stability field of kaolinite. This indicates the occurrence of weathering processes which cause the formation of secondary min-

Gess=x (I-X) [& + A1(2~-1)]

(4) where x represents the mole fraction of pure dolomite whereas & + A1 are two constants derived fiom asymmetrical solid-solutions model (Guggenheim, 1937). The solubility constant (Kss) in a CalciteDolomite binary solid solution as a fbnction of Mg moles (x) can be calculated as follows: lnK,, = x (I-x) R-'T-'[& + A1(2x-l)] + (I-x) In [K, (1-x)] + x In (x Kd) (5) Solubility constant values, calculated in the range 0.005 10-2atm.) are due to dissolution of CO2 formed by plant root respiration and microbial decomposition of organic matter in the soil zone. Weathering of basaltic rocks is known to produce large amounts of cations, bicarbonate and dissolved silica due to the high reactivity of basaltic minerals. The slopes of the respective cation v's bicarbonate ion (converted to equivalent units) can be used to estimate the relative amounts of each cation released during the weathering process (Fig. 1). The slopes are approximately: Na/HC03 = 0.20; Ca/HC03 = 0.44; Mg/HC03 = 0.36 (the sum of these equals l), indicating that practically all the total cation and bicarbonate increase is due to weathering of silicate minerals. The effects of evaporation in the soil zone are thought to be minimal because the chloride concentrations are very low (8- 12 mg/L) and are independent of cation concentrations. High dissolved Si02 concentrations (up to 80 mg/L) adds further evidence supporting the proposed hypothesis for silicate weathering control of the chemistry of basaltic aquifers.

to maturity and slow reactions involving clays and Fe-A1 oxyhydroxides which are the residue of the weathering process. The basaltic rocks in the Atherton area range from basanite to alkali olivine basalts (essentially defined by high Mg/Fe ratio). Mineralogy is predominantly: olivine, plagioclase, and clinopyroxene. A summary of whole rock chemistry for a basanite in the Atherton region reported by Stephenson (1995) is summarised as follows: Si02 47.2%; A1203 - 15.4%; CFe - 10.2%; MgO - 8.6%; CaO - 9.4%; Na20 - 3.8%). If all the A1 and Fe in these minerals is conserved in the solid phase through formation of secondary minerals, at least on time scales < 10,000 years (Eggleton et a1 1987, Nesbitt & Wilson 1992) then the notional weathering reactions involving Na, Ca and Mg bearing minerals can be written involving the three principal component minerals, olivine, clinopyroxene and plagioclase. Dissolution of the most reactive mineral (considered to be the Mg-olivine end-member, forsterite), reacts congruently with high CO2 bearing waters according to reaction 1:

Clinopyroxene is a Ca-Mg aluminosilicate which dissolves incongruently to produce kaolinite plus cations according to the following idealized stoichiometry :

+ 1.1H4Si04 + 3 .4HC03-

(2)

3 DISCUSSION

Plagioclase comprises a solid-solution between a sodium and calcic end- members, which occurs in sub-equal amounts in alkali basalts, with the Ca endmember being preferentially weathered relative to Na (i.e., the Na/Ca of the plagioclase increases as the reaction proceeds). Dissolution of the two end members, anorthite and albite can be written separately according to equations (3) and (4);

3. I Weathering reactions in basaltic aquifers

Anorthite: CaAlzSi2Os + 2C02 + 3H20 =

The cations and bicarbonate in the Atherton Tablelands aquifers are considered to be almost entirely derived by weathering of the basaltic minerals and this occurs primarily within the aquifer where waters are in contact with fresh basaltic minerals. The highly porous nature of the soils imply short residence times within the unsaturated zone. Interactions in the weathered zone are minimal due

= A12Si205(OH)4 + Ca2++2HC03'

Albite: 2NaAlzSi308 + 2C02 + 11H20

(3) =

The weathering reactions outlined above yield kaolinite as the solid end-product (except Mg-

530

clinopyroxene in basalts are considered to occur at grossly similar rates (Nesbitt & Wilson 1992), therefore we can use the relatively well established plagioclase weathering rate data to approximate the whole rock weathering rate. Although olivine is more easily weathered than plagioclase and clinopyroxene, it is a relatively small fraction of total mineralogy. Neglecting the higher rate of Mg2+ release rate from olivine dissolution results in a slight underestimation of the rate of release of Mg2+ (by assuming all Mg is derived at the weathering rate of Cpx), which in turn would tend to slightly overestimate groundwater residence time. There have been many laboratory studies of the kinetics of dissolution of a variety of rock forming minerals (see review by Sverdrup & Warfinge 1995). Laboratory experiments show that the weathering rates are highly pH dependent, but the rates are lowest (and occur at roughly the same rate at pH's between 5 - 8) which is the range of most groundwaters at Atherton. Although many earlier studies of laboratory reaction kinetics indicated that laboratory rates were much higher than field weathering rates, Sverdrup & Warfinge (1995) suggest that these data can be used to calculate field rates given certain chemical conditions and appropriate calibration. An estimated uniform linear dissolution rate constant derived from experiments on weathering of plagioclase of 10-17moles/cm2/sec (at 25°C) is used which is about 1 order of magnitude slower than laboratory rates. Gislason & Eugster (1987) derived an expression relating the residence time of groundwater to the concentration of ions:

Figure 2. Activity-activity diagram for the Na+-H'-Si02 system for the Atherton Tableland basaltic aquifers. The data show a tendency to approach equilibrium with respect to smectite minerals with increasing concentration. Similar trends are also observed for CaZ'-H'-Si02 and Mg2+-H+-SiO2systems.

olivine which dissolves congruently) and Na+, CaZ+, and Mg2+are all released on a 1:l equivalent ratio relative to alkalinity (as HC03-) which concurs with the observed amounts of cations released relative to alkalinity discussed in section 2 above. Mineral stability relationships (shown for the Na+-H+-Si02system in Fig. 2) show that the waters generally plot within the kaolinite stability field but evolve towards equilibrium with respect to smectite minerals with increasing cation and HC03- (or decreasing H 3 concentrations. Therefore, the above chemical reactions, and the proposed stoichiometries, are a reasonable representation of the processes in the Atherton basalt aquifers which can now be used to develop the residence time calculations presented in section 3.2. 3.2 Estimation of gross weathering r tes and groundwater residence time

The premise of the following discussion is that for a given rock weathering rate, uniform mineralogy, and other aquifer parameters being equal, higher solute concentrations correspond to longer groundwater residence times. Although this may be self evident in a qualitative sense, the estimation of more precise groundwater residence time from solute concentrations may appear to be fraught with too many uncertainties to be realistically achievable. However, the highly reactive nature of basaltic minerals, warm soil temperatures and high water fluxes soil pCO2's in this tropical environment can create an ideal setting to test the use of this approach in a more quantitative way against independent dating techniques. If the weathering rates of plagioclase and

t=-

c-CO K,* S

(5)

Where t = water residence time (time), C = initial concentration (mass/volume), CO = final concentration (mass/volume), K1 = weathering dissolution rate constant (mass/(area/time), and S = surface area of rock in contact with unit volume of water (aredvolume). Because the term S is difficult to estimate, the expression can be rewritten in terms of porosity and grainsize in the form:

where r = mean radius of rock grains (length) and no = effective active porosity. These calculations assume zero order kinetics for the dissolution process. Substituting an estimate of mean rock grain size of 2mm, and estimates of porosity of 0.1 and 0.2 into equation 6, we can derived a relationship between bicarbonate concentration and groundwater

531

be used as a proxy for age, and there is potential to calibrate it with CFC ages at the young end of the age spectrum ( ~ 4 years). 0 4 CONCLUSIONS

Figure 3. Groundwater residence time as a function of alkalinity for porosity of 0.1 and 0.2 respectively (see Eq. 6). Mean rock grain size is 0.2 cm.

residence time (Fig. 3) for the Atherton Tableland aquifers. Apparent groundwater residence time ranges from -5 years up to -120 years for the highest HCOi concentrations (2.25 meq/L). The data less than an alkalinity of 0.5 meq/L are ignored because they are predominantly from the metamorphic fractured rock aquifers where weathering rates are very slow and poorly known. Given that the apparent residence time is directly proportional to bicarbonate concentration, we can refine these estimates once better constraints on mean porosity and grain size can be obtained. Groundwater ages for the Atherton Tableland basaltic aquifers calculated from CFC-11 concentrations range from 2 - 40 years (the upper limit of CFC dating) Cook et a1 (2001), with older ages (lower CFC-11 concentrations) tending to correlate with higher HC03- concentrations (Fig. 4). Although one cannot directly compare CFC-11 groundwater ‘ages’ with chemical ‘residence times’ there is some indication that the bicarbonate ion can

Groundwater chemistry from basaltic aquifers in high rainfall areas of northern Australia is completely dominated by weathering of plagioclase, clinoproxene and olivine. Bicarbonate ion concentrations can be used to estimate groundwater flow rates using a simple model based on weathering rates, porosity and mean grain size. Groundwater residence times based on this model is of the order of 5 - 100 years. Calibration with independent methods such as chloroflourocarbons (CFC’s) means that routinely collected groundwater chemistry can be used as a proxy for quantitative groundwater residence time in certain circumstances. REFERENCES Cook, P.G., Herczeg, A.L. & K.L. McEwan 2001. Groundwater recharge and contribution to stream baseflow: Atherton Tablelands, Nth. Queensland. CSIRO Land & Water, Tech. Rep. (In press). Eggleton, R.A., Foudoulis, C. & D. Varkevisser 1987. Weathering of basalt: Changes in rock chemistry and mineralogy. Clay & Clay Minerals 35: 161-169. Gislason, S.R. & H.P. Eugster 1987. Meteoric water-basalt interactions. 11: A field study in N.E. Iceland. Geochim. Cosmochim. Acta 51: 2841-2855. Nesbitt, H.W. & R.E. Wilson 1992. Recent chemical weathering of basalts. Am. J. Sci. 292: 740-777. Stephenson, P.J. 1989. Northern Queensland (Part of Chap 3 East Australian Volcanic Geology) In: Intraplate Volcanism in eastern Australia and New Zealand (Ed. R. W. Johnson), Chap 3, Cambridge Univ. Press, pp. 89-96. Sverdrup, H. & P. Warfvinge 1995. Estimating field weathering rates using laboratory kinetics. In: Chemical Weathering Rates of Silicate Minerals (eds. A.F. White & S.L. Brantley), vol. 31, Chap. 11, Reviews inMineralogy, p. 485-541.

Figure 4. CFC-11 concentration versus alkalinity. Lower CFC-11 concentrations correspond to older ages (total age range -5 - 30 years).

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Water chemistry at Snowshoe Mountain, Colorado: mixed processes in a C o m o n bedrock A.R.Hoch Lawrence University, Appleton, Wisconsin, U S A .

M .M .Reddy U S . Geological Survey, Lakewood, Colorado, U S A .

ABSTRACT: At Snowshoe Mountain the primary bedrock is quite homogeneous, but weathering processes vary as waters moves through the soils, vadose zone and phreatic zone of the subsurface. In the thin soil, physical degradation of tuff facilitates preferential dissolution of potassium ion from glass within the rock matrix, while other silicate minerals remain unaltered. In the vadose zone, in the upper few meters of fractured bedrock, dilute water infiltrates during spring snowmelt and summer storms, leading to preferential dissolution of augite exposed on fracture surfaces. Deeper yet, in the phreatic zone of the fractured bedrock, Pleistocene calcite fracture fillings dissolve, and dioctahedral and trioctahedral clays form as penetrative weathering alters feldspar and pyroxene. Alkalinity is generated and silica concentrations are buffered by mineral alteration reactions. 1 INTRODUCTION Typically, soil and ground waters with longer residence times in a given rock become more solute rich, but hydrochemical data are often used to interpret solute sources without looking at the reacting minerals. To understand the chemical evolution of water chemistry from dilute snowmelt to spring and stream effluent, we studied water chemistry, hydrology and rocldmineral characteristics in the soil, vadose and phreatic mnes developed in an Oligocene welded-tuff near Creede Colorado. Field analyses and experimental results demonstrated the dominance of various solute producing processes in different hydrologic reaction environments formed from a single bedrock type.

2 BACKGROUND

aged, moderate-density Engelmann spruce (Picea engelmanni) and shallow regolith (0.0 - 1 .O m). The mean annual temperature is 2" C, and mean annual precipitation is about 50 cm, about half of which is in the form of snowfall. 2.2 Locatiori ailcl sampling sites Snowshoe Mountain is located in the eastern San Juan Mountains of southwestern Colorado (37'45' N, 57'30'W). Sampling sites are summarized in Table 1. 3 METHODS

Regolith or "rock pit" water samples were collected and sampled seasonally by zero-tension flow through a collection device described by Claassen et al. (1 986).

2.1 Geological Setting Table 1. Sampling sites from highest to lowest elevation

Snowshoe Mountain was originally chosen as a site for a dual watershed study because of the homogeneity of the bedrock (Bates & Henry 1922). It is the eroded central dome of a caldera complex, which was crystallized about 26.5 My before present (Steven & Lipman 1976). High-angle normal faults trend sub-parallel (N-NE) to the Deep Creek fault system in the rock comprised of augitsbiotite, quartz latite, welded tuff. This hard, siliceous tuff has been described in detail by Reddy et al. (1994) and Hoch et al. (1 999, 2000). The Seven Parks area is at an elevation of approximately 3500 m and is covered with even-

Elev. (in)

Rock Pit

3505

Zero-tension collector below 1 in of regolith.

Well 4 (vadose)

3487

13 i i i well, water level varies by 10 m.

Well 1 (phreatic)

3475

Site 48 (springs)

3399

20 ni well, water level is always near surface. Perennial spring, down gradient from well 1.

Deep Creek gauge 2698

533

Description

Site

Sampling station at the base o f the watershed,

Wells, springs, and surface waters were sampled seasonally from 1983 to 1998 and analyzed for cations, anions, and silica. (Hoch 1997). Core was sampled in 3 foot intervals for porosity and permeability tests; thin sections were made from rhodaniine dye impregnated billets for optical study of weathering and pore structure (Hoch et al. 2000). Interstitial glass in soil fragments was analyzed and verified by transmission electron microscopy (Hoch et al. 1999). Relative dissolution rates of Snowshoe Mountain tuff and augite were determined with a flow-through reactor system (Hoch et al. 1995, 1996). Mean residence time (MRT) of several points in the shallow aquifer was estimated from seasonal isotope variations in input water (rain and snow) and surface and ground water. The MRT can provide information about ground water recharge and flow paths in the subsurface. Seasonal variations of oxygen-18 in water are attenuated during transit in the shallow subsurface aquifer. This attenuation is related to recharge pathways and the MRT of the shallow subsurface aquifer. This work is in progress.

4 RESULTS AND DISCUSSION 4.I Regolith processes

Rock fragments undergo continual physical disaggregation in the soils. Internal fracturing related to freezing and thawing in soils disaggregates rock fragments with negligible mineral alteration. Phenocrysts such as plagioclase and augite remain unaltered, due to the short duration of wetting by dilute meteoric water (3 to 44 days observed residence time). Interstitial glass is more easily dissolved than reactive phases like biotite and augite (Fig. 1). Dissolution of a residual potassium-rich glass phase (< 5% by volume) produces an anomalous spike of potassium in waters collected from the regolith zone (Hoch et al. 1999).

Figure 1 . I n the soils, interstitial glass is preferentially dissolved and crystalline phases are left unaltered. (qtz = quartz; kspar = potassium feldspar).

rimmed with relatively inert iron oxide that protects the minerals from weathering. Dissolution of these minerals is not considered to contribute significantly to the chemistry of infiltrating waters. In fractures near the land surface that contain calcite and a pathway for water, calcite is dissolving but the tuff along the fracture lining remains unaltered, which suggests that the calcite dissolution has only recently begun (Hoch et al. 2000). 4 , 3 Phreatic zone processes

In Well 1 , water levels remain near the land surface year-round (estimated residence time 74 1-895 days in the subsurface). In contrast to the vadose well, we observe significant penetrative weathering deeper than the fracture surface (Hoch et al. 2000) and secondary mineralization in core samples. XRD analysis indicates that both dioctahedral and trioctahedral clays form after feldspar/glass and pyroxene, respectively (Fig. 3). These fill the rock fractures but do not totally impede permeability. (Hoch et al. 2000)

4.2 Vadose zone processes

In Well 4 we observe that water is in contact with the full length of the borehole only 2 to 3 months per year, during and after Spring snow melt. Petrography shows that augite is dissolved in open fractures in core samples from the vadose zone (Fig. 2). We also see preferential dissolution of augite in outcrops where previously subsurface fractures are exposed. Ca/Si and Mg/Si ratios from augite dissolution experiments mimic ratios in vadose zone waters, supporting the petrographic observation that augite dissolution is the dominant process there (Hoch 1997). Contrary to other studies that show biotite as a highly reactive phase (Blum et al. 1993), both biotite and hornblende in Snowshoe Mountain tuff are

Figure 2. Augite is preferentially dissolved in open fractures in core samples from the vadose zone.

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Figure 5 . Dissolved silica at sampling sites Figure 3. Photomicrograph showing secondary mineralization wif/ii/i an open fracture (long segmented material) and

penetrative weathering (bright areas in rock matrix). I n samples from the vadose and soil the bulk rock is unaltered.

4.4 Water chemistry

Potassium is notably enriched in the waters draining the regolith due to dissolution of freshly-exposed glass in the rock matrix (Fig. 4, Hoch et al. 2000). Deeper in the bedrock, fresh surface area production is not a dominant process, interstitial glass is not as available for dissolution and potassium concentrations remain low relative to other dissolved solids. Potassium may also be consumed in the soils by microbial activity. Dissolved silica concentrations are constant in ground waters and surface waters, suggesting that they are buffered by secondary formation of silica minerals or clays (Fig. 5). Solid silica phases were detected along with secondary clays in rock fractures, but SiO, activities do not correspond to any commonly reported minerals (Hoch 1997). Bicarbonate ion concentrations increase with weatheringhesidence time (Fig. 6). This results from dissolution of augite along with calcite in fractures where they are present. The pH (Fig. 7) increases only slightly in samples collected deeper than the soil horizon, due to strong buffering by bicarbonate in this pH range. Calcium ion concentrations (Fig. 8) show little increase in the rock pit (regolith) waters because the main phase dissolving is a potassiumrich glass. At

Figure 4. Potassium ion at sampling sites. Vertical bars indicate concentration range, short horizontal bars are means. Data are from Hoch ( I 997).

Figure 6. Bicarbonate at sampling sites

Figure 7. pH at sampling sites.

sites with longer residence times, calcium concentrations are greater as a result of augite and calcite dissolution. For all dissolved constituents shown, well 4 is intermediate between the rock pit and well 1, probably because its waters are a mixture of

Figure 8. Calcium ion at sampling sites.

535

infiltrating fresh water and concentrated, phreatic water.

deeper,

more

Hoch, A.R. 1997. Mechanisms and dissolution kinetics of clinopyroxene and a welded tiiff with implications for southwestern Colorado. Ph.D. Dissertatioti, University of Wyoming, Laraniie, Wyoming. Hoch, A.R., Reddy, M.M., & J.I.Drever 1999. The importance of mechanical disaggregation in chemical weathering in a cold alpine environment, San Juan Mountains, Colorado. Geological Society of Anzerica Bulletiri 1 1 I (2): 304-3 14. Hoch, A.R., Reddy, M.M., & M. Heymans 2000. Transient calcite fracture fillings in a welded tuff, Snowshoe Mountain, Colorado. Applied Geocheriiistry 15(10): 1495-

5 SUMMARY

From the rock pit at the top of the mountain to the stream gauge at the bottom (Table l), snowmelt input waters are modified in different ways depending on extent of infiltration and available phases for dissolution: - In the soil or regolith, dissolution of K-rich glass from freshly created rock surfaces is the prevalent chemical weathering process. - In the vadose zone, chemical weathering occurs on exposed fractures. The most commonly observed weathering process is the preferential dissolution of augite along open fractures, while other minerals, such as plagioclase, biotite and hornblende remain essentially unaltered. - In cores from the vadose well, calcite was absent, but in those from the phreatic well, calcite is a fracture-filling material and was observed to dissolve near the land surface. - In the phreatic well, alteration of feldspar, pyroxene, and interstitial glass to clays by penetrative rock weathering was the dominant process, possibly accompanied by reprecipitation of a silica-containing phase that buffers dissolved silica concentrations in springs and creeks.

1504.

Rcddy M.M., Claassen H.C., Rutherford D.W. & C.T. Cliiou 1994. Welded ttiff porosity using incrcury intrusion, nitrogen and ethylene glycol nionoethyl ether sorption and epifluoresccncc microscopy. Applied Geoc/ie/iii.sfry9: 49 1 499. Steven, T.A. & P.W Lipman 1976. Calderas of the San Juan volcanic field, southwestern Colorado. Uiiited States Geological Siirvey Professioiinl Paper 958, 35 p.

6 CONCLUSION

Although we identified separate and petrographically interesting dissolution processes in the regolith and vadose zone, spring and stream water chemistry is dominated by deeper, slower, penetrative weathering observed in the phreatic zone where residence times are longest. REFERENCES Bates, C.G., and Henry, A.J. 1922. Streamflow experiment at Wagon Wheel Gap, Colorado. United Slates Departtmnt of Agriculture Weatlier Bureau. Morithly Weatlier Review 17: 76 p. BItim, J.D., Erel, Y . , & Brown, K. 1993. 87Sri86Sr ratios of Sierra Nevada stream waters: Implications for relative mineral wcathering ratcs. Geocliirizica et C'os/iiiclii/iiico Acta 58(21122): 501 9-5025. Claassen, H.C., Reddy, M.M., & D.R. Halm 1986. Use of thc chloride ion i n determining hydrologic-basin water budgets - a 3 year case study in the San Jatin Mountains, Colorado, U.S.A. Jour/ial of Hydrology 85: 49-71, Hoch, A.R., Claasscn, H.C., Drcver, J. I., & M.M. Rcddy 1995. Dissolution stoichiometry and near-surface mineral chemistry o f augite after 1.5 years in a flow through reactor, with implications for reaction mechanism and watershed-scale mass balance calculations. Geological Society of America annual meeting, New Orleans. Hoch, A.R., Reddy, M.M., & J.I. Drevcr 1996. The cffect of iron content and dissolved O2 on dissolution rates of clinopyroxene at pH 5.8 and 25°C: preliminary results. Clieniical Geology 132(1-4): 15 1 - 156.

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Evidence for brine circulation in a groundwater discharge zone N.M.Howes & C.Le Gal La Salle Department of Physics, Chemistry and Earth Sciences, The Flinders University of South Australia, Australia

A .L.Herczeg CSIRO Land and Water, Adelaide, Australia

ABSTRACT: Groundwater pumping schemes in semi-arid south-eastern Australia intercept saline groundwater flowing into the River Murray and divert it to natural or man-made basins where super-concentrated brines (>350g/L) form. These brines infiltrate through the sediments of the basins, potentially mixing with the underlying groundwater (-5OgL). This study uses environmental tracers to provide field evidence to support the hypothesis of mixing of such brines with regional groundwater. A system of natural salt lakes is studied, as a long time-scale analogue to man-made evaporation ponds. Significant downward movement, and mixing with regional groundwater is observed in the geochemical signatures of the brines and an evaporation-leakage model suggests large amounts of groundwater through-flow for the salt lakes. Regional groundwater compositions at depths > 40m also indicate mixing with the brines. Geochemical mass transfer processes inferred from PHRQPITZ mixing model calculations suggest a substantial influence of re-mixing on brine compositions.

1 INTRODUCTION Stream salinisation, caused by rising water tables, is becoming a critical problem in the Murray-Darling Basin of south-eastern Australia, where its impact on water quality in the River Murray is worsening. Groundwater pumping schemes that aim to reduce groundwater mounds and intercept saline groundwater as it flows into the River Murray are a popular solution to this problem. Saline groundwater from these pumping schemes is disposed of into either naturally occurring or man-made basins and allowed to evaporate in the semi-arid conditions. A superconcentrated sub-surface brine, and often a surface salt crust, is subsequently formed. These brines infiltrate through the sediments of the basins, and can mix with the underlying groundwater (-5Og/L). A similar situation occurs below evaporation ponds for solar salt harvesting schemes. Both salt disposal and salt harvesting operations are likely to become more common in the Murray Basin, raising questions about their impact on regional groundwater quality. Natural groundwater discharge zones are also common in the Murray Basin. Here, groundwater is discharged by evaporation from shallow water tables in the semi-arid to arid climate. Brine bodies form due to the evapo-concentration of this groundwater in what have previously been considered to be closed systems. It is now recognised that natural salt lakes and engineered evaporation ponds can act as open systems for salt (e.g. Barnes et al., 1990), with

transport of the brines occurring by such proposed mechanisms as downward convection driven by the density inversion of a dense brine overlying a less dense groundwater system (Simmons & Narayan, 1997). Little field evidence for such transport processes exists but can be expected to be found in the hydrochemical signatures of the resulting brines. This study seeks to evaluate the use of environmental tracers to provide evidence for mixing of saline brines with regional groundwater. A range of natural salt lakes, forming part of a groundwater discharge complex, is used as a long time-scale analogue for processes that can be expected to occur in more recently comissioned evaporation ponds. 2 THE SITE Raak Plain, in south-eastern Australia, is a groundwater discharge complex containing approximately 50 small (natural) playa lakes in an area of approximately 400 km2 (Figure 1). The “lakes” rarely contain surface water, however water tables are generally within 10 cm of the ground surface. This allows evaporation of groundwater through the unsaturated zone and the formation of evapo-concentrated brines. The playa deposits, on which the salt lakes sit are directly underlain by the Blanchetown Clay, a 30 m thick sequence of tight clays, silts and minor fine sands (Macumber, 1992). This is underlain by the 537

Figure 2. Regional groundwater concentrations versus distance along Transect AA’.

Figure 1. Site map.

regional Parilla Sand aquifer, a 60-70 m thick sandstone layer. Regional groundwater flow is approximately from east to west. Three of the natural salt lakes, Western Salina, Main Salina and Salt Lake (Figure l), were chosen for the study based on variability in lake size, position along the regional groundwater flowpath, and the presence/absence of surface water and salt crusts. This was done to ensure that the influence of a range of variables was investigated but, because the lakes are in the same geological and climatic setting, these two variables can be eliminated.

3 GROUNDWATER AND BRINE CHEMISTRY 3.1 Regional Groundwater Groundwaters from bores installed in the regional Parilla Sands aquifer, between 40m and 70m depths, were sampled in November, 1999. The chemical composition of the groundwater (see Table 1) is similar to that of seawater, resulting from its origin in the accession of cyclic seasalt via rainfall (Macumber, 1992). Na and C1 are the dominant ions, along with high concentrations of Mg and SO4. Figure 2 shows a general increase in the groundwater solute concentrations as it flows beneath the groundwater discharge zone, from east to west along transect AA’ (see Figure 1). This addition of salts may be the result of mixing with the concentrated brines associated with the discharge zone. 3.2 Brines The ranges in ion concentrations and 6l80 and 62H enrichments of the brines beneath the salt lakes at Raak Plain are presented in Table 1, along with a typical chemical composition of the near-surface brine in the centre of each lake. The brines are up to

10 times the concentration of seawater and, like the regional groundwater, are dominated by Na and C1, with high Mg and SO4. The ratios of heavy to light isotopes of H and 0, which form the water molecule are a good tracer of the water molecule itself. Their signatures are not complicated by geochemical reactions. The brines at Raak Plain are highly enriched in the heavy isotopes, with 6l80 values ranging between -1.9 and 9.6 %O and 62Hbetween -6.2 and 32.1 %O (Table 1). 4 BRINE LEAKAGE INFERRED FROM AN EVAPORATION/ LEAKAGE MODEL A basin evaporation/leakage model (Gonfiantini, 1986), based on an I8O mass balance for a wellmixed, constant volume lake body (in this case, the brine body), was used to determine the degree of basin leakage (groundwater through-flow) from the three natural salinas at Raak Plain. The results, shown in Table 1, indicate that only a small proportion of the groundwater that flows into the basins leaves by evaporation, and hence there is a large amount of groundwater through-flow. This throughflow provides the necessary mechanism for the transport of salt away from the discharge zone. In comparing the percentage of inflow that leaves the basin as through-flow, it is found that Salt Lake (7078% through-flow) is a more closed basin than Western Salina (86-89%), but the Main Salina (6575%) is the most closed to leakage of all three. The relative closure of Salt Lake to brine leakage explains high brine salinities below 1 m depth ([Cl] = 150 000 mg/L) and hence the preservation of a salt crust. In comparison, brine salinities below 1 m at Western Salina are around 100 000 mg/L C1- with no surface salt crust. A study by Wood and Sanford (1990) explains this, finding that the amount of brine leaking from a system via groundwater outflow influences brine concentrations and the formation of evaporite deposits. A more closed system to groundwater outflow allows higher brine concentrations to build up, whereas large amounts of throughflow inhibit this.

538

Table 1. Major ion concentrations, stable isotope enrichments and estimated through-flow. Ranges Regional Groundwater Brines 25 100-50 100 76000-196000 33 -47 10 - 100 328 - 814 120 - 1 000 150 - 300 390 - 1 400 4600-11000 1 470 - 2 620 13 000-28 100 44000-114000 13002470 8000-35 000 34 - 143 150 - 630 0 - 1 000 5.92-6.59 4.4 - 7.5 -1.9 -9.6 -6.2 - 32.1 N/a N/a

The Main Salina appears as a contradiction to this as it is estimated by the model to be the most closed to leakage of all three salt lakes but salinities there are comparable to those at Western Salina, much lower than at Salt Lake. This may be due to the lower concentration of groundwater inflow compared to the Western Salina for example. Hence, the salinas at Raak Plain can be considered to be through-flow systems, with the amount of throughflow varying and possibly influencing the concentrations of the brines, although other factors are also involved in this.

5 SALINITY PROFILES To further investigate downward movement of the brines below the salt lakes at Raak Plain, the vertical extents of the brines below the Western Salina and Main Salina were investigated. This was done by examining the chloride concentrations versus depth profiles from deep cores taken in March, 1999. These profiles are shown in Figure 3 (a and b). Site West4 is located approximately 3m from the edge of the Western Salina. Chloride concentrations

Shallow Brines at the Centre of Each Lake Western Salina Main Salina Salt Lake 177 500 172 600 189 500 30 69 21 307 254 149 770 938 1285 11 700 14 400 24 900 95 000 103 400 87 900 25 500 34 700 48 700 430 306 620 416 0 90 6.89 4.33 6.56 9.47 4.85 5.82 32.12 23.88 86 - 89 65 - 75 70 - 78

above 8.5 m depth are up to 115 000 mg/L, below which there is a decrease in concentrations between 8.5 m and 14 m depth to between 25 000 mg/L and 55 000 mg/L. The sharpness of this decrease is unknown due to loss of sample in this region. Concentrations below this remain relatively constant to a depth of 33 m, where sampling stopped. The regional groundwater C1 concentration at bore 50074, located next to West 4, was 50 100 mg/L in November, 1999 and historical data (Department of Natural Resources and Environment, 1996) from the period 1982 to 1986 indicates fluctuations between 31 000 mg/L and 52 000 mg/L. This salinity profile indicates that a distinct boundary between the brine body and regional groundwater is located between the depths of 8.5 m and 14 m. At site Main3, at the edge of the Main Salina, the maximum chloride concentration at a depth of 0.5 m below the surface is 132 000 mg/L. This drops quickly to 42 000 mg/L and then gradually to approximately regional groundwater concentration (between 27 000 mg/L and 29 000 mg/L) at 5 m depth, which is the boundary between the brine and regional groundwater. This brine has a lower salinity than the brine at Western Salina, probably due to the lower salinity of the parent groundwater. The salinity profiles show that the concentration and vertical extent of the brine at the edge of the Main Salina is much less than at the Western Salina. These results are in agreement with the results of the evaporation/leakage model, which indicates that brine leakage is much less below the Main Salina than at the Western Salina. It is possible that the brine bodies penetrate to greater depths at the centres of the lakes, however difficulties with sampling at the centres of the lakes have prevented the collection of such data. 6 GEOCHEMICAL MODELLING

Figure 3. Chloride versus depth profiles.

To further test the hypothesis that brines from groundwater discharge zones interact with their par-

539

Figure 4. Results of PHRQPITZ mixing simulations. Regional groundwater is evaporated to halite saturation and mixed with (a) regional groundwater (b) shallow groundwater below recharge areas (c) precipitation (rainfall).

charge zone. As yet, the mechanism for this is unknown but large (variable) amounts of groundwater through-flow from the lakes provides the necessary transport mechanism and, in fact, appears to be one of the controls on the salinities and vertical extents of the brines themselves. Furthermore, re-mixing of the brines with regional groundwater (rather than a lower degree of evaporation) accounts for the observed brine compositions, indicating some interaction between the brine bodies and the regional groundwater. Further investigation into the mineralization processes occurring is required to confirm this. The brine bodies at the edges of the salinas can extend to depths of up to 15 m. This may be even greater near the centres of the salinas. Distinct boundaries between the brines and the underlying groundwater can be identified from chloride versus depth profiles.

ent regional groundwater, a simple geochemical model for this was tested. The geochemical model, PHRQPITZ (Plummer et al., 1988) was used to simulate: 1 evaporation of regional groundwater to either halite or gypsum saturation (to investigate the importance of the degree of evaporation of the brines) 2 re-mixing of the resulting solution with (a) regional groundwater, (b) shallow groundwater from surrounding recharge areas and (c) rainfall. The resulting solution compositions were compared with observed brine compositions from shallow (up to 3 m depth) piezometers in the lakes, using Na, K, Mg, Ca, C1, SO4 and Br as the ions of interest. An error was calculated between observed brine compositions and the model predictions by averaging the partial errors for each species. The partial errors (E) between observed concentration ([X]o,s) and the predicted concentration ([XlCalc)are calculated using the following formula:

REFERENCES Barnes, C.J., Chambers, L.A., Herczeg, A.L., Jacobson, G., Williams, B.G. & Wooding, R.A., 1990, Mixing processes between saline groundwater and evaporation brines in groundwater discharge zones., International Conference on Groundwater in Large Sedimentary Basins. Department of Natural Resources and Environment, 1996, Groundwater Database, h ttp ://ww w .dce.vic .gov. addnrelgrnd wtrlgrnd wtr .h tm. Howes, N.M., 1998, Geochemistry and Hydrologic Processes in the Evolution of Hypersaline Brines at a Groundwater Discharge Area, Raak Plain, Murray Basin, Australia., Honours Thesis, The Flinders University of S.A. Macumber, P.G., 1991, Interaction Between Ground Water and Surface Systems in Northern Victoria., PhD. Thesis, Victorian Department of Conservation and Environment, Australia. Macumber, P.G., 1992, Hydrological processes in the Tyrrell Basin, southeastern Australia., Chemical Geology, 96, pp. 1-8. Plummer, L.N., Parkhurst, D.L., Fleming, G.W., and Dunkle, S.A., 1988, A computer program incorporating Pitzer’s equations for calculation of geochemical reactions in brines: U.S. Geological Survey Water-Resources Investigations Report 88-4153,310 p. Simmons, C.T. & Narayan, K.A., 1997, Mixed convection processes below a saline disposal basin., J. of Hydrol., 194, pp 263-285. Wood, W.W. & Sanford, W.E., 1990, Groundwater control of evaporite deposition., Economic Geol., 85, pp 1226-1235.

[xlobs The results of the preliminary modelling exercise for brines evaporated to halite saturation (as these provided the best matches with observed data) are shown in Figure 4. The diagram shows that, despite high errors between observed and predicted brine compositions, higher proportions of mixing of the brines with fresher solutions yield brine compositions closer to those observed at Raak Plain. The large errors are caused by discrepancies in one or two chemical species, specifically identifying the mineralization processes not accounted for in the model. Further development of the model is planned to better account for these processes. This exercise provides further evidence that interaction of the brines with regional groundwater is an important process in determining their chemical evolution. 7 CONCLUSIONS Regional groundwater flowing below Raak Plain accumulates salt as a result of interacting with the dis540

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Origin of sodium-bicarbonate waters in the south-eastern part of the Great Artesian Basin: Influx of magmatic CO, J.Jankowski & W.McLean UNSW Groundwater Centre, School of Geology, University of New South Wales, Sydney, N.S. W., Australia

ABSTRACT: Sodium-bicarbonate groundwaters occur in the south-eastern margin of the Great Artesian Basin (GAB), New South Wales, Australia. These waters experience enrichment of 613C compared to waters from other parts of the basin, where carbon influx is from biogenic CO2 and dissolution of carbonates. Cainozoic volcanic activity in the eastern part of the GAB modifies the 613C signature of groundwaters due to the possible influx of heavy carbon-13 from magmatic activity. Chemical and isotopic composition of these waters indicate that GAB groundwaters are of meteoric origin with an external source of CO2 modifying 613C values.

Most of the hydrogeochemical and isotopic studies in the GAB were done in central and western parts of the basin. Airey et a1 (1979) investigated stable and isotopic data from the Queensland portion of the GAB. Airey et a1 (1983) and Calf & Habermehl (1984) first indicated that the anomalous enrichment of 6l3C (average 613C =: -7%0) occurs near the outcrop of the Pilliga Sandstone aquifer in eastern part of the basin. Studies by Calf (1978) in the same area in the alluvial aquifer system of the Lower Namoi River catchment found groundwaters from a depth interval of 75-100 m to have 613C values of -10 to 6%0. Groundwaters from depths >100 m have values -9.9 to -5.4%0. Torgersen et a1 (1987, 1992) used helium ratios to suggest that waters from the eastern part of the GAB contain elevated 3He. They concluded that this is associated with the influence of mantle derived gasses which is a result of Cainozoic volcanic activity. Also Collerson et a1 (1988) found that the unradiogenic juvenile 87Sr/86Srcompositions are attributed to interaction between basin waters and relatively young mafic intrusions. Studies by Baker at a1 (1995) indicate a magmatic source of carbon in sedimentary rocks in the Bowen-Gunnedah-Sydney Basin system in eastern Australia. Recent detailed hydrochemical and isotopic studies (Lavitt, 1998; Schofield 1998; Schofield & Jankowski 1998) in the area immediately to the east of the GAB eastern margin confirm that influx of CO2 of magmatic origin is Troducing Na-HC03 groundwaters with enriched 6l C signatures, in both alluvial and bedrock aquifers.

1 INTRODUCTION Several hydrogeochemical studies into isotopic composition and groundwater chemistry from the GAB have been published. These studies showed that groundwaters in the majority of the basin belong to the Na-HC03 chemical type. Habermehl (1980, 1983, 1986) concluded that Na-HC03 groundwaters originate from several processes beginning with generation of CO2 in soil zone which in turn produces carbonic acid. This acid is a driving force for the dissolution of CaC03 producing inorganic carbon in the form of HC03. Other processes include reduction of SO4 to H2S and ion exchange between Ca and Na. Herczeg et a1 (1991) reviewed chemistry of groundwaters in the central and western part of the basin and had a similar conclusion to those of Habermehl (1986). The latter authors suggested that soil generated CO2 dissolves carbonates and silicates, producing dissolved inorganic carbon @IC) which has a carbon isotopic signature of -12%0 ( 6 ' 3 C ~ ~Mass ~ ) . balance reactions show that sodium is introduced to the groundwater system by a threestage processes involving; carbonate dissolution to produce Ca, Mg and HC03, cation exchange of Na for Ca to supply Na, and reaction of Na-rich kaolinite to form a Na-smectite. This reaction consumes alkalinity and releases protons, which buffers pH, allowing continued carbonate dissolution and ion exchange. As these processes progress CO2 is reduced to produce C&, and as the aquifer system is open to C02, in situ anaerobic fermentation produces CO2 which in turn enriches the 613C~1c. 541

2 GEOLOGY AND HYDROGEOLOGY The studied area is located within the Surat Basin, a subbasin of the GAB, and separated from the larger Eromanga Basin by the Triassic Nebine Ridge. The south-eastern part of the Surat Basin is further subdivided into the Coonamble Embayment that is the focus area of this study. The main aquifer in the area is the Jurassic Pilliga Sandstone, with minor contributions from the Purlawaugh Formation and Orallo Formation. Other units are Mooga Sandstone and the marine Bungil and Wallumbilla Formations. The confining beds comprise parts of the Bungil Formation and the major confining unit, the Rolling Downs Group. The Pilliga Sandstone is a uniform unit consisting of medium to very coarse well-sorted quartzose sandstone. The sandstone is generally quite porous and permeable but often pore spaces have been infilled with kaolinite formed by the post depositional breakdown of feldspars and mica. Fresh siderite and calcite cement has been observed in thin sections of the Pilliga Sandstone and more porous zones in the sandstone can be classified as primary or secondary porosity produced by leaching of carbonate-cemented zones (Arditto 1983). The main recharge zone for the GAB occurs in the southeast of the basin where the Pilliga Sandstone outcrops. The aquifers are generally well separated from highly saline waters in the marine Cretaceous units. The Cainozoic magmatism in eastern Australia is indicated by a series of large shield volcanoes, basaltic lava fields and small isolated plugs and sills. Wellman & McDougall (1974) dated Miocene volcanic rocks in eastern Australia using K-Ar methods. Volcanic rocks from the Dubbo area were dated at 12-15 Ma (Dubbo Province) and at 13-17 Ma (Warrungable Volcano), from the Mooki area at 32-36 Ma (Liverpool Volcano) and from the Namoi River catchment at 17-21 Ma (Nandewar Volcano). The magmatic activity is related to the northerly movement of the Australian continent over one or possibly several hotspots.

Figure 1. Major ion composition.

Figure 2 is a Deffeyes style plot illustrating the relationship between total alkalinity and DIC (Deffeyes 1965). From this graph, the contribution to total alkalinity and DIC by CaC03 dissolution and various redox processes can be deduced. Groundwater of Na-Cl chemical type and mixed waters plot close to 1:l line. These waters experience reduction of SO4 to reduced S and as most of these waters have pH about 6, the reduction of SO4 will produce both H2CO3 and HCO3 in nearly equilibrium molar concentrations according to reaction:

Na-HC03 groundwaters show an increase in DIC without an associated increase in total alkalinity indicating that an external source of CO2 is contributing to the increase in DIC. The Mooki and Dubbo waters with elevated DIC due to ingassing of geogenic CO2 show this incremental shift along the DIC axis. The GAB waters (insert), however, plot along the 1:l line, indicating that DIC and total alkalinity are in equilibrium. Only one GAB sample showed an increase in DIC suggesting an influx of magmatic co2.

3 GROUNDWATER CHEMISTRY Na-HC03 water samples from two aquifers in Dubbo and Mooki areas representing fractured bedrock aquifer and deep alluvial aquifer respectively have been compared to GAB groundwaters from the southeast part of the basin. The dominant ionic species in all GAB waters are Na and HCO3 comprising >85% of these ions in solution (Fig. 1). The C1, Ca, and Mg concentrations are low; SO4 very often is not present in strongly negative redox conditions. These waters are mixed with Na-CI waters present in strata of marine origin or shallow alluvial aquifers to produce Na-HC03-Cl type groundwaters.

Figure 2. Total Alkalinity versus DIC.

542

Figure 3. Deuterium versus oxygen-18.

Figure 5. 6I3C versus l/DIC.

Fi ure 3 shows the relationship between 6D versus 6* 0. All waters plot close to the Global Meteoric Water Line (GMWL) indicating a meteoric origin. Na-C1 waters from the Mooki and Dubbo areas lie to the right of the GMWL. The enrichment in 6D and 6l80 is associated predominantly with groundwaters of higher salinity and can be attributed possibly to evaporative concentration or secondary evaporation prior to infiltration. Na-HC03 waters plot along or to the left of the GWML. The position of these waters on the diagram is associated with isotopic exchange. The more likely exchange might involve C02(gas) andor HZS(gas) entering the system from deep sources and progressively mixing with the groundwater. This shift would reflect isotopic exchange reactions resulting from the readjustment of isotopic equilibrium between the gas phase and the groundwater. Figure 4 shows the carbon evolution of groundwaters in the three areas. In all waters evolution commences with isotopically light carbon (613C =: -14 to -12%0). This signature is derived from the intermixing of primary biogenic carbon (6I3C = -23%0) and inorganic carbon from the dissolution of CaC03 (613C =: -6 to O%O). The evolution is also clearly visible on the plot 613C versus 1DIC (Fig. 5). In all

groundwaters the isotopic evolution from low to high HC03 waters is clearly different. Na-C1 and mixed waters with low HC03 concentrations have light 613C values ranging from -16 to -1l%0 for all areas. Na-HC03 groundwaters with low 1/DIC (that is high DIC) have enriched 613C values and plot as group with 613C values ranging from -4 to +3%0. Groundwaters from the Dubbo and Mooki regions receiving an influx of CO2 lie in this group. GAB groundwaters plot in a distinct group between the formerly described groups. The calculated average 613C value of GAB groundwaters is -8.9%0. A mass balance calculation for 613C taking mantle 613C as -5.9%0 (Rollinson 1993, Schofield 1998) and the average 613C value for groundwaters with mixed biogenichnorganic carbon as -12.7%0 produces a value of 613C for GAB waters of -9.3%0, which is consistent with the calculated average.

F

4 GEOCHEMICAL PROCESSES To test the origin of Na-HC03 waters and assess hydrogeochemical processes the computer program NETPATH (Plummer et al. 1991) has been used to calculate a mass-balance model, which describes this groundwater system. Chemical analyses of fresh recharge groundwaters and the flowing artesian GAB waters were used as end members. The massbalance modelling, which is supported by 613C data describes the chemical evolution and processes of the Na-HC03 groundwaters. In exchange reaction the species listed first exchanges to the clay and the second species is released into solution. The selection of phases is based on geological, mineralogical and hydrogeochemical assumptions, field observations and understanding of hydrogeology and hydrochemistry in the basin. The mass balance is as follows: Fresh Groundwater +1.41C02 (gas) + 5.47 CaC03 + 0.02 K-Feldspar 5.58 Ca/Na-Exchange + 0.66 CaMg(CO& + 0.13 Kaolinite + 0.002 FeC03 + GAB Water

*

Figure 4. 613C versus Total Alkalinity.

543

The dominant processes in the GAB aquifer are dissolution of 5.47 mmolkg CaC03, exchange of 5.58 mmolkg Ca for Na and influx of 1.42 mmol/kg C02(,,,,. The isotopic composition computed by NETPATH of the final water is -7.71 for 613C, which is consistent with the observed data of 7.63%0 for the representative sample of the GAB water.

5 DISCUSSION AND CONCLUSIONS Hydrogeochemical studies in the south-eastem margin of the GAB suggest that the origin of HC03rich groundwaters is associated with the influx of COzgasfrom magmatic activity. Addition of CO2 with 613C of -5.9%0 elevates DIC and enriches 613C. Ingassing of CO:! and subsequent dissociation to form carbonic acid, lowers the pH of groundwaters and enhances dissolution of carbonates. Subsequently carbonate dissolution reactions remove acidity, elevating pH to values of 7.6-7.9. This theory for the origin of Na-HC03 groundwaters in alluvial and bedrock aquifers of the GAB is supported by several geological and geochemical factors: (1) Miocene magmatic activity in this part of the GAB; (2) Groundwaters from the Dubbo and Mooki areas which are located on the boundary of GAB have heavy 613C with high concentration of DIC and are experiencing an influx of magmatic COz; (3) The principal Pilliga Sandstone aquifer in this area is a coarse quartzose sandstone with small amount of organic matter; (4) Hydrogeochemical modelling which is consistent with modelled and observed 613C. Chemically and isotopically similar groundwaters, however, can be produced by bacterial generation of methane by CO2 reduction. The redox reaction involving anaerobic fermentation producing CO2 can enrich 613C and increase values of DIC. Methane production has been found in the GAB central and western parts in much older and deeper groundwaters than these south-eastern GAB waters. To resolve the problem of the origin of carbon in south-eastem part of the GAB further studies are necessary including collection of gases and radiogenic isotopes. REFERENCES Airey, P.L., G.E. Calf, B.L. Campbell, P.E. Hartley, D. Roman & M.A. Habermehl 1979. Aspects of the isotope hydrology of the Great Artesian Basin, Australia. In Proc. Int. Symp. on Isotope Hydrology, IAEA, Neuherberg, Fed. Rep. Germany, 19-23 June 1978: 205-219. IAEA, Vienna, 1979. Airey, P.L., H. Bentley, G.E. Calf, S.N. Davis, D. Elmore, H. Grove, M.A. Habermehl, F. Phillips, J. Smith & T. Torgersen 1983. Isotope hydrology of the Great Artesian Basin, Australia. In Proc. Int. Con$ on Groundwater & Man, Syd-

544

ney, Australia, 5-9 December 1983: 1-11. Aust. Water Resour. Counc. Conf. Ser. No. 8. Aust. Gov. Pub. Service, Canberra, 1983. Arditto, P.A. 1983. Mineral-groundwater interactions and the formation of authigenic kaolinite within the southeastern intake beds of the Great Australian (Artesian) Basin, New South Wales, Australia. Sedimentary Geol. 35: 249-261. Baker, J.C., G.P. Bai, P.J. Hamilton, S.Z. Golding & J.B. Keene 1995. Continental-scale magmatic carbon dioxide seepage recorded by dawsonite in the Bowen-GunnedahSydney Basin system, Eastern Australia. J. Sed. Res. A65(3): 522-530. Calf, G.E. 1978. An investigation of recharge to the Namoi Valley aquifers using environmental isotopes. Ausf. J. Soil Res. 16: 197-207. Calf, G.E. & M.A. Habermehl 1984. Isotope hydrology and hydrochemistry of the Great Artesian Basin, Australia. In Proc. hit. Symp. Isotope Hydrology in Water Resources Development, IAEA, Vienna, Austria, 12-16 September 1983: 397-413. IAEA, Vienna, 1984. Collerson, K.D., W.J. Ullman & T. Torgersen 1988. Ground waters with unradiogenic 87Sr/86Srratios in the Great Artesian Basin, Australia. Geology 16:59-63 Deffeyes, K.S. 1965. Carbonate equilibria: A graphic and algebraic approach. Limnol. Oceanog. 10: 4 12-426. Habermehl, M.A. 1980. The Great Artesian Basin, Australia. BMR J. Aust. Geol. Geophys. 5( 1): 9-38. Habermehl, M.A. 1983. Hydrogeology and hydrochemistry of the Great Artesian Basin, Australia. In Proc. Int. Con$ on Groundwater & Man, Sydney, Australia, 5-9 December 1983: 83-98. Aust. Water. Resour. Counc. Conf. Ser. No 8. Aust. Gov. Pub. Service, Canberra, 1983. Habermehl, M.A. 1986. Regional groundwater movement, hydrochemistry and hydrocarbon migration in the Eromanga Basin. In D.I. Gravestock, P.S. Moore & G.M. Pitt (eds.), Contributions to the geology and hydrocarbon potential of the Erornanga Basin. Spec. Pub. No 12: 353-376. Geol. Soc. Aust. Inc. Herczeg, A.L., T. Torgersen, A.R. Chivas & M.A. Habermehl 1991. Geochemistry of groundwaters from the Great Artesian Basin, Australia. J. Hydrol. 126: 225-245. Lavitt, N. 1998. Integrated approach to geology, hydrogeology and hydrogeochemistry in the Lower Mooki River catchment. PhD Thesis, UNSW. Plummer, L.N., E.C. Prestemon & D.L. Parkhurst 1991. An interactive code (NETPATH)for modelling NET geochemical reactions along aflow PATH. USGS Water Res. Invest. Rep. 91-4078. Rollinson, H.R. 1993. Using Geochemical Data: Evaluation, Presentation, Interpretation. Singapore: Longman. Schofield, S. 1998. The geology, hydrogeology and hydrogeochemistry of the Ballimore region, central New South Wales. PhD Thesis, UNSW. Schofield, S. & J. Jankowski 1998. The origin of sodiumbicarbonate groundwaters in a fractured aquifer experiencing magmatic carbon dioxide degassing, the Ballimore region, central New South Wales, Australia. In G.B. Arehart & J.R. Hulston (eds), Proc. qhInt. Symp. Water-Rock Interaction-WRI-9, Taupo, New Zealand, 30 March-3 April 1998: 27 1-274. Rotterdam: Balkema. Torgersen, T., W.B. Clarke & M.A. Habermehl 1987. Helium isotopic evidence for recent subcrustal volcanism in eastern Australia. Geophys. Res. Lett. 14(12):1215-1218. Torgersen, T., M.A. Habermehl & W.B. Clarke 1992. Crustal helium fluxes and heat flow in the Great Artesian Basin, Australia. Chem. Geol. 102: 139-152. Wellman, P. & I. McDougall 1974. Potassium-argon ages on the Cainozoic volcanic rocks of New South Wales. .IGeol. . SOC.Aust. 21: 247-272.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Chemical evolution of groundwater in the Tularosa Basin in Southern New Mexico, USA Thomas G .Kretzschmar Departamento de Ingenieria Civil y Ambiental, Universidad Autdnoma de Cd. Jua'rez, Chihuahua, M&xico

Dirk Schulze-Makuch Departainent of Geological Sciences, University of Texas at El Paso, Texas, USA

Ignacio S .Tones-Alvarado Centro de Investigacioiz en Energia, UNAM, Temixco, Morelos, Mexico

ABSTRACT: The Tularosa basin of Southern New Mexico bears a wide range of water quality, ranging from low mineralized waters at the flanks of the adjacent mountain ranges to the highly saline water at Lake Lucero. The chemical and microbiological composition of the waters from the Tortugas Mountain Geothennal Area indicate that the sampled hydrothermal water derives from a mixture zone of deep water with meteoric water from the alluvial fans. Comparing the SI obtained through a simple cooling process of the hydrothermal waters with SI values from dilute waters, important similarities were found with respect to dolomite, calcite and magnesite in the waters HTA-3, W-5 and SW-1OA. Waters from the Tularosa basin indicate a constant supersaturation for dolomite and equilibrium with calcite. Both areas are lying in the stability field of dolomite, which might strengthen the microbiological evidence for a connection between these two systems. 1 INTRODUCTION The purpose of this study is to compare two hydrologic systems, the Tularosa basin and the Tortugas Mountain Geothermal Area, which might be connected despite their geographical separation (Fig. 1). 1.1 Geological Setting

rocks and Paleozoic sedimentary rocks form the San Andres Mountains, whereas the Paleozoic rocks are the same as exposed in the Sacramento Mountains. The Organ Mountains consist of Paleozoic, Cretaceous and Tertiary sedimentary rocks to the south and masses of tertiary intrusive rocks (quartz monzonite) to the north (Seager 1981). The Tortugas Mountain Geothennal Area (TMGA) is located on the western flank of the Organ Mountain Range. It is associated with the Rio GrandeRift

The Tularosa Basin is a faulted, north-south oriented, intermountain depression in southern New Mexico. The basin fill consists of an up to 1000 m thick sequence of alluvial and lacustrine deposits. Earlier investigations showed that most of the sediments are saturated by very saline water (Orr & Myers 1986). The alluvial fans on the east and west flanks of the basin however host low mineralized waters. Groundwater flow in the basin is directed from the flanks of the San Andres and Organ Mountains in the west and the Sacramento Mountains in the east towards Lake Lucero, which is located in the deepest part of the basin. The rocks exposed in the Sacramento Mountains range from Precambrian to Permian. The Precambrian rocks include slightly metamorphosed sandstone, siltstone, shale and some diabase sills. The Paleozoic sedimentary rocks are mainly limestone, dolomite, shale and sandstone (Pray 1961). Precambrian granites, Precambrian metamorphic

Figure 1. Location of the study area. I 3 Tularosa Basin Drinking Water, 0 Tularosa Basin high TDS, Tortugas Mountain Low TDS, Tortugas Mountain Geothermal water.

*

545

+

and is partly located on the Jornada del Muerto Fault Zone (Schulze-Makuch & Kennedy 2000). It has been proposed to be the result of high regional heat flow and deep circulation of groundwater in a bedrock-hosted regional groundwater flow system (Morgan et al. 1981, Barrol & Reiter 1990). Analysis of waters from the Lake Lucero area of the Tularosa basin also resulted in the identification of thermoacidophilic Archaea that were interpreted to derive from an upflow component of the Rio Grande Rift system (Schulze-Makuch, pers. comm.). 2 HYDROCHEMICAL DATA Groundwater samples were collected from the Tularosa Basin and the TMGA. Additionally, hydrochemical data from earlier investigations has been taken into consideration (Schulze-Makuch & Kennedy 2000, Cruz 1985).

Figure 2. Piper diagram of the studied waters. For an explanation of the symbols see Fig. 1.

2.1 Tularosa Basin The water sample GWBP-D represents the groundwater at Lake Lucero. It was taken at a depth of 1.5 m below the playa surface. The TDS content of about 180,000 mg/l is over five times above the mean TDS value for seawater. The chemical composition is dominated by sodium and chloride, with relatively high values of magnesium and sulfate (Table 1). The samples representing the “low” TDS type of waters of the Tularosa Basin are literature data (MAR, SW-lOA, HTA-3) of drinking water from the White Sands Missile Range facilities (Cruz 1985), or were collected within the basin (WHSA-MW2, WHSA-MW3, PZ-4). The composition of these water samples is also summarized in Table 1. The Table 1. Hydrochemical data of selected wells in the Tularosa Basin. Concentrations are in mg/l. Sample

GWBPO WHSA-

WHSA-

M

Temperature(C) PH Cond.(pS/cm)

17 7.17 132000

W 2 m 13 21 7.47 7.84 5150 14500

PZ4 23.6 8.1 NA

53200 370 14300 2240 1 NA

1780 590 510 3.9 0.11 0.73

400 506 21 1 2.94 0.05 20

3090 483 1810 72 0.7 NA

2000 69 370 27 NA NA

22 33 7.5 1.8 NA NA

64 93 22 0.8 0.004 NA

Chloride Sulfate Nitrate Bicarbonate

82000 24300

2870 3570 CO.05 80.2

244 2570 CO.05

2500 9490 CO.05 764

670 5700 3.8 120

13 47 cO.05 97

NA 140 3.8 235

% 70 MC). Deeper samples (> 40 m) have higher salinities (up to 5000 mg/L) are older (< 5 % MC), with an increased contribution to salinity from dissolution of carbonate minerals over time. Mixing between these two end members has been trigged by increase recharge to the system since the clearing of native vegetation 100 years ago. Enhanced recharge of low salinity “young” water has induced back diffusion of solutes from the matrix to the fractures resulting in a net export of salt via groundwater discharge to streams.

-

1 INTRODUCTION Anthropogenic impacts such as contamination spills and acid rain and their deleterious impact on groundwater resources are well established. In the Australian continent the removal of native vegetation to make way for European style agriculture has similarly resulted in an adverse impact on the water quality in streams and regional sedimentary groundwater systems (Allison et a1 1990). In this study we examine the impacts that clearing of native vegetation has had on fractured rock aquifers in a sub-humid catchment in the Clare Valley, South Australia. Our study site is intensively instrumented with a set of 10 vertically nested piezometers in a fractured carbonaceous dolomite aquifer. There are very few studies of this type where geochemical process are examined in vertical profiles in low permeability media (e.g., Hendry 2000). Assuming that groundwaters sampled at great depth correspond to increasing age, we can evaluate both natural and anthropogenic evolution over long time scales.

2 FIELD SITE AND METHODS The Clare Valley is located 100 km north of Adelaide, South Australia (33’50’s; 138’37’E) and contains Proterozoic rocks (600 -800 million years old). These are exposed as indurated and fractured

quartzites, shales and dolomites with low porosity that have been subjected to low grade regional metamorphism. Mean annual precipitation varies from 590 to 650 m d y r throughout the study area and is winter dominated from June to August. Average annual evaporation is about 1975 m d y r . Groundwater of variable quality (500 - 700.0 mg/L,) and low yield (0.5 to 20 Usec) is stored in these low porosity rocks. Today the dominant land use is vineyards and pasture with only a minor amount of native vegetation remaining. Large scale clearing of native vegetation occurred approximately 100 years ago as a result of the introduction of European style agriculture. In other parts of Australia vegetation clearing has resulted in an order of magnitude increase in groundwater recharge (Allison et a1 1990). Pre-clearing recharge to the Clare Valley fractured rock aquifer was low (< 5 m d y r ) . Since clearing, recharge has increased by an order of magnitude. A well was drilled in 1997 (200 mm in diameter; to a depth of 117.5 m) and completed as an uncased open hole. Vertical profiles of electrical conductivity (EC) were taken using a down hole logging probe. Samples for major ion data were collected in situ in the unpurged open borehole by means of a bailer. This well, and another well drilled to a depth of 55 meters, 2 meters to the north, was later converted to a series of nested piezometers. The nested site has ten different well completions intervals located between 2 and 100 m below the 557

water table. The piezometers are open to the aquifer by slotted casing at their base, with slotted intervals from 2 to 6 meters. Individual piezometers are gravel packed around the slots and are separated by cement and bentonite seals. Piezometers were sampled for environmental tracers I4C, 6I3C, 62H, 6'*0. Porosity of a rock core from at 8.2 m was measured to be 4.7 % using helium porosimetry. The vertical hydraulic gradient at the site is less than 5 x10-3. Two flow systems have been recognized at this field site (Love et a1 1999). An upper flow system (< 40 m) is characterized by high horizontal flow, high fracture density with apertures ranging from 200 600 pm. The deeper flow system (> 40 m) is characterized by, low horizontal flow, larger fracture spacing and smaller apertures, with minimal hydraulic connection between the two systems. 3. RESULTS AND DISCUSSION 3.1 Grouizdwater Mixing Electrical conductivity (EC) profiles have been used to infer the location of fracture flow into the open borehole (Fig. la). Sharp increases in EC of between 0.3 to 1.5 mS/cm occur over vertical distances of only 1-3 metres. These discontinuity's in EC occur at depths of 36, 52, 80 and 85 meters below the water table. We believe these represent

locations of major groundwater inflow to the bore via fractures. Deuterium versus depth ranges from -21.5 %O at 2 m below the water table to -31.1 %O at 100 m below the water table (Fig.lb). This increasingly negative signature with depth (i.e, a -10 %O shift in 62H over 100 m) corresponds to an EC increase of - 2.6 mS/cm. This is the opposite to what is normally observed throughout the world where more negative 62H values are often associated with lower EC concentrations due to colder climatic conditions at the time of recharge. 613C concentrations become progressively enriched relative to I2C, increasing from -14.1 %O at 2 meters below the water table to - 3 %O at 100 meters (Fig.lc). Such a large shift in 6I3C (1 1 %O in 100 m) indicates geochemical control on the groundwater composition. The 6I3C data suggests a discontinuity between piezometer 4 and 5 where 6I3C changes from -12.8 to -6.8 %O over a vertical distance of only 6m. This most likely represents a divide between the upper and lower flow systems as was inferred from the EC profile. The pH decreases down profile from 7 to 6.35. The major ion data plotted against chloride (Fig. 2 a-f) all display positive linear correlations, with the concentrations of all dissolved ions increasing with depth (correlation coefficients (r2) = 0.86 - 0.97). The linear correlation for all ions versus C1 may be the result of a number of possible processes; 1) variable evapotranspiration of a single input water

Figure 1. Electrical Conductivity, deuterium and 6 l 3 C D I C profiles Wendouree a) EC profile open bore; b) Deuterium profile from piezometers; c) 6 ' 3 c D 1 C profile from piezometers. The depths are shown as below standing water level.

558

between these two end members; 1) low salinity waters of a relatively modern origin (i.e 14C >70 %MC) in the top 40 m and; 2) higher salinity older groundwater (i.e. 14C < 5%MC) with an increased contribution from water/ rock interactions at depth. We suggest that the observed linear correlations in chemical and isotopic data between these two end members is a result of mixing between older more saline water in the matrix and fresher younger water in the fractures. With increasing depth down the profile there tends to be a greater contribution from the older (matrix) water component. We believe that this mixing is triggered by increased recharge to the system due to vegetation clearing - 100 years ago. This would result in the increased flux of lower salinity water to the groundwater system. Under this scenario, greater flushing has occurred in the upper flow system due to higher horizontal flow, larger apertures and closer fracture spacing. Low rates of flushing persist for the lower flow system which maintain their long term ( 103- 104kyr) chemical and isotopic signatures.

during recharge, 2) progressive addition of ions via waterhock interactions, 3) mixing of two different water bodies with different end member compositions. In this case the mixing process would be by diffusion, where older more saline immobile water in the matrix would mix by diffusion into the relatively younger fresher more mobile water in the fractures. The stable isotopes of the water molecule plot on or above the meteoric water line (Fig. 2g) with a positive linear correlation that suggests possible mixing between two end -members. The slight offset to the left of the MWL for deep groundwaters maybe a result of isotopic exchange of oxygen between silicate minerals in the aquifer and the groundwater. The deep groundwaters satisfy the most important criteria for exchange to occur; a low water-rock ratio and aotentially long residence times. The I3C versus C data (Fig. 2h) also shows a remarkable linear correlation with 14C concentrations at 90 % MC at 2 meters below the water table to background ( 13 mmol/L (ie, > 36 m) both Na and SO4 are above the ion/CI marine ratio suggesting that addition of these ions to the groundwater above the seawater dilution line is by w aterhock interactions. The major source of dissolved salts to the upper flow system is from the concentration of rainfall cyclic salts due to evapotranpiration. The shallow near surface end member has modern 14C with a 6I3C composition consistent with equilibrium with the soil CO2 reservoir. For the deeper flow system there is a greater

559

contribution to the water chemistry from waterhock interactions. The deep end member has background 14C and relatively high 6I3C of -3 %o, which is consistent with equilibrium with the dolomite. To reconcile the origin of the more saline end member we propose two possible scenarios; the first involves the dissolution of dolomite via reaction with high CO2 and dissolution with gypsum. The second involves via reactions with protons generated through reduced sulfur mineral oxidization (SMO). increase in We note that both Ca2+ and SO': equimolar proportions (Fig. 3a) suggesting that dissolution of gypsum may be an important process in scenario 1. This would involve dissolution of gypsum (CaS04.2H20) to produce Ca. Increasing Ca concentrations in solution results in CaC03 to precipitate which in turn results in a decrease in pH and an increase in CO2 according to reaction 1:

Implicit in reaction (3) is the need for a source of oxygen to generate Fe3+. The dissolved oxygen maybe transported via advective flow through the fracture network. The generation H+ will cause dolomite to be undersaturated and result in increases of Mg and Ca in solution. As with scenario 1, we would expect some calcite to re-precipitate when dolomite saturation is approached. CONCLUSIONS Increased recharge as a result of the removal of native vegetation has triggered mixing between low salinity water and saline matrix water in a fractured dolomite aquifer. This has resulted in back diffusion of salt from the matrix into the fractures. Two distinct groundwater end members are recognized. The chemistry of the shallow ( Na+K, and Ca > Mg and Na > K. The anion concentrations also show a clear trend. In low TDS waters HCO; is the dominant anion. High TDS waters are generally rich in SO:and Cl-. On the basis of the relative proportion of the major ions 3 types of waters can be distinguished:

(2) Ca-SOq-HC03-water These waters show high Ca concentrations and very low Na concentrations. SO4 is the dominant anion but HC03 is important too. The C1 concentrations are very low. These waters were collected out of fractures and exploration drillholes. TDS is still quite low but significantly higher than in group (1).

(3) Ca-S04-C1-water These waters were collected in the deepest parts of the mine out of exploration drillholes. TDS is significantly higher (log-scale) than in all other samples. Na concentrations increased too and are wth 20-40 meq% significantly higher than in any other group before. The main anions are so4 and C1 with together about 60 meq%. HC03 remained nearly unchanged compared to group no. 2.

(1) Ca-HC03 water Precipitation, like rain and snow water, and most surface waters and some other low TDS-waters from within the mine out of exploration drillholes belong to this group. HC03 is with more than 70 meq% the dominant anion. There is not much difference between precipitation-, surface- or waters out of drillholes. 586

me@

Mg

Ca

Ca+Mg

NatK

CI

HC03

SQ

Mg

Ca

CatMg

NatK

CI

HC03

SQ

Mg

Ca

Ca+Mg

Na+K

CI

HQ

l

i

l

l

l

SO,

100

10

waers from explorationdrllhdes 01

I

I

I

I

I

I

I

0 01

O 0 I

0 001

Figure 2. Schoeller-Diagramm of the waters in the Clara mine, Black Forest, Germany.

From type-one to type-three the electrical conductivity increases from about 100 p / c m to more than 1000 ps/cm.

2NaAlSi308+ 2C02 + 11H20 = A12Si205(OH)4 (albite) kaolinite) + 4H4Si04 + 2Na' + HC03-

4 THERMODYNAMIC CALCULATIONS

2CaAlSi308 + 4 CO2 + 8 H20 = Al~Si20j(OH)4 (anorthite) kaolinite) + 4H4Si04 + 2Ca2' + 4HCO3'

We calculated the component activities and the saturation state of the sampled groundwater using the code PHREEQE (Parkurst et al. 1990). Generally, the waters of the Clara mine are weakly oversaturated with respect to quartz ( 0 . N SI < 0.5). The saturation with respect to carbonates depends on the TDS. Calcite saturation rapidly increases with increasing TDS and reaches equilibrium when TDS exceeds 200 mg/l. With respect to dolomite a similar correlation can be observed, however, few waters reach dolomite saturation. All waters are undersaturated with respect to gypsum. The saturation indexes increase parabolically with increasing TDS. During the water-rock interaction process along the flow path of the infiltrating meteoric water, secondary minerals, predominantly kaolinite and illite, form from the primary minerals of the gneiss, predominantly plagioclase and biotite. The feldspar processes can be described by the reactions: 2KAiSi3O8+ 2C02 + 11H20 = A12Si20j(OH)4 (K-feldspar) (kaolinite) + 4H4Si04 + 2K' + HC03-

Figure 3 shows the stability-diagrams for feldspars and their weathering products (Bowers et al. 1984, Berman 1988). The dashed lines represent quartz and amorphous silica saturation, respectively. All waters are weakly oversaturated with respect to quartz. Silica concentration increases with increasing TDS. The diagram shows that all surfaceand groundwaters are consistent with kaolinite as the prime Al-rich residue of the feldspar and mica decomposition. Mine waters with high pH values, however, fall near the boundary to K-feldspar. 5 CONCLUSION Water from the Clara mine is local meteoric water. Its chemical evolution is controlled by water-rock reactions. Feldspar and biotite weathering and dissolution of soluble minerals, e.g. calcite, gypsum, sellaite, etc. are important processes. Plagioclase weathering forms abundant kaolinite and creates a Ca-HC03-type water. This water descends along fractures and the barite-fluorite veins. Calcium 587

Figure 3. Stability diagrams of feldspars and their weathering product minerals (15"C, 1 bar).

and sulphate increases ratidly as a result of dissolutibn of secondary calcite and gypsum. The water changes to a Ca-S04-HC03- type. With increasing depth, the water grades finally into a CaS04-Cl-water with a dramatical increase in Na too. Chloride is mostly contributed by opened fluid inclusions in primary minerals. Stober & Bucher (1999a, 1999b) showed that the salinity of deep crystalline basement water in the Black Forest is mainly of marine origin. I

Stober, I. 1995. Die Wasserfiihrung des kristallinen Gmndgebirge. Verlag, StUttiW. 191PP. Stober, I. & K. Bucher 1999a. Deep groundwater in the crystalline basement of the Black Forest region. Appl. Geochem. 14: 237-254. Stober, I. & K. Bucher 1999b. Origin of salinity of deep groundwater in crystalline rocks. Terra Nova 11 (4): 181185.

I

ACKNOWLEDGEMENTS We thank Dr.K.-H.Huck from the Clara mine for support in the field and S. Hirth- Walther and E. Lutz for assistance with the analytical work.

REFERENCES Berman, R.G. 1988. Internally-Consistent Thermodynamic Data for Minerals in the System: NazO- KzO- CaO- MgOFeO- FezO3- A1203- SOz- TiOz- HzO- COz. Journal of Petrology 29: 445-522. Bowers, T.S., Jackson, K.J. & H.C. Helgeson 1984. Equilibrium Activity Diagrams for Coexisting Minerals and Aqueous Solutions at Pressures and Temperature to 5kb and 600°C. Springer-Verlag Berlin Heidelberg. Huck, K.H. 1986. Clara am Schwarzbruch. In: Bliedtner, M. & Martin, M.: Erz- und Minerallagerststten des Mittleren Schwarzwaldes. Geologisches Landesamt BadenWurttemberg, Freiburg iBr., 366-399. Parkhurst, D.L., Thorstenson, D.C. & L.N. Plummer 1980. PHREEQE ' - a computer program for geochemical calculations. Water Resources Investigations, 80-96, 2 10 pp., U S . Geological Survey.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Hydrologic controls on groundwater salinisation, Murray Basin, Australia I.P.Swane, T.R.Weaver & C .R.Lawrence Hydrogeology & Environmental Research Group, School of Earth Sciences, University of Melbourne, Victoria, Australia

I .Cartwright Hydrogeology & Environmental Research Group, Department of Earth Sciences, Monash University, Victoria, Australia

ABSTRACT: Groundwater flows through a multiple aquifer system near the regional recharge area of the southern Murray Basin. Despite low topographic relief (-70m elevation / 15,000 km2) local recharge and discharge processes control groundwater quality. Total dissolved solids (TDS) contents in the upper aquifer (-30m deep) range from 650 to 105,000 mg/L over short distances (27 km). In the deepest aquifer at -180m, TDS ranges from 600 to 17,000 mg/L over the same area. In this semi-arid environment, recharge of fresh water occurs along palaeo-sand ridges and infiltrates the upper aquifer (-650 mg/L). However, in a nearby palaeo-river channel, groundwater in the shallow aquifer is highly saline (>100,000 mg/L) due to evaporative concentration in salt lakes and playas. This discharge area compromises water quality throughout the entire aquifer system, and indicates that brines from salt lakes and playas may penetrate by reflux to depths of over 180m, potentially transporting salt from the surface to aquifers relatively rapidly and over short distances. 1 INTRODUCTION This paper focuses on a hydrostratigraphically complex section of the Wimmera Region, southern Murray Basin. This region is characterised by a multiple aquifedaquitard system and is near the regional recharge area for the southern Murray Basin (Fis. 1). The Wimmera region extends over 15,000 km , and comprises three major aquifers: the Parilla Sand, the Murray Group Limestone and the basal Renmark Group (Fig. 2). In this region the Parilla Sand and the Renmark Group aquifers extend throughout most of the Wimmera, however, the Murray Group aquifer is present only west of the Wimmera River. This boundary represents part of a major facies change from marginal marine-lagoonal to shallow-marine and includes the Douglas Depression, a narrow, shallow palaeo-channel that extends from the Dundas Plateau in the south to beyond Lake Hindmarsh in the north (Fig. 1). Groundwater use in the Wimmera is primarily from aquifers west of the Douglas Depression, which include the Murray Group Limestone aquifer and some sections of the Parilla Sand and Renmark Group aquifers. Groundwater use east of the Douglas Depression in the area shown in Figure 1 is minimal, with most water supplied from reservoirs via an overland channel system. aquifers is from the area recharge to adjacent to the basin margin (Fig. 1) on the northern flanks of the Dundas Plateau (part of the Southern

Highlands) and Grampians. Direct infiltration of rainfall across the Wimmera Region is likely to provide some recharge to the unconfined Parilla Sand. Locally, the Parilla Sand also has the potential to be

Figure 1. Location of the study area within the southern Murray Basin; showing the Douglas Depression and sampling locations.

589

2 FIELD AND ANALYTICAL METHODS Nested monitoring wells located within and around the Douglas Depression and the palaeo-sand ridges of the central Wimmera were sampled for this research (Fig. 1): Gerang Gerung, Duchembegarra, Quantong, Natimuk, Jilpanger and Jungkum. With the exception of Gerang Gerung (which has been sampled only once), groundwater elevations and field parameters were measured, and groundwater samples were collected during three sampling rounds during 1999 - summer, winter and spring. Field parameter measurements included pH, temperature, redox (Eh), electrical conductivity (EC), dissolved oxygen (DO) and alkalinity. Cation, anion and stable isotope chemistry was determined by laboratory analysis. Groundwater was extracted from the wells using a 50mm diameter QED Well-Wizard submersible pneumatic bladder pump, utilising low-flow sampling methods. Groundwater pumped to the surface passed through a flow-cell for continual monitoring of pH, temperature, Eh and EC, and measurements were recorded once these parameters had stabilised. Samples were collected for immediate analysis of DO and alkalinity by titration. Samples for cation analysis were filtered (0.45 pm), preserved with ultra-distilled HNO3 to pHNa+>Mg2+Cl->SO>->K+; while in arkosic aquifers Mg2+ is lower than S042-. Some deep wells (S14 and Sl5) show a special composition, being C1->HC03--Ca2+ >Na+-Mg2+>SO>- >K+. Both aquifers are characterised by the highest TDS content and silica concentration as well as a good correlation between Na:Ca, Na:CI and Ca:Cl(r>0.7) showing a change in slope when C1T 110 mg/l (Fig. 3). Waters from arkosic aquifers have thus similar chemical characteristics as those from granites, even though they show a lower mineralisation. High SO-: contents can be related to Quaternary paleosols or human pollution input (supported by NO3-occurren-

1.52 0.7 0.3 0.3 0.9 0.7 0.5 1.5 0.4 0.7

'

L qJ F .1

30.4 21.3 12.2 11.3 16.4 11.4 13.3 12.5 13.2 10.3 12.0 11.6 12.9 10.6 10.4 4.7 6.8 10.0 10.0 5.7 5.6 6.8 5.2 7.9 7.3 9.4 7.2 5.3 7.2 6.9 6.5 6.1 8.5

C1-

SO4"

NO3-

mg/l 40.6 78.7 76.1 37.5 36.8 60.8 65.2 71.6 40.4 I09 47.9 27.6 94.6 245.0 247.0 22.0 61.5 41.2 77.5 127 26.6 28.9 62.5 40.9 33.1 39.9 87.5 94.3 33.0 44.4 78.4 51.5 117 89.8 59.8 55.6 36.9 60.0

mg/l 6.6 3.0 19.2 18.6 63.2 65.0 67.4 51.2 15.3 44.3 31.7 31.4 58.1 74.9 37.4 52.1 7.5 39.83 71.9 31.4 44.9 79.6 76.7 70.7 6.2 20.4 137.7 117.9 49.1 77.3 146.1 52.7 40.7 173.4 82.1 105.1 16.8 41.7

mg/l 0 37.0 0 0 0 50.0 71.5 22.5 0 0 0 15.7 22.5 0 0 17.0 0 34.0 0 0 0 9.5 40.5 8.1 -

42.2 64.2 45.0 29.0 11.0 0 0 0 18.3 36.9 -

HC03- S.I. cal mg/l 1503.1 0.03 3673.5 -0.06 1912.4 -0.06 1417.6 -0.17 888.0 -0.1 299.1 -0.1 1 200.9 -0.45 273.9 -0.08 383.5 0.32 438.6 0.22 417.6 0.32 214.2 -0.11 259.5 -0.23 376.0 0.12 303.6 0.07 225.7 -0.07 348.2 -0.26 323.7 0.13 278.4 -0.07 304.1 0.07 394.9 0.48 381.0 0.41 431.6 0.32 369.0 0.17 219.0 0.32 285.1 -0.13 426.0 0.51 401.3 0.46 277.8 0.14 278.3 0.29 347.6 0.39 349.6 0.59 382.0 0.73 352.2 0.92 329.9 0.16 431.6 0.5 200.9 -0.28

ce). Before considering water-rock interaction, we need to remark that some individual concentrations in chloride as well as other element concentrations, can vary over relatively short distances, illustrating the importance of emplacement in fractures rather than depth. Granitic and arkosic hydrochemistry is then attributed to the irreversible dissolution of primary silicate minerals and formation of secondary silicate minerals. HC03- originated by meteoric CO2 input, is formed during silicate dissolution. In such aquifers the molar HC03-:Ca2+ratio is, in general, greater than 2 suggesting the evidence of silicate hydrolysis (Fig. 4). Also, a general slight undersaturation in calcite (calculated with computer program PHREE C-2; Parkhurst & Appelo 1999) proves the slow C j + release from silicate minerals. Nevertheless the presence in calcite in some veins is related to its oversaturation by Ca-plagioclase dissolution. These reactions are relevant in deepest wells with highest

602

Figure 4. Plot of HC03-versus Ca" molar content. Legend as in Fig. 2.

Figure 2. Piper diagram for the different water samples.

chloride and cationic content, indicating a much longer residence time.These waters show relatively moderate HC03- concentration. It suggest that some equilibrium is reached, probably related to a limited supply of CO2 or to calcite precipitation (Fig. 5). The easiest weathering minerals of this granite are albite and biotite.The greater molar Naf:Ca2+ratio is attributed to acid hydrolysis of albite and oligoclase, and biotite alteration could account for the build up of C1- and Mg2+(Edmunds et a1 1984) (Fig. 3). Due

Figure 3. Ca, Mg and Na plotted against C1. Legend is the same as in Figure 2. Solid lines show correlation for granitic and arkosic water samples only.

to the slow rate of biotite dissolution, high Mg2+and C1- concentrations are attributed to long residence times, and the depletion of Na' relative to C1- also supports this idea and excludes the possibility of trapped water. Also, this relatively low Na' value, constitutes evidence of partial equilibrium with albite, since albite hydrolysis is possible during its evolution to equilibrium (Paces 1973). Furthermore, some depletion of Mg2+can occur as a result of observed biotite alteration to chlorite and by its uptake by secondary products that results from incongruent plagioclase dissolution. On stability-field diagrams representing partial geochemical systems, samples plot within the smectite stability field although kaolinite has also been noticed in the field (Roquk 1993) (Fig. 6). Smectite is a usual weathering product in arid climates with low flushing rate and relatively longer residence times (Tardy 1971) although is difficult to assess if true equilibrium is attained. The similar chemistry between arkosic and granitic water samples indicates an hydraulic connection between bedrock and detrital overlaying material, and therefore a potential vertical recharge. Nevertheless, the presence of C1-rich waters indicates that such recharge is locally limited; otherwise, mixing would reduce this higher C1- concentration. The par-

Figure 5. Possible approach of equilibrium trend in groundwater charged with COz. Legend as in Fig. 2.

603

plained by mixing with recent meteoric waters. Molar Na+:Cl-ranges from 1.2 to l l , and are poorly correlated, suggesting enhanced water-rock interaction by low pH and by extended residence time. These waters are oversaturated or near saturation in chalcedony, If chalcedony control is accepted, this geothermometer gives a temperature between 40 to 85 "C. 4 CONCLUSION

Figure 6. Stability diagram for albite, smectite, kaolinite and gibbsite at 25°C and 1 atm (Tardy 1971). Legend as in Fig. 2.

ticular hydrochemistry of this water is useful as an indicator of flow-paths within this crystalline massif. Waters from metamorphic Palaeozoic rocks that outcrop at the north part, show a variable cationic content which is difficult to correlate because of lithological variations over relatively short distances. These Ca-HC03 waters have relatively low TDS values. This lower mineralisation is a response of the slow kinetic alteration of metamorphic rock minerals (Rogers 1989). They are also undersaturated in calcite and in chalcedony; nevertheless some high silica contents can be related to quartzitic levels. Waters from sandstones and limestone Paleogene aquifers are characterised by high Ca2' and Sod2-, and low silica content, and low Mg2+:Ca2' molar ratio. Ca:HC03 molar ratio is in agreement with stoichiometric carbonate dissolution. All waters are saturated or slightly oversaturated in calcite, and most of them also in dolomite. The chemistry of these waters is dominated by reversible carbonate dissolution. These reactions are established very quickly, and they can mask previous aluminosilicate dissolution happened at the recharge area. They show however, a great range of variation in some cases, correlated with surface pollution input (NO3-, K'). Sandstones aquifers are poor in Sr2+and Mg2+, and the samples nearest to Palaeozoic massif are enriched in some elements (Mg2+,Sr2+,Li', Pb2+). The COz-rich waters are characterised by a high mineralisation (TDS between 2300-5500 mg/l) and pH values below 6.5. HC03- is the predominant anion (900-3700 mg/l), and the main cation is calcium or sodium. Four samples are undersaturated in calcite, and only one with a higher calcium content has reached saturation. Molar HC03-:Ca2' ratios (Figure 4) vary from 2.8 to 8 showing that HC03- is in excess with respect to the carbonate system pointing out a limited income of Ca2+from silicate hydrolysis. High pC02 enhances the ability to react and buffers the water at relatively low pH (Fig. 5). Water trend from Na to Ca-bicarbonate type could be ex604

Hydrochemical analysis provides a fair distinction of groundwater pathways within the Gavarres massif. Even though most of the sampled waters are of CaHC03 type, specific relationships allow characterizing of their origin. Within the massif, waters from granites are distinguishable from those of metamorphic terrain. In particular, silicate hydrolysis enhanced by meteoric COz, is responsible for their chemistry. In granites, results show the evidence of preferential flow through fractures together with a heavy mineralisation along flow-paths as residence time increases. Water chemistry of the arkosic aquifer in the southwestern basin is in agreement with that of granites. In the north part, carbonate dissolution in Paleogene aquifers masks the chemical signature of infiltrated water in metamorphic recharge area. Finally, CO2 rich waters indicate the existence of deep flow-paths. ACKNOWLEDGMENTS The authors are grateful to W.M. Edmunds and P. Shand, from B.G.S, for their assessment and fruitful discussions. This study was supported by grant CICYT-HID98-0366 (Research agency of the Spanish Government).

REFERENCES Edmunds, W.M., J.N. Andrews, W.G. Burgess, R.L.F. Kay & D.J. Lee 1984. The evolution of saline and thermal groundwaters in the Carnmenellis granite. Mirzeralogical Magazine. 48: 407-424. Paces, T. 1973. Steady-state Kinetics and equilibrium between ground water and granitic rock. Geoclzinz. Cosnzochim. Acta. 37: 2641-2663. Parkhurst, D.L & C.A.J. Appelo 1999. PHREEQC-2. A Computer program for speciation, batch-reaction, onedimensional transport, and inverse geochemical calculations: U.S. Geol. Surv. Water Research Invest Report 994259,312 p p . Rogers, R.J. 1989. Geochemical comparison of ground waters in areas of New England, New York, and Pennsylvania. Ground Water. 7(5):690-7 11. RoquC, C. 1993. Litomorfologia dels massissos de les Gavarres i de Begur. Tesi Doctoral, UAB. 494 pp. Inkdita. Tardy, Y. 197 1. Characterization of the principal weathering types by the geochemistry of waters from some European and African crystalline massifs. Chenz. Geol.. 7: 253-27 1.

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Hydrogeochernistry of shallow groundwaters from the northern part of the Datong basin, China Runfu Wang Institute of Geo-environniental Monitoring of Shanxi Province, Taiyuan, Shanxi; China University of Geosciences, Wuhan, P.R. China

Yanxin Wang & Huarning Guo China University of Geosciences, 430074 Wuhan, P.R. China

ABSTRACT: Located in the arid-semi arid region of northwestern China, the Datong Basin is one of the Cenozoic rift basins of the Shanxi Rift System. Groundwater is the most important source for water supply at Datong. Three hydrogeochemical zones were delineated: basalt-water interaction zone, municipal and industrial wastewater-affected zone and irrigation wastewater-affected zone. Monitoring results showed that, over the past 20 years, the levels of shallow Quaternary groundwaters at urbanized areas declined continuously and groundwater quality at the urbanized and agricultural areas deteriorated. Over-exploitation of groundwater has not only caused shortage of water resource but affected groundwater quality. Correlation analysis indicates that there is a close relationship between contaminant (nitrate, chloride, sulfate, and fluoride) concentration and groundwater drawdown. 1 INTRODUCTION

middle aquifer which is composed of sand and sandy gravel is 70- 160 m, with hydraulic conductivities of 0.31-9.10 d d . The buried depth of groundwaters in. the lower aquifer consisting of fine sand and silt is more than 200 m, with the permeability coefficient of about 0.06-4.69 d d . Changes in the groundwater levels of the three aquifers respond well with each other, indicating that they have close hydraulic interconnections. The general direction of groundwater flow in the northern region is from north to south, while that of the rest of the regions are from west to east (Fig. 1)

Studying hydrogeochemical evolution of regional groundwaters is important for scientific management and sustainable development of water resources. Groundwater monitoring is often the basis for the study. The hydrochemistry of aquifers can also be well defined by properly analyzing groundwater monitoring data (Makhnach et al. 2000, Grobe et al. 2000). Located in the arid-semi arid region of northwestern China, the Datong Basin is one of the Cenozoic rift basins of the Shanxi Rift System (Wang & Shpeyzer 2000). Groundwater is the most important source for water supply at Datong. Groundwater monitoring work started in 1981. A total of 71 water table monitoring wells and 65 water quality monitoring wells have been in use since then in the northern part of the Datong Basin. The mean annual evapotranspiration is around 1980 m d a at Datong, far exceeding the mean annual precipitation of 377 mm. The thickness of Cenozoic sediments ranges from 50 m and 2500 m at Datong. Shallow groundwaters occur in the Quaternary alluvial, alluvial-diluvial and alluvial-lacustrine aquifers. The Quaternary groundwater systems can be basically divided into three groups: upper, middle and lower aquifer. The buried depth of groundwaters in the upper aquifer that consists of coarse sand and gravel is 5-60 m, with hydraulic conductivities of 1.4314.89 d d . The buried depth of groundwaters in the

2 GROUNDWATER MONITORING RESULTS 2.1 Groundwater level Since 1981, groundwater levels of the monitoring wells near the Datong city have been declining. One typical case is the monitoring well K23 where the buried depth changed from about 20 m in 1981 to about 80 m in 1999. Two depression cones, one to the southwest of the Datong city and the other to the northeast, have been formed and tend to join with each other. The lowering of groundwater levels is clearly related to increased withdrawal of groundwater around the city. 2.2 Groundwater quality Affected by hydrodynamic conditions and hydrogeochemical processes, regional groundwater 605

general tendency of increase with time. The concentration of nitrate of water from borehole No.172 in 1994 was 30 times higher than that in 1981. 3 DISCUSSION

3.1 Delineation of hydrochemical zones

Figure 1. Hydrogeology of the northern part of Datong Basin. A, B, C respectively represents the basalt-water interaction zone, the municipal and industrial wastewater-affected zone and the irrigation affected zone.

chemistry changes in different parts of the groundwater flow system. From recharge area to discharge area, the concentrations of various components except silica in shallow groundwaters increase (Fig.2) and the hydrochemical type changes from bicarbonate water to chloride-sulfate water, as demonstrated by the hydrochemical section of the Shili River pluvial fan (Fig.3). The concentrations of water constitutions are relatively low in the upper part of the fan, as a result of deep groundwater table, small evaporation and high hydraulic conductivity. Decrease in groundwater buried depth and aquifer permeability and increase in evaporation down to the lower part of the fan result in high concentrations of dissolved components of groundwater. It is noticeable that the content of silica is higher in the recharge area than that in the discharge area. Human activity has considerable impact on groundwater quality. Different types of industrial, municipal and agricultural wastes have been released into groundwater systems at Datong. The contents of water components have been increasing since 1981. The changes in concentrations of sulfate, nitrate, total salinity, chloride, silica and bicarbonate with time were plotted in Figure 4. It can be seen that, except silica, the other components show a

Monitoring data indicate that the groundwater at Datong has been contaminated, and the contaminants were different at different locations. Nitrate is the major pollutant in the southwestern part, while sulfate, nitrate and chloride are the chief contaminants near Datong city. Almost no contaminants have been detected in groundwaters from the northern part of the study area. Considering the distribution and extent of contamination source, hydrogeological setting, meteorological condition and anthropogenic factors, we divide the study area into three hydrogeochemical zones (Fig.1): A. basalt-water interaction zone, B. municipal and industrial wastewater-affected zone, and C. irrigation wastewater-affected zone. A. Basalt-water interaction zone: This zone includes the highlands in the eastern, western, and northern parts of the region. Aquifer is composed of sandy gravel and medium-coarse sand, with groundwater buried depths of 20-100 m. Lateral flow of fracture witter in the basalt and metamorphic rocks that are vulnerable to pollution (Eaton & Alexander 1997, Paczynski 1997), is the primary recharge source for the groundwaters in the Quaternary aquifer. The quality of the groundwater, mainly HC03-CaMg water, is good with total dissolved solids of less than 300 mg/L. There are very few contamination sources in this zone. Groundwater is used chiefly for daily life of rural villages. B. Municipal and industrial wastewater-affected zone: This zone is situated in the urban area with high population density. The aquifer is chiefly composed of multiple thin layers of Middle Pleistocene alluvial fine sands that contain huge amount of water. Municipal and industrial wastewaters are not properly treated before discharge, which makes it easy for contaminants to reach groundwater through vertical infiltration of meteoric and surface water. Unfortunately, the intrinsic remediation capacity of the aquifer is limited. Besides, groundwater has been over-exploited. Long-term over-exploitation results in formation of three depression cones: in the downtown area, around Zhijiabu, and around No.428 factory respectively. Groundwaters are seriously contaminated in this zone. The concentration of nitrate is sometimes more than 40 mg/L and total dissolved solids are usually higher than 500 mg/L. 606

Concentration (mgh) Only fo ilica

Concentration (mgh)

Concentration (mg/L)

15.4

2000.0 7

I

1600.0

1200.0

800.0

400.0

2ooo 1500

1-I

P

500

15

0

0.0

78

63 101 Boring No.

27 -.

Figure 3. The groundwater components from the upper part to the lower part of the pluvial fan of the Shili River in 1994 (cross section B-B' in Fig.1). The legend of symbols is the same as in Figure 2.

Figure 2. Variations of groundwater chemistry from recharge zone to the center of the Datong basin in 1994 (cross section AA' in Fig.1). Legend of the symbols: *TDS; +bicarbonate; x silica; @chloride; onitrate; Asulfate

groundwater flow system. The aquifers are in the outskirts of pluvial fans and therefore made up of alluvial-lacustrine and alluvial-diluvial fine sand and silt, Vertical infiltration of meteoric water, surface

C. Irrigation wastewater-affected zone: This zone is the agricultural region area of the study area and is also the discharge area of the regional

750.0 -

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600.0

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1980 1984 1988 1992 1996 2000 60.0

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1980 1984 1988 1992 1996 2000

I

1

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R

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-1

0.0 1980 1984 1988 1992 1996 2000

1980 1984 1988 1992 1996 2000

Figure 4. Changes of bicarbonate, chloride, nitrate, silica, sulfate, and total salinity concentrations(in mg/L) with time. Symbol for boreho1es:O borehole No. 172, A borehole No.63, 0 borehole No.28.

607

water and irrigation water is the major way of groundwater recharge. Buried depths are 1 5 - 3 0 m. The wastewater irrigation was the main source of groundwater contamination. The major contaminants are nitrate and sulfate. Groundwater quality is the worst in the study area, with the total dissolved solid of over 800 mg/L, the sulfate concentration of above 250 mg/L, and nitrate content of over 40 mg/L.

groundwater and causes abundant contaminants to be introduced into groundwater system. That’s why a drop in the groundwater table due to overexploitation aggravates deterioration of groundwater quality.

3.2 The relationship between hydrochemistry and hydrodynamics

The research work has been financially supported by the National Natural Science Foundation of China (Grant No.49832005) and the Ministry of Science and Technology (Grant No.95-pre-39).

Wastes of human activities are important contamination sources for groundwaters. Overexploitation of groundwater is an indirect factor accelerating water quality deterioration. As stated above, groundwater exploitation has been increasing in recent years, resulting in a continuous decline of groundwater levels of shallow Quaternary aquifers at urbanized areas. In the meantime, groundwater quality at the urbanized and agricultural areas worsens. Therefore, studying the relationship between hydrochemistry and hydrodynamics of groundwaters is important for understanding the hydrogeochemistry of regional groundwaters. Correlation analysis of monitoring data from a well in the municipal and industrial wastewateraffected zone for 1 2 years (1 988- 1999) indicates that there is a close relationship between contaminants (nitrate, chloride, sulfate, and fluoride) concentration and groundwater drawdown. The correlation coefficients are 0.80, 0.73, 0.65, 0.37 respectively for nitrate, chloride, sulfate, and fluoride and drawdown. In addition, we also made the multiple linearity regression analysis for these monitoring data used for the correlation analysis. The multiple linearity regression analysis equation is shown as following:

ACKNOWLEDGEMENTS

REFERENCES Eaton, T.T. & Z. Alexander 1997. Evaluation of groundwater vulnerability in an urbanizing area. In Chilton et a1 (eds), Groundwater in the Urban Environment: Problem, Processes and Management: 577-582. Rotterdam: Balkema. Grobe, M., Machel, H. G. & H. Heuser 2000. Origin and evolution of saline groundwater in the Munsterland Cretaceous Basin, Germany: oxygen, and strontium isotope evidence. Journal of Geochemical Exploration 69-70: 5-9. Makhnach, N., Zernitskaya, V., Kolosov, I. et al. 2000. 6 I8O and 613C in calcite of freshwater carbonate deposits as indicators of climatic and hydrological changes in the Late-Glacial and Holocene in Belarus. Journal of Geochemical Exploration 69-70: 435-439. Paczynski, B. 1997. Valuation of groundwater: Future trend in vulnerability mapping. In Chilton et a1 (eds), Groundwater in the Urban Environ-ment: Problem, Processes and Management: 64 1-645. Rotterdam: Balkema. Wang, Y.X. & G.M. Shpeyzer 2000. Hydrogeochemistry of Mineral Waters from Rift System on the East Asia Continent: Case Studies in Shanxi and Baikal. Beijing, China Environmental Science Press.(in Chinese with English abstract)

D=1.578Cni~ra~e-.487Cma~esium+O.66Ocfr,,,,d,t0.408csr/ica+2.495 csu~ate+4.041Cca/cium+ 0 7 146cbicarbona1e-0 342cchloride *

Where D is groundwater drawdown (m);

Cnitrate,

Cmagnesrum Cfluorrde, Csrhca, csulfntet Ccalcium, Cbicarbonate, and Cchloride are the concentration of nitrate,

magnesium, fluoride, silica, sulfate, calcium, bicarbonate and chloride respectively. The result clearly shows the close relationship between drawdown and chemistry. The reasons are considered as follows. The increase of drawdown resulted in the increase of thickness of the aeration zone, which changes the hydrogeochemical condition and promotes oxidation of minerals in the matrix and ion exchange. On the other hand, the hydraulic gradient between the groundwater level and contaminated surface water table increases due to the decline of the groundwater level, which enhances the flow rate from the surface water to the 608

Decoupling solute distributions from groundwater flow in low permeability media T.R .Weaver Hydrogeology and Environment Research Group, University of Melbourne, VIC 3010, Australia

S .K .Frape & J .A.Cherry University of Waterloo, Waterloo, Ontario N2L 3G1, Canada

ABSTRACT: Solute transport modelling of a 40 m thick aquitard in Ontario, Canada, clearly demonstrates the decoupling that can occur between hydraulic conditions and groundwater chemistry in low permeability media. In this glacially-deposited aquitard (K-10-” d s ) , 62H and 6”O values indicate that the groundwater was recharged during the last glaciation (-10 ka). The distribution of 0 and H isotopes in the aquitard is controlled primarily by diffusion whereas the current head distribution has developed as a result of exploitation of underlying water and petroleum resources over the last century. The distribution of isotopes and, therefore, solutes in this low permeability system is largely unrelated to the current head distribution: instead it reflects hydraulic gradients present over geological time scales. As long as aquitards are not penetrated by fractures or improperly abandoned boreholes, the relatively rapid hydrodynamic response of aquitards can be ignored. However, if these preferential pathways for groundwater do exist, the hydrodynamic responses of aquitards to changing hydraulic conditions must be considered. region is underlain by 15-40 m of low-permeability clay-rich glacial till that was deposited during the Wisconsinan glaciation (eg Chapman & Putnam 1984). The sites discussed here are located where the glacial till comprises 2 units and is over 30 m thick. The lower till (occurring below depths of 14.5 m),

1 INTRODUCTION Throughout the world, low permeability formations (aquitards) are relied upon to protect water-supply aquifers from contamination and to separate aquifers of different water quality. Identifying the effectiveness of aquitards as isolating units becomes particularly important 1) when the aquitard is sufficiently transmissive to allow advective transport to occur across it over a period of years to decades; and 2) when fractures or leaky boreholes penetrate the aquitard, effectively by-passing the aquitard (e.g., Lacombe et al. 1995). Solute and stable isotope distributions have been used to evaluate the hydraulic conductivity of aquitards (eg Desaulniers et al. 1981) and to support their use as “protecting” units; however, unless hydrodynamic responses of aquitards through time are also determined, the potential for solutes and contaminants to migrate across aquitards that are penetrated by fractures, leaky boreholes, or other flaws cannot be predicted accurately. Here we demonstrate the clear decoupling between isotope and solute distribution and hydraulic gradients that can occur in low permeability formations.

1.1 Geology and Hydrogeology An is provided from Lambton County in southwestern Ontario, Canada (Fig. 1) where the

Figure 1. Location of the study area, Dow, and Laidlaw sampling sites.

609

deposited about 13 ka (Bamett 1993) is overlain by the St. Joseph till, the upper zone of which is now weathered and fractured (eg Ruland et al. 1991). Only the unweathered part of the till (below about 8 m depth) is discussed here, and is referred to here as the clay-till aquitard. Beneath the glacial deposits is a basal-till aquifer of clayey sands and gravel which is 1-3 m thick and that separates the aquitard from underlying bedrock. The interface aquifer is underlain by a series of Devonian black shales (Kettle Point Formation) and carbonate-rich shales (Hamilton Group). In central Lambton County, carbonates of the oil-bearing Dundee Formation underlie the Hamilton Group at a depth of about 140 m. Desaulniers (1986) and Ruland et al. (1991) determined that downward hydraulic gradients existed in the till at several sites in Lambton County. The distribution of stable isotopes clearly indicates that the unweathered, unfractured, clay-till aquitard is dominated by diffusion (Desaulniers et al. 1981, 1986; Weaver 1994; Husain 1996). Groundwater flow in the underlying interface aquifer is predominantly to the west and southwest (eg Vandenberg et al. 1977; Beaton 1994). This aquifer supplies water to rural Lambton County.

'-U'Y

I I

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conditions

m E

Y

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rquifer

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E

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200 300 400 500 600 Days since July 1991 sampling event

Iapproximate stea_dy-stateconditions

samolina

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aquifer bedrock Elapsed days since sampling event

Figure 2. Hydrographs of clay aquitard, interface aquifer, and bedrock at Laidlaw (upper) and Dow (lower) over 22 months.

7x10-" to 4x10-'* m f s (0.002 to 0.01 mfyr), further indicating that solute transport in this region is dominated by diffusion. Fluctuations in water levels in the interface aquifer at each site were much greater than in the overlying clay-till aquitard (Fig. 2). At Laidlaw, water levels in the interface aquifer fluctuated by 0.2-1.1 m over several days to several weeks and at Dow, they fluctuated by 0.1-0.4 m over the same time frame. Water levels in the lower clay-till aquitard at both sites varied by ~ 0 . 0 1m, within measurement error for standpipe piezometers. Hydraulic diffusivity (Dh) values were calculated based on the following relationship

2 METHODS AND RESULTS Water-level elevations in the lower clay-till aquitard, the interface aquifer, and the upper bedrock were measured over 22 months at two sites, Laidlaw (LD) and Dow (DW). Hydraulic gradients, hydraulicconductivity values, and the degree of hydraulic connection between the clay-till aquitard and the interface aquifer were determined from water-level responses at both sites. Specific-storage values were estimated from consolidation tests on four samples of clay till taken from depths of 19.5-29 m at the Laidlaw site (McKay 1991). The sites (Fig. 1) are located away from current petroleum-production areas (Weaver et al. 1995).

2.I Current potentiometric conditions After near steady-state conditions were reached, groundwater elevations in the lower clay-till aquitard at both Dow and Laidlaw were consistently lower than those in the underlying interface aquifer (Fig. 2). Under these conditions, the groundwater flow direction would be upward toward the clay-till aquitard. Hydraulic gradients (Vh) in the lower-most clay till ranged from -0.05 to -0.3 mfm (negative is upwards), maximum hydraulic conductivity (K) is 5x10-" mfs, and porosity is assumed to be 0.38 (Desaulniers 1986). Under these conditions, groundwater flow rates through the matrix of the lower clay-till aquitard would be negligible, about 610

(e.g., Bear 1972) where K, is vertical hydraulic conductivity ( d s ) and S, is specific storage (m-'). Using the range of hydraulic conductivit data derived from field tests (3.0x10-" to 5.8~1070 d s : Desaulniers 1986; Weaver 1994), a specific storage value of l ~ l O -m-*, ~ hydraulic diffusivity values range from 3 . 0 ~ 1 0 -to~ 5 . 8 ~ 1 0 m2/s - ~ for the lower clay-till aquitard at the Laidlaw site. These values are low enough to prevent the large fluctuations in head measured in the aquifer from propagating into the overlying clay-till aquitard during the monitoring period. The propagation of changes in hydraulic head in the interface aquifer to the lower clay-till aquitard was modeled using a 1-D analytical solution

(SFLOW2; Neville pers. comm.), to the following groundwater flow equation, assuming a semi-infinite aquitard

s

a2h ah --Aa z 2 - K, at

(2)

(e.g., Freeze and Cherry 1979). Even when the maximum Dh value of 6 . 7 ~ 1 0m2/s - ~ is assumed for the aquitard, groundwater elevations at the base of this unit would have increased only by a maximum of 0.02 m after 1 month in response to a change in hydraulic head in the aquifer of 1 m.

I

to 1970s aquifer bedrock

2.2 Hydraulic disequilibrium at the interface aquqer

Figure 3. Changing hydraulic gradients across the aquitard and aquifer sequence.

At the Laidlaw site, hydraulic gradients in the aquitard were upward to a depth of less than 22 m, and were downward above this depth. This pattern is consistent with downward hydraulic gradients identified in the upper 10-30 m of clay till elsewhere in the study area (Desaulniers et al. 1981; Desaulniers 1986; Ruland et al. 1991). These data indicate that, in the lower clay-till aquitard just above the interface aquifer, hydraulic gradients reverse from downward to upward. If this reversal represented equilibrium conditions, a zone capable of conducting lateral groundwater flow would need to be present in the lower aquitard over the central Lambton County area. No coarser-grained zones were identified in core collected from the base of the clay-rich till at either site (Weaver 1994). In the absence of these units, the current reversal in vertical hydraulic gradients is considered to represent hydraulic disequilibrium in the lower clay-till aqui tard.

Groundwater in the interface aquifer was also developed as a resource during this period, leading to long-term depressurisation of the aquifer (Fig. 32). In the last 30 years, however, groundwater use has been superseded by surface-water pipelines (Husain 1996) and groundwater extraction from the interface aquifer has declined. This has allowed the recovery of groundwater elevations in the interface aquifer. The reversal in hydraulic gradients identified in the lower clay-till aquitard during this research probably results from the pressure increase in the interface aquifer propagating in the overlying aquitard (Fig. 3-3). As head in the interface aquifer declined due to groundwater extraction and petroleum production, the change in head would have propagated through the overlying clay-till aquitard. Using the aquitard properties discussed above and SFLOW2, hydraulic gradients across the entire clay-till aquitard could be reversed on the order of 50-100 years. The downward hydraulic gradients currently present in the upper clay-till aquitard could have developed rapidly, within 1 or 2 decades of initial development of the interface aquifer, and may have persisted as the aquifer continued to be developed. Therefore, the downward gradients measured in the clay-till aquitard are most likely a recent phenomenon and do not represent long-term conditions in the aquitard.

3 CHANGING HYDRAULIC GRADIENTS Since current hydraulic gradients in the aquitard indicate non-equilibrium conditions, it is important to understand the changes in head distribution through time in the region. The development of petroleum and groundwater resources from depths of up to 150 m in the study area began in the mid-19th century. Petroleum seeps and saline water at the surface and in the interface aquifer indicate that hydraulic gradients were upward before development (Fig. 3-1). Oil wells drilled before 1880 into bedrock immediately below the interface aquifer were gushers with heads of up to 30-50ft (-10-15 m) above ground surface (Fairbank 1953), also indicating upward hydraulic gradients (Weaver et al. 1995). As petroleum resources were exploited, these gradients declined. Currently, petroleum is pumped from the oil-bearing formations.

4 DISCUSSION AND IMPLICATIONS Although the distribution of solutes in the aquitard indicates a diffusion controlled environment, groundwater and petroleum extraction in underlying units has produced profound, and rapid, changes in hydraulic gradients across the till. The implications of this are critical in situations where lowpermeability media are relied upon to protect and isolate hydrogeological units since aquitards may be 61 1

fractured or may be penetrated by improperly abandoned wells. Groundwater chemistry in aquitards can be used to indicate if a system is diffusion dominated. However, the occurrence of diffusion profiles for solutes or stable isotopes alone cannot be used to indicate that the aquitard does not respond to changing hydrogeological conditions. Therefore, it is important to model the response of the aquitard to changing groundwater elevations in surrounding aquifers or aquitards. Potentiometric conditions, and, therefore, hydraulic gradients, in thick low-permeability, moderately-compressible formations can change over time frames of only 1 to 2 decades in response to changing potentiometric conditions in minor, subjacent aquifers. By constrast, the solute distribution in these thick aquifers reflect long-term, pre-development hydraulic conditions in the aquitard. Current hydraulic conditions across aquitards will control solute movement where preferential pathways for fluid movement exist. By contrast, the distribution of solutes in the matrix will be controlled by the hydrogeological conditions that dominated the groundwater flow system for the hundreds or thousands of years during which the solute distribution in the matrix developed.

Lacombe, S., Sudicky, E.A., Frape, S.K. & A.J.A. Unger 1995. Influence of leaky boreholes on cross-formational groundwater flow and contaminant transport. Water Resources Research 31 ( 8 ) : 1871-1882. McKay, L.D. 1991, Groundwater flow and contaminant transport in a fractured clay till: Unpub. Ph.D. Thesis, University of Waterloo, Ontario, Canada, 4 10 p. Ruland, W.W., Cherry, J.A. & Stan Feenstra 1991, The depth of fractures and active ground-water flow in a clayey till plain in southwestern Ontario: Ground Water, 29 (3): 405-417. Vandenberg, A., Lawson, D. W., Charron, J.E. & B. Novakovic 1977. Subsurface waste disposal in Lambton County, Ontario. Fisheries and Environment Canada Technical Bulletin 90, 64 p. Weaver, T.R. 1994. Groundwater flow and solute transport in shallow Devonian bedrock formations and overlying Pleistocene units, Lambton County, southwestern Ontario: Unpub. Ph.D. Thesis, University of Waterloo, Ontario, Canada, 401 p. Weaver, T.R., Frape, S.K. & J.A. Cherry 1995, Recent crossformation fluid flow and mixing in the Michigan Basin, southwestern Ontario. Geological Society of America, Bulletin 107: 697-707.

REFERENCES Barnett, P.J. 1993. Quaternary geology of Ontario. Geology of Ontario, Ontario Geological Survey, Special Volume 4, p. 101 1-1088. Bear, J. 1972. Dynamics of fluids in porous media. New York, Dover Publications, 764 p. Beaton, W.A. 1994. An isotopic and geochemical survey of the Fresh Water Aquifer. Unpub. M.Sc. Project, Department of Earth Sciences, University of Waterloo, Ontario, Canada, 88 p. Chapman, L.P. & D.R. Putnam 1984. Physiography of southern Ontario (3d ed.). Ontario Geological Survey, Special Volume 2,270 p. Desaulniers, D.E. 1986. Groundwater origin, geochemistry and solute transport in three major glacial clay plains of eastcentral North America. Unpub. Ph.D. Thesis, University of Waterloo, Waterloo, Ontario, 445 p. Desaulniers, D.E., Cherry, J.A. & P. Fritz 1981. Origin, age and movement of pore water in argillaceous quaternary deposits at four sites in southwestern Ontario: Journal of Hydrology, v. 50, p. 23 1-257. Desaulniers, D.E., Kaufinann, R.S., Cherry, J.A. & H.W. 5~~ in a diffusion-controlled Bentley 1986. 3 7 ~ 1 - 3variations groundwater system. Geochimica et Cosmochimica Acta 50: 1757-1764. Fairbank, C.O. I1 1953. The Petrolia story: Northwest Oil Journal, Edmonton, Alberta, v. 2, p. 5-13. Freeze, R.A. & J.A. Cherry 1979. Groundwater: Englewood Cliffs, New Jersey, Prentice-Hall Inc., 604 p. Husain, M.M. 1996. Origin and persistence of Pleistocene and Holocene water in a regional clayey aquitard and underlying aquifer in part of southwestern Ontario. Unpub. Ph.D. Thesis, University of Waterloo, Ontario, Canada, 336 p.

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Sedimentary basins

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Water-Rock hteraction 2001, Cidu (ed.), 02001 Swets R Zeitlinger, Lisse, ISBN 90 2651 824 2

Fluid-sediment interaction and clay authigenesis along the flank of the Juan de Fuca Ridge M.B .Buatier, M.Steinmann & C.Bertrand EA2642, Ge'oosciences Universite'de Franche Comte',France

A .M Karpoff CGS-EOST Strasbourg, France

G .L.Fruh-Green ETH Zurich, Switzerland

ABSTRACT: This study focuses on the mineralogy and chemistry of sediments that overly basaltic basement through which seawater circulates (Leg168, flank of the Juan de Fuca Ridge). Silicate authigenesis was observed in the sediment layer just above basement at sites located more than 30 km from the ridge axis. This sediment alteration is particularly abundant at ODP Sites 1031 where the authigenic formation of Fe-Mg rich smectite (montmorillonite and saponite) and zeolite, and the dissolution of biogenic calcite, are observed. Vertical advection of basement fluid through the sediment section is required to produce this alteration. These processes are still active at Site 1031, based on systematic variations in pore-water profiles. A SEM and TEM-EDX study of the authigenic clays and zeolites enable the chemistry and mineralogy of the hydrothermal phases to be characterized and their mechanism of formation to be determined. This mineralogical study has been complemented by REE analyses on clay fractions in order to further constrain the origin of the authigenic clays and their environment of formation. l INTRODUCTION

ish-gray and light olive-gray clayey silt. The mineralogical composition of the basal sediments is marked by a higher phyllosilicate content related to an increase in the relative abundance of smectite (Fig. 2,

During Leg ODP 168, ten sites were drilled in a 100 km east-west transect located along the eastern flank of the Juan de Fuca Ridge (Fig. 1). Geophysical and chemical investigations suggested that fluid circulation is mostly restricted to the upper part of the basaltic basement and that this circulation is still active far from the ridge axis (Davis et al. 1997). Vertical advection of fluids through sediment is focused on a topographic high (Sites 1031 and 1030) where the sedimentary cover does not exceed 40 meters, compared to 200 to 600 meters-thick at other sites. Basal sediments at Site 1031 are less than 1 Ma old suggesting that sediment deposition started when the basement was more than 12 km from the ridge axis. Sediments located just above basement display thin beds of dusky yellowish-green silt mixed with green clays alternating with green-

Figure 2 : Mineralogical distribution with depth at Site 1031 (mbsf : meters below sea floor) (SM = smectite, MI = mica, CHL = chlorite, FS = Felspar, QTZ = quartz

Figure 1. Location of the drilling transect along the eastern flank of the Juan de Fuca Ridge

615

Davis et al. 1997). The relative abundance of feldspar and quartz is almost constant with depth (Davis et al. 1997). Analysis of the clay fraction confirmed that the sediments located just above the basement are smectite-rich compared with other sediments from the same lithology along the sedimentary COlomn (Buatier et al. 2001). The increase in the relative abundance of smectite in the clay fraction can be correlated with an increase in the total Mg and Fe content in the basal sediments. XRD analyses on clay fractions show smectite interlayers are composed of a mono-layer of water, i.e. the d(001)-spacing of smectite is about 12-138, under air-dried natural conditions and shifts to about 178, after glycolation.

2 RESULTS 2.1 SEM observations Secondary electron scanning images of hemipelagic sediments enabled the mineralogical assemblages of the sediments to be characterized. Most sediments are composed of detrital quartz, feldspar and clays. The detrital quartz and feldspars display rounded edges and anhedral morphologies. Calcareous microfossils are well preserved and in some levels they are the major component of the sediment. In contrast, sediments lying just above basement have clay minerals as the major component. Smectite with a honeycomb-like morphology can also be observed. In the same sample, coccoliths are either partially dissolved or pseudomorphosed by authigenic Fe-Mg rich clay minerals (Fig.3). Zeolite crystals are locally observed, associated with authigenic clays. Their morphology is characterized by euhedral to sub-euhedral prismatic crystals up to 10 or 15 pm.

2.2 Transmission Electron Microscopy (TEM) observations and analytical (AEM)data Smectite is the dominant clay in the basal sediments and occurs as veil-like particles of about 0.25 to 1 pm width. Small particles with fibrous morphologies are also abundant in the same sample. Most fibers are 0.5 pm long and only about 50 nannometer wide. They form aggregates and are associated with veillike smectite from which they appear to have grown. The structural formulae of the smectite particles are calculated on the basis of 22 negative charges. AI, Fe and Mg are present in octahedral sites for all the analyses although in varying relative abundance. The cations present in the interlayer are dominated by K with a very small proportion of Ca and Na. AEM analyses are reported in Figure 4 with AI, Fe and Mg rich end-members (Fig. 4) together with the bulk sediment composition. Three groups of smectite can be distinguished (open circle, triangle and square) in this diagram based on their chemistry, octahedral content and layer charge: (1) AI-rich montmorillonite, (2) Fe- and Mg-rich montmorillonite and (3) saponite. Smectite chemistry can explain the Fe and Mg enrichment of the basal sediment, i.e., the non-altered sediments and the AI-rich smectite (supposed to be detrital in origin) have similar Al/(Fe+Mg+AI) ratios, whereas the basal sediments are enriched in Fe and Mg and plot in the same part of the diagram as the authigenic Fe-Mg rich smectites. AEM analyses on the 10-2Opm sized-fraction revealed the presence of phillipsite, and Ba-rich zeolite is also detected. 2.3 Mass balance calculations In order to quantify the element mobility related to smectite and zeolite authigenesis in the basal sediments, we compare the bulk sediment chemistry of non-altered and altered sediments. The bulk sediAV(AI+Fe+Mg)

/ Fe/(AI+Fe+Mg)

Fe-Mg saponite saponitesaponite

\

Mg/(AI+Fe+Mg)

Figure 4. AEM analyses of individual smectite particles and bulk sediments from Site 1031.

Figure 3. SEM image of basal sediment showing partially dissolved coccoliths and authigenic clays.

616

ment data used for these calculations were averages from 4 samples of non-altered and 4 samples of altered sediments. The mobility of each element was calculated assuming that A1 was conservative. Figure 5 shows the gains and losses of various elements for the altered sediments. It can be seen that the alteration of sediment implies a loss of Ca and a gain of Mg and Fe. The same behavior has been observed by Alt and Teagle (1999) for seawater altered basalt. However, the roughly equal gain of Mg and loss of Ca observed for basalt alteration is not observed here. The gain of Mg is about 3 times the loss of Ca. The quantification of the Ca is difficult because the initial Ca content of the sediment (essentially in biogenic carbonate) is very variable in the hemipelagic sediment (Buatier et al. 2000).

I-

+1031-40,

1 :

1031-40.

2

8

.

w 0.1: a .

Residues site 1031 NMORB-normalized

1

I ( J 1 ( 1 ( 1 1 1 1 ( 1 1 1 1 1

0.01

L a C e P r Nd

Sm EU Gd Tb D y HO E r Rn Yb L u

Figure 6. REE pattern from analyses of the fine clay fraction after leaching (normalized to NMORB).

2.4 REE analyses Analyses of rare earth elements (REE) included < 2 pm fractions of authigenic clays that were leached with 1 M HC1 for 15 minutes at room temperature in order to remove secondary rare earths adsorbed on the clay surfaces. In the following, the cleaned clay fractions will be called “residues”, and the HC1 1 M solutions “leachates”. The REE distribution patterns of the residues normalized to NMORB are shown in Fig. 6. The patterns are flat for the heavy rare earths (HREE, DyLu) and enriched in the light rare earths (LREE, LaSm). ZREE is furthermore positively correlated with A1203 (r2 = 0.678, n = 9) suggesting that the rare earths are structurally integrated into the clays. The REE distribution patterns of the leachates when compared to the residues are depleted in Ce and enriched in the middle rare earths (MREE, Sm-Tb). ZREE of the leachates is positively correlated with P205 (r2 = 0.929) suggesting that the rare earths are not fixed in the silicate phase but in phosphates in contrast to the residues. Ce can occur as Ce3+or Ce4+.The latter is less soluble in HCI 1 M than the other REE which only occur in the trivalent oxidation stage under surface conditions. The Ce-depletion observed for the leachates therefore points to the presence of Ce as Ce 4+ and suggests that the authigenic clays have been formed under oxidizing conditions (> 100 mV at pH 8).

Si

AI

Mg

Ca

Fe

element

Mn

Ti

Na

K

Figure 5 : Chemical changes of altered sediment related to non altered sediments (data from Buatier et al, 2001)

3 DISCUSSION 3.1 Evidence of sediment alteration Mass balance calculations show that the bulk sediments next to the basalts have been enriched in Mg and Fe, and depleted in Ca. Fe-enrichment in oceanic sediments can be related to various processes. For example, hydrothermal metal accumulation related to plume fall out from active hydrothermal vents has been described in late Quaternary sediments close to the Valu Fa Ridge in the Lau Basin (Daessle et al. 2000). Fe and Mg enrichment in basal sediments can also be related to the alteration of the young basaltic crust by seawater. This process called ‘palagonitisation’ gives rise to the formation of Mg-smectite (saponite) and phillipsite replacing basaltic glasses (Honnorez 1981). In this study, the Fe-Mg enrichment of the basal sediments are related to the presence of Fe-Mg richsmectites. The Fe-Mg rich sediments contain partially dissolved calcareous fossils and authigenic clay minerals in-filling the porosity and replacing microfossils. AEM analyses confirm that the authigenic clays are smectites that are chemically distinct from the detrital smectites. The absence of basaltic glass and the presence of calcareous microfossils suggest that the basal sediments are not metalliferous deposits related to the alteration of the basaltic crust or precipitated below the Carbonate Compensation Depth (CCD). The authigenic minerals described here are clearly related to the interaction between a fluid and sediment. The basalt-like REE pattern of the leached clays suggests that the fluids from which the clay precipitated previously interacted with basalt.

3.2 Relation between alteration and upjlow The intensity of alteration in sediments from the flank of the Juan de Fuca Ridge does not depend on 617

the distance from ridge axis and there is no relation between the degree of sediment alteration or the age of the basaltic basement andor the age of sediment (Buatier et al. 2001). Significant alteration (with Fe and Mg enrichment in the basal sediments) has been observed at Sites 1031 and 1029, on 1.5 Ma and 2 Ma crust respectively. These observations suggest that alteration cannot be explained by diffusional transport of the solutes from the basalt though the sediment, but that advection of fluid is necessary to explained the highly altered sediment observed at this site. The Mg and Ca concentration profile in the pore water reflects vertical advection of crustal fluid at Site 1031. (Mottl and Wheat 1994; Davis et al. 1997). These fluids are relatively young (80% illite, the presence of an expandable component and changes in its abundance are difficult or impossible to detect by X-ray diffraction (XRD) using conventional polar interstratification compounds. Although high-grade diagenetic and very low-grade metamorphic (VLGM) samples do not expand in response to ethylene glycol treatment, the illite crystallinity of illitic materials improves significantly with increasing diagenetic/metamorphic grade. Previously, peak broadening for anchizonal to epizonal phyllosilicates was thought to be related entirely to stacking faults. The present study reexamines the contribution of expandable components to IC in shale and slate samples of Gaspe Peninsula in the high-grade diagenetic and VLGM zone using XRD patterns and high-resolution transmission electron microscopy lattice-fringe images after treatment with n-alkylammonium cations. These reveal the presence of expandable components including RI - and R3-ordered structures and “expandable illite”. was related to an increase in the crystallite size and a decrease in lattice strain. On the other hand, dehydration and collapse of smectite layers to 10 A layers during HRTEM investigations of ion-milled VS or illitic samples may have resulted in small amounts of expandable components having been overlooked. Furthermore, high-charge expandable components cannot be differentiated from non-expandable discrete illite in both conventional XRD and HRTEM (Vali et al. 1991). The present study aimed to assess qualitatively the presence of expandable phases in high-grade diagenetic-VLGM rocks of Lower Paleozoic sedimentary sequences in the Quebec Appalachians and its impact on IC measurements using conventional polar interstratification compounds; e.g., solvation with ethylene glycol (EG), and nalkylammonium intercalation.

1 INTRODUCTION

Illite crystallinity, measured as full width at half maximum (FWHM) of the 10 A X-ray diffraction peak of illitic materials, is an empirical index used to characterize the grade of diagenesis and incipient metamorphism in terranes where index minerals and diagnostic mineralogical assemblages are not available. Because of the simplicity of measuring illite crystallinity (hereafter IC), this property has been widely used in the past, generally in conjunction with other parameters of maturity, to establish diagenetic and very low-grade metamorphic (VLGM) conditions. Yet, the physical meaning of IC has remained largely unexplained. Arkai and Toth (1983), Nieto and Sanchez-Navas (1 994), and Warr and Nieto (1998) studied ion-milled samples with the high-resolution transmission electron microscope (HRTEM) to evaluate lattice strain, i.e., the various types of lattice imperfections such as stacking faults, edge dislocations, layer terminations and defects that give rise to non-periodic features of the crystal structure. These investigators found that expandable components did not play a role contributing to the lattice strain beyond the high diagenetic zone, probably because anchizonal illite shows no reaction to glycol treatment in XRD analyses. It was inferred that smectite-group phases and expandable mixed layers are absent in the anchizone, following Merriman et al. (1990). The continuous improvement of IC with increasing grade

2 MATERIALS AND METHODS Samples of argillaceous rocks from deep-water flysch deposits of the Ordovician Deslandes Formation and the Cambro-Ordovician Cap-desRosiers Group of the Taconian belt of Gaspe Peninsula ranging from high-grade diagenesis (8 samples) to VLGM (7 samples) from the vicinity of McGerrigle Mountains Pluton were chosen with the aid of IC and organic-matter maturation maps of Hesse and Dalton (1 99 1). 655

According to St. Julien and Hubert (1975), the external domain of the Taconian orogenic belt of the Gaspe Peninsula in the Quebec Appalachians belt comprises Cambro-Ordovician sediments derived from various portions of a former passive continental margin. The younger Ordovician rocks of this belt were deposited in a foreland basin created during the Taconian Orogeny. As far as igneous activity in the Taconian belt is concerned, the McGerrigle Mountains Pluton is an Upper Devonian-Lower Carboniferous post-orogenic intrusion surrounded by an exceptionally (up to 20 km) wide anchimetamorphic halo. Experimental settings of the XRD and HRTEM work are given in Shata (2000).

arrangement generally have a d(001) spacing of 29 to 31 8, (e.g., a -10 A 2:1 layer silicate + a 19 8, nc = 18 expanded interlayer) in XRD, whereas in HRTEM images, d(001) is measured as -24-25A. Treatment with alkylammonium cations (nc 2 12) produced two characteristic XRD reflections; a 10 8, peak representing non-expandable illite (packets of discrete non-expandable 2:1 layer silicate in TEM) and a 29-31 A peak (uniform coherent expandable 2:1 layers of 25 A spacing in HRTEM) representing highly charged "expandable illite" (Fig. 3).

3.2.2Vermiculite Only one sample (TM29) exhibits an expansion in response to nc=8 as well as nc=18 intercalation. The sample is characterized by three sharp XRD reflections, i.e., 29.4,14.5,and 9.97A after nc-18 treatment, whereas it expanded to 18.6 A upon intercalation with nc=8 (in contrast to expandable illite, which does not expand in response to nc=8). This may be interpreted as intermediate- to highcharge vermiculite with a paraffin-type arrangement of the alkylchains Lattice-fringe images for this sample (TM29) show a coherent pattern of hlly expandable interlayers with uniform spacing of 25 A (Fig. 4) alternating with a few very thin non-expandable 10 A 2:1layers.

3 RESULTS

3.1 XRD-characterization of glycolated and n-alkylammonium treated illitic and chloritic materials In air-dried samples, the high-grade diagenetic samples show a fairly broad asymmetrical basal (001) reflection at 10 A.After glycolation, the 10 8, reflection shifts slightly toward higher 28 angles, indicating the presence of expandable layers. On the other hand, the FWHM of the 10 8, peak narrows from the high-grade diagenetic zone to the anchizone, and within each zone with increasing diagenetic and metamorphic grade, indicating some structural modification. Anchimetamorphic illite of Gaspe Peninsula shows no significant XRD-shift upon glycol treatment (Figs 1 , 2). The details of the response of the 10 A reflection of illitic phases of Gasp6 Peninsula phyllosilicates to glycolation in both the coarse and fine fractions are given in Shata (2000,Tables 3-1,-2). The expandability of anchizonal illite was evaluated by comparing lattice-fringe images of 2:1 clay minerals in the 1.60) alteration were identified in Leg 115 ODP at Hole 7 15A. Oxidizing environment results from water-rock ratio in vesicular margins of lava flows. It is primarily marked by dark brown and brownish dark color of basalts. Oxidative and “non-oxidative” alteration occurred simultaneously as it was determined for basalts from Ninetyeast Ridge (Artamonov et al. 1998). Contrary to basalts from the East Indian Ocean (Legs 26 and 121), where the highest degree of alteration was determined in oxidized basalts and especially in highly vesicular varieties, we have not recognized this peculiarity in basalts from Hole 715A. Maximum H20’ content is up to 2.93 wt% in non-oxidized sample and 2.53 wt% in oxidized basalt. Secondary mineral assemblage in basalts from Leg 115 is significantly poorer than that from Ninetyeast Ridge. Secondary minerals are represented mainly by smectites and calcite, Fe-oxides and hydroxides, or sometimes sulfides. We have not recognized zeolites, K-feldspar, clay minerals with chloritic structures, and quartz. This poor association of secondary minerals corresponds with a lowtemperature alteration. Burns et al. ( I 990) determined low temperature of carbonate formation in these basalts. This is: 2 to 13*C at Hole 7 15A and 12 to 2OoC at Hole 707C. It is probable that all holes of Leg 115 penetrated basalts from marginal parts of edifices that are most distant from the eruptive centers. In these terms they are similar with Hole 7561) from Ninetyeast Ridge. As it was determined for the Ninetyeast Ridge, the various ages of sub-sea volcanic edifices of Mascarene Plateau (34 Ma, Hole 706C, and 64 Ma, Hole 707C), Chagos Bank (49 Ma, Hole 713A), and Maldives Ridge (57 Ma, Hole 715A), did not change significantly the degree of basalt alteration. Basalts from all holes of Leg I 15 are slightly altered. We have calculated the mass balance of major and minor elements in the studied basalts (Kazitzin & Rudnik 1968). This atomic-volume method implies recalculation of chemical analyses with due regard for their porosity and real packing of atoms in minerals. Alteration of basalts from holes drilled in Leg 115 resulted in slight mobility of chemical elements. In general, the behavior of chemical elements is similar to that in basalts from the Ninetyeast Ridge (Artamonov et al. 1998). The main peculiarity is in that oxidative alteration is accompanied mainly by accumulation of major and minor elements (including REE) while the non-oxidative alteration is accompanied mainly by loss of elements. 670

4 RESULTS AND DISCUSSION At present, investigators have not decided on a single point of view on formation of intraplate rises in Indian Ocean. It is supposed (Morgan I98 I , Duncan 1981, 1990), that the island of Mauritius, Mascarene Plateau, the Chagos Bank, the Maldive and Laccadive Ridges, and the Deccan traps (western India) are produced by Reunion stationary hotspot due to the motion of the Indian plate northward during Tertiary. The main pulse of the Deccan flood basalts occurred rapidly at about the Cretaceous/Tertiary boundary (Courtillot et al. 1986). Alternative models suggest that the lineament is a product of volcanic activity along a transform fault associated with Tertiary seafloor spreading (Fisher et al. 1971, McKenzie & Sclater 1971). Mayerhoff & Kamen-Kaye (198 1) consider, for example, that the Mascarene Plateau is a submerged island arc. At different times many hypotheses for formation of Ninetyeast Ridge were suggested (Francis & Raitt 1967, Le Pichon & Heirtzler 1968, McKenzie & Sclater 1971, Sclater & Fisher 1974, Neprochnov et al. 1979). There is a widely accepted hypothesis that the Ninetyeast Ridge, the Broken Ridge, the Kerguelen Plateau, and the Rajmahal Traps (eastern India) represent volcanic products of the longexisting (during last 120 1n.y.) hotspot which is located near the Kerguelen-Heard Islands (Luyendyck & Rennick 1977, Duncan 198 1, Morgan 1981). The hypothesis about predominant role that mantle plumes play in forming intraplate rises in Indian Ocean is commonly accepted. However, it is difficult to explain the block morphology (Udinzev 1987) and geochemical features of basalts of this rises within the scope of this hypothesis. We suggest that the formation of the studied rises occured by powerful streams of the basaltic lava in the large fault zones over the already formed oceanic crust. The geochemical data suggest that majority of basalts of intraplate rises are derived from mantle melts generated at more depth than melts of N-MORB, but at less depth than melts of ocean island tholeiites (OIB). Fracture zones are developed from continent to middle parts of the ocean. In this direction occur movement of volcanic activity and change of the depth of generation of primary melts. Basalts from aseismic structures of the oceanic floor are altered to various degrees. Alteration of basalts varies from slight to high in both intrasites and intersites. This patchy pattern in alteration of basalts results from various permeability and crystallinity of basalts. Two types of low-grade alteration “nonoxidative” and “oxidative” occur in basalts from aseismic ridges, rises, and plateaus. Basalts altered in oxidizing environment demonstrate stable association of secondary minerals as follows: brown and greenish-yellow smectites, Fe-oxides, and hydrox-

ides, calcite, and sporadically K-feldspar. Basalts altered in non-oxidizing environment demonstrate predominance of trioctahedral green smectites. Basalts from highly permeable vesicular lava flows erupted in subaerial environments suffer alteration in both oxidizing (along vesicular margins of flows and along cracks) and “non-oxidizing” (inner parts of lava flows and unoxidized fragments at margins of flow units) environments. Basalts from lava flows erupted in deep-water environment suffer non-oxidative alteration as effusive bodies lack margins highly permeable for the seawater. Alteration of basalts occurred during the waning stage of volcanic activity at various temperatures depending on the distance from eruptive centers. Various ages of the studied subaqueous volcanic edifices demonstrate no significant influence on alteration of basalts. Main alteration of basalts occur under the influence (and during the existence) of hydrothermal systems within volcanic edifices. After the edifice cools, hydrothermal alteration of basalts ceases almost completely. Oxidative type of alteration is accompanied by gain of most major and minor elements, including REE. In contrast, “non-oxidizing” type of alteration is characterized by loss of the matter. The obtained results probably reflect real redistribution of elements at alteration of basalts in aseismic structures of the oceanic floor. ACKNOWLEDGMENTS We thank ODP for providing the samples. This study was financially supported by the Russian Foundation for Fundamental Research (grants 98-05-64856 and 99-05-65462). REFERENCES Artamonov, A.V., Kurnosov, V.B., & B. P. Zolotarev 1998. Alteration of basalts from the Ninetyeast Ridge, Indian Ocean (ODP data). In G.B. Arehart & J.R. Hulston (eds), Water-Rock Interuction: 7 I 1-7 14. Rotterdam: Balkema. Backman, J., Duncan, R.A., et al. 1988. Proc. ODP, Init. Repts., 115. College Station, TX (Ocean Drilling Program). Baxter, A.N. 1990. Major and trace element variations in basalts from Leg 115. In R.A. Duncan, J . Backman,. L.C. Petreson et al. (eds), Proc. ODP. Sci. Results, 115: 11-21. College Station, TX (Ocean Drilling Program). Bums, S.J., Swart, P.K., & P.A. Baker 1990. Geochemistry of secondary carbonates in Leg I IS basalts: tracers of basaldseawater interaction. In R.A. Duncan, J. Backman, L.C. Petreson et al. (eds), Proc ODP. Sci. Results, I IS: 93- 10 I . College Station, TX (Ocean Drilling Program). Courtillot, V., Besse, J., Vandamme. D., Montigny, R., Jaeger, J.-J., & H. Cappeta 1986. Deccan flood basalts at the Cretaceous-Tertiary boundary? E ~ i r t hPIunet. Sci. Lett. 80: 36 1 374. Duncan, R.A., 1981. Hotspots in the southern oceans - an absolute ftame of reference for motion of the Gondwana continents. Tectonophysics 74: 29-42.

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Duncan, R.A., 1990. The volcanic record of the Reunion hotspot. In R.A. Duncan, J. Backman, L.C. Petreson et al. (eds) Proc. ODP, Sci. Results, 115: 3-10. College Station, TX (Ocean Drilling Program). Duncan, R.A. & R.B. Hargraves 1990. 4oAr/39Argeochronology of basement rocks from the Mascarene Plateau, the Chagos Bank, and the Maldives Ridge. In R.A. Duncan, J. Backman, L.C. Petreson et al. (eds) Proc. ODP, Sci. Results, 1 15: 43-5 1. College Station, TX (Ocean Drilling Program). Fisher, R.L.. Sclater, J.G., & D.P. McKenzie 1971. Evolution of the Central Indian Ridge, Western Indian Ocean. Geol. Soc. Atn. Bull. 82:553-562. Fisk, M. R., Upton, B.G.J., Ford, C. E., & W.M. White 1988. Geochemical and experimental study of the genesis of magmas of Reunion Island, Indian Ocean. J. Geophys. Res. 93: 4933-4950. Francis, T.J.C. & R.W. Raitt 1967. Seismic refraction measurements in the Norhwest Indian Ocean . J. Geophys. Res. 7 1 : 427-449. Kazitzin, Y . & V. Rudnik 1968. Guide to estimation of mass balance and inner energy during tnelnso- matic rock formation. Nedra: Moscow. Le Pichon X. & J.R. Heirtzler 1968. Magnetic anomalies in the Indian Ocean and sea- floor spreading. J. Geophys. Res. 73: 2109-21 17. Lightfoot, P. & C. Hawkesworth 1988. Origin of Deccan Traps lavas: evidence from combined trace element and Sr-, Ndand Pb-isotope studies. Earth Planet. Sci. Lett. 9 1 : 89- 104. Luyendyck, B. P. & W. Rennick 1977. Tectonic history of aseismic ridges in the eastern Indian Ocean. Geol. Soc. Am. Bull. 88: 1347-1356. McKenzie, D.P., & J.G. Sclater 1971. The evolution of the Indian Ocean since the Late Cretaceous. Geophys. J. R. Astron. Soc. 25:437-528. Meyerhoff, A.A. & M. Kamen-Kaye I98 I . Petroleum prospects of the Saya de Malha and Nazareth Banks, Indian Ocean. Am. Assoc. Pet. Geol. Bull. 65: 1344-1347. Morgan, W.J. 1981. Hotspot tracks and the opening of the Atlantic and Indian oceans. In C. Emiliani (ed.), The Sea, 7: 443-487. Wiley: New York. Neprochnov, Y.P., Merlin, L.R., Shreider, A.A., Sedov, V.V.& I.N. Elnikov 1979. Structure of the East Indian Ridge based on the integrated geophysical studies. Oceanology 4: 644657. Sclater, J.G. & R. L. Fisher 1974. The evolution of the east central Indian Ocean with emphasis on the tectonic setting of the Ninetyeast Ridge. Geol. Soc. Am. Bull. 85: 683-702. Udinzev, G.B. 1987. Relief and strucfure of ocean floor. Nedra: Moscow.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Multiple fluid-flow events and mineralizations in SW Sardinia: an European perspective M.Boni & A.Iannace Dipartimeizto Scienze della Terra Universita di Napoli, Italy

I .M .Villa Isotopengeologie, Bern, Switzerland

L.Fedele & R.Bodnar Department Geological Sciences Blacksburg, VA, USA

ABSTRACT: Between the end of Variscan and the beginning of Alpine, SW Sardinia was affected by several hydrothermal phases, comparable with those occurring in other parts of Europe. They resulted in a widespread hydrothermal dolomitization of Lower Palaeozoic carbonates, and in a range of base metal-Ba-F vein mineralizations. By comparing the available geological and geochemical data with the absolute geochronology of hydrothermal silicates, at least two main hydrothermal stages can be hypothesized: A) the first one at +270 Ma (Middle Permian). This phase has been recognized in many other hydrothermally affected areas of western and central Europe; B) the second one can only be narrowed down to the time interval from Upper Permian through the Mesozoic (probable age for Sardinia: ' ~ 2 3 0 )This . event, characterized by high salinity fluids, occurred in slightly different ages throughout Europe, depending on the paleogeographic position in regard to the future Tethyan margins.

1 GEOLOGICAL SETTING The evolution of the Sardinian Palaeozoic basement shows structural, stratigraphical and geochemical analogies with other European Variscan Belts. In the Iglesiente-Sulcis area (Fig. l), considered as the external zone of the Sardinian Variscan chain, an incomplete Palaeozoic succession, spanning in age from Early Cambrian to Devonian, underwent at least two compressional and one extensional phase of deformation, followed by granite intrusions and late-Variscan basement uplift. The widespread, calcalkaline granitoid bodies are attributed both to synand post-collisional Variscan stages, with leucogranite types marking the end of the sequence (throughout 330 to w 290 Ma, Late Carboniferous to Permian, Del Moro et al. 1975, Guasparri et al. 1984, Secchi et al. 1991, Boni et al. 1999). The post-Variscan sedimentary record, though poor, starts with a clear trend toward crustal attenuation and tearing, possibly related to a large transcurrent megashear zone (Cassinis et al. 1999), and resulting in the inset of Upper Carboniferous-Permian continental basins, containing coeval magmatic products of still calc-alkalic affinity. Porphyry stocks, ignimbrite flows, rhyodacitic lavas and basaltic dykes occur throughout the whole island within an age-interval estimated between 280 and 250 Ma (Atzori & Traversa, 1986, Beccaluva et al. 1981, Cozzupoli et al. 1984, Edel et al. 1981, Lom-

bardi et al. 1974, Vaccaro et al. 1991). It must be taken into account, however, that the large spread of age data could result either from the widespread hydrothermal alteration of the magmatites, or from the uncertainity due to imprecise and/or inaccurate dating techniques. The Mesozoic successions, though incomplete and sporadic, point to further crustal thinning (a prelude to the Alpidic rifting), evidenced by the Middle-

Figure 1. Geological sketch map of SW Sardinia with some localities mentioned in text. The main mineralized area is located in the Palaeozoic lithologies around the town of Iglesias (from Boni et al. 1999).

673

Upper Triassic marine sediments. The magmatic counterparts of this extension could be identified in the sub-alkalic and alkalic dykes (Atzori & Traversa, 1986), with a mantle Sr-signature, occurring mostly in the northern part of the island, and dated by Vaccar0 et al. (1991) at about 230 Ma. The condensed Triassic successions extend upward into thicker (mostly in northern and eastern Sardinia) Jurassic and Cretaceous carbonates. This briefly sketched sedimentary and magmatic setting in Sardinia could be interpreted in a broader evolutionary frame of stepwise changes from lateVariscan still orogenic regimes (Middle Permian), to post-orogenic, Alpidic (Middle Triassic) tensional stages. This evolution is better documented along the marginal domains of the European plate (where the future Tethys ocean was bound to develop), e.g. in the Carnian Alps, along the Insubric line and in the sedimentation realm of the Betics, Calabria and Maghrebides. In the more cratonic domains of central and western Europe, the change to Alpidic tensional stages, involving the inset of widespread marine sedimentation and higher subsidence rates, should be rather shifted to Jurassic.

magmatic biotites and feldspars, and (b) determining the age of newly segregated Ba-silicates. The calculated age of this hydrothermal phase Ma, at least 30 Ma later than corresponds to ‘ ~ 2 7 0 the youngest Variscan granitoid intrusions, at a time when the shallow intrusives had no residual heat left to fuel the hydrothermal circulation system. This same age, recorded by geochronological methods also in several magmatic dykes of calc-alkalic character throughout Sardinia (see Vaccaro et al. 1991 and references therein), could fit with the Permian post-orogenic extensional phase and wrench fault tectonics, which caused widespread hydrothermal alteration in the basement rocks, locally coupled with important mineralizations in all of Europe. Among the recent literature on this subject, we can mention Boni et al. (1992 and references therein), Calvet et al. (2000 and references therein), Gbmez-Fernindez et al., (2000), Schneider & Haack, (2000a), Tornos et al. (2000). We think that this Permian hydrothermal stage produced in SW Sardinia, besides barium silicates (celsian & armenite), high temperature-low salinity, heavily radiogenic Fe-Cu-Zn-Pb-F-Ba vein ores (Su Zurfuru, Santa Lucia, Montega etc.). This conclusion is mainly based on the geochemical analogies of the precipitating fluids, as deduced from the fluid inclusions and stable isotopes. (2) The pervasive hydrothermal dolomitization (locally known as “Dolomia Geodica”), replacing extensive areas of the Cambrian and Ordovician carbonates, is indicative of a large fluid flow, controlled in SW Sardinia by the Variscan foliation and cleavage planes (Boni et al. 2000a). The “Dolomia Geodica” is comparable to similar kind of dolomites, occurring in northern Spain (GbmezFernindez et al. 2000, Boni et al. 2000b), Belgium (Nielsen et al. 1998), Ireland (Gregg et al. 1999), and other areas of continental Europe. No absolute dating was possible: the relative age of the “Dolomia Geodica” can only be inferred by the crosscutting relationships of younger Pb-Ag-Ba low temperature veins on the epigenetic dolomites. Though the low temperature-high salinity nature of the dolomitizing fluids (similar to that of the younger veins), in consideration of the pervasive nature of this phenomenon and of its lithologic control, we are inclined to assign also a Permian(?) age to the hydrothermal dolomitization, as in other European late Variscan domains (Gbmez-Fernindez 2000, Nielsen et al. 1998). (3) 39Ar/40Aranalysis showed also that younger hydrothermal phases, related to further episodes of fluid flow, might have occurred through the Mesozoic, apparently as late as Cretaceous (Boni et al. 1999). No unambiguous age determination of these younger phases was possible so far. The

2 REVIEW OF THE CHARACTERISTICS OF THE HYDROTHERMAL STAGES As already reported in Boni et al. (1992, 1999, 2000a), several quite distinct hydrothermal systems were supposedly active in SW Sardinia not only at the lower time fringe of the Variscan magmatism (skarn and retrograde contact-metamorphism), but spanned from Permian to Mesozoic, resulting in a variety of hydrothermal products, including ore deposits. Due to the established poor sedimentary record of this time interval in the southwestern area of the island, the exact timing of each of the fluid flows can be only hypothesized. Therefore, a possible dating of these events in Sardinia should be based not only on the geochemical characterization of the ores and on absolute geochronology, but also on the comparison to time-related flows in other European areas (e.g. north-eastern Spain and southern France), sharing with Sardinia a similar geologic evolution. Further hydrothermal activity, not discussed in this paper, is also related to the Tertiary magmatism, resulting in the recently discovered epithermal Low sulfidation and High sulfidation gold and base metal mineralization. (1) One first prominent episode of hydrothermal (Boni mineral formation has been dated by 39Ar/40Ar et al. 1999). That study analyzed granitic and vein minerals from SW Sardinia with the aim of (a) checking the effects of the hydrothermal fluids on

674

39Ar/40Arresults demonstrate a very complex

slightly different ages depending on the location of the deposits in regard to their paleogeographic position. The older events (Middle-Upper Triassic) seem to have happened in areas nearer to the future Tethyan margins, as Sardinia and Eastern Spain. The younger (Jurassic) ones occurred in the more cratonic areas, like France and Germany.

superposition of hydrothermal alteration events at different times (including Tertiary). However, we note that an age of 230 Ma was assigned to the alkali dykes by Vaccaro et al. (1991), and interpreted by the same authors as the first evidence of Alpidic rifting in Sardinia. We propose to relate this rifting phase to the inset of the low temperature-high salinity, poorly radiogenic Pb-AgBa vein- and paleokarst deposits, occurring in Iglesiente and Sulcis (the “Ricchi Argento” ores, but also part of the Montevecchio vein system), controlled by a younger set of fractures. Unfortunately, every attempt of direct dating ore and gangue minerals related to this hydrothermal phase has failed so far. This kind of low temperature-high salinity ores are ubiquitous in Europe, even if their ages appear to range from Triassic to Jurassic, depending on their original position with regard to the European intraplate geometry (Boni et al. 1992 and references therein; Canals & Cardellach 1997 and references therein, Meyer et al. 2000, Schneider & Haack 2000b, Tornos et al. 2000).

REFERENCES Atzori, P. & G. Traversa 1986. Post-granitic Permo-Triassic dyke magmatism in eastern Sardinia (Sarrabus p.p., Barbagia, Mandrolisai, Goceano, Baronie and Gallura). Per. Mineral. 55: 203-231. Beccaluva, L., Leone, F., Maccioni, L. & G. Macciotta 1981. Petrology and tectonic setting of the paleozoic basic rocks from Iglesiente-Sulcis (Sardinia, Italy). N.Jb.MinerAbh. 140(2): 184-201. Boni, M., Balassone, G. & I.M. Villa 1999. Age and evolution of granitoids from South West Sardinia: genetic links with hydrothermal ore bodies. Proc. Fifth Biennial SGA Meeting “Mineral Deposits: Processes to Processing” Stanley, C.J. et al. Editors. V01.2, London, August 1999: 1255-1258. Boni, M., Iannace, A., Koppel, V., Hansmann, W. & G. FriihGreen 1992. Late- to post-Hercynian hydrothermal activity and mineralization in SW Sardinia. Econ.Geo1. 87(8): 21 132137. Boni, M., Parente, G., Bechstadt, T., De Vivo, B. & A. Iannace 2000a. Hydrothermal dolomites in SW Sardinia (Italy): evidence for a widespread late-Variscan fluid flow event. Sedimentary Geology 131(3-4): 181-200. Boni, M., Iannace, A., Bechstadt, T. & M. Gasparrini 2000b. Hydrothermal dolomites in SW Sardinia (Italy) and Cantabria ( N W Spain): evidence for late- to post-Variscan fluid flow events. Proceedings GeofZuids Meeting, Barcellona July 2000, Journ.Geochem.Expl.69-70: 225-228. Calvet, F., Canals, A., Cardellach, E., Carmona, J.M., G6mezGras, D., Parcerisa, D., Bitzer, K., Roca, E. & A. Travt. 2000. Fluid migration and interaction in extensional basins: application to the Triassic and Neogene rift in the central part of the Catalan Coastal Ranges, NE Spain. Field Trip Guidebook, Geofluids 111 2000, July 2000 Barcellona: 58 PP. Canals, A. & E. Cardellach 1997. Ore lead and sulphur isotope pattern from the low-temperature veins of the Catalonian Coastal ranges (NE Spain). Mineral. Deposita 32(3): 243249. Cassinis, G., Cortesogno, L., Gaggero, L., Pittau, P., Ronchi, A. & E. Sarria 1999. Late Palaeozoic continental basins of Sardinia. Field Trip Guidebook International Field Conference on “The continental Permian of the Southern Alps and Sardinia (Italy). Regional Reports and general correlation. 15-25 September 1999, Brescia: 116 pp. Cozzupoli, D., Gerbasi, G. Nicoletti, M. & C. Petrucciani 1984. Eta WAr delle ignimbriti permiane di Galtelli (Orosei, Sardegna Orientale). Soc.Mineralog.Petrolog.lta1iana Rend. 39: 471-476. Del Moro, A., Di Simplicio, P., Ghezzo, C., Guasparri, G., Rita, F. & G. Sabatini 1975. Radiometric data and intrusive sequence in the Sardinian Batolith. N.Jb.Miner.Abh. 126: 28-44. Edel, J.B., Montigny, R. & R. Thiuzat 1981. Late Palaeozoic rotations of Corsica and Sardinia. New evidence from Paleomagnetic and WAr studies. Tectonophysics 79: 210223.

3 CONCLUSIONS

Between the end of Variscan orogeny and the beginning of the Alpine cycle, southwest Sardinia was the site of several hydrothermal phases, comparable with those occurring in other parts of central and western Europe. They resulted in a widespread hydrothermal dolomitization (“Dolomia Geodica”) of the Lower Palaeozoic carbonates, and in a range of base metalBa-F vein- and paleokarst mineralizations showing distinct characteristics. By comparing all the available geological and geochemical data, as well as the absolute geochronology of both the hydrothermal minerals and of the primary magmatic phases measured in the late Variscan - Early Alpine bodies, at least two main hydrothermal stages could be identified: A) The first one took place at about 270 Ma (Middle Permian). This phase has been recognized in many other hydrothermally affected areas of eastern and central Europe, and seems to mark quite similar geodynamic conditions among even distant domains. B) The second one can only be narrowed down to the time interval from Upper Permian through the Mesozoic. Because the first Alpidic rifting stages in Sardinia (evidenced by the alkaline dykes and marine transgression) have an age set of 230 Ma (Middle Triassic), we think that this could be also the age for the second main hydrothermal event. It is interesting to note that this hydrothermal event, characterized by highly saline fluids, and producing comparable ore deposits all over Europe, has 675

Gomez-Fernandez, F., Both, R.A., Mangas, J., & A. Arribas 2000. Metallogenesis of Zn-Pb carbonate-hosted mineralization in the Southeastern region of the Picos de Europa (Central Northern Spain) Province: geologic, fluid inclusion and stable isotopes studies. Econ.Geo1. 95(1): 1940. Gregg, J.M., Shelton, K.L., Johnson, A.W., Somerville, I.D. & W.R. Wright 1999. Diagenetic and hydrothermal dolomitization of the Waulsortian Limestone (Carboniferous) in the Irish midland (Abstract) Geol.Soc.of America, Abstracts with Programs 31: n.7. Guasparri, R., Riccobono, F. & G. Sabatini 1984. Considerazioni sul magmatismo intrusivo ercinico e le connesse mineralizzazioni in Sardegna. Rend.Soc.Ita1. Mineral. e Petr. 32: 17-52. Lombardi, G., Cozzupoli, D. & M. Nicoletti 1974. Notizie geopetrografiche e dati sulla cronologia WAr del vulcanesimo tardo-paleozoico sardo. Per.Mineralogia 43: 221 -312. Meyer, M., Brockamp, O., Clauer, N., Renk, A. & M. Zuther 2000. Further evidence of a Jurassic mineralizing event in central Europe: WAr dating in hydrothermal alteration and fluid inclusion systematics in wall rocks of the Kafersteige fluorite vein deposit in the northern Black Forest, Germany. Mineral. Deposita 35(8): 754-761 Nielsen, P., Swennen, R., Muchez, Ph. & E. Keppens 1998. Origin of zebra dolomites from the Dinantian south of the Brabant-Wales Massif, Belgium. Sedimentology 45: 727743. Schneider, J. & U. Haack 2000a. A different kind of Pb-Pb age. Proc. 78. Jahrestagung Deutsche Mineral. Gesellschaji Meeting, September 2000, Heidelberg, Abstract: p. 188. Schneider, J. & U. Haack 2000b. Direct Rb-Sr dating of sandstone-hosted sphalerites. Proc. 78. Jahrestagung Deutsche Mineral. Gesellschaji Meeting, September 2000, Heidelberg, Abstract: p. 189. Secchi, F.A., Brotzu, P. & E. Callegari 1991. The Arburese complex (SW Sardinia, Italy). An example of dominant igneous fractionation leading to peraluminous cordieritebearing leucogranites or residual melts: Chem.Geo1. 92: 213-249. Tornos, F., Delgado, A., Casquet, C. & G. Galindo 2000. 300 million years of episodic hydrothermal activity: stable isotope evidence from hydrothermal rocks of the Eastern Iberian Central System. Mineral. Deposita 35(5): 551-569. Vaccaro, C., Atzori, P., Del Moro, A., Oddone, M., Traversa, G. & I.M. Villa I.M. 1991. Geochronology and Sr-isotope geochemistry of late Hercynian dykes from Sardinia. Schweiz.Minera1.Petrogr.Mitt. 71: 227-235.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

High-pressure melting and fluid flow during the Peterrnann Orogeny, central Australia 1.S .Buick Department of Earth Sciences, La Trobe University, Bundoora, Vic. 3086, Australia

D .Close, I .Scrimgeour & C .Edgoose Northern Territory Geological Survey, Alice Springs, N.T. 0871, Australia

J.Miller & C .Harris Department of Geological Sciences, University of Cape Town, Rondebosch 7700,South Africa

I .Cartwright Department of Earth Sciences, Monash University, Clayton, Vic. 3168, Australia

ABSTRACT: During the -0.56 Ga Petermann Orogeny, anhydrous -1.3 Ga granulites, -1.2 Ga granites, and post-1 .l Ga mafic dykes were deformed in an intraplate transpressional setting. At the deepest levels, these rocks were variably deformed at sub-eclogite-facies conditions (-1 1-13 kbar and -750 "C) but remain unmelted. In contrast, at higher crustal levels, similar deformed country rocks can be traced into discrete shear zones that contain hydrous migmatites formed at 7-1 1 kbar and -700-650 "C. At the deepest levels, granites and mafic dykes have 618O(w~lvalues (-7-8%0 and 6-8%0, respectively) appropriate to likely precursors. values (-4 %o However, at higher crustal levels, metagranites and, to a lesser extent, metabasites have 6180(W~) to +6 %O SMOW and 4 to 5%0 SMOW, respectively) that reflect interaction with low-'*O fluids. The patterns of isotopic resetting do not closely relate to the extent of shearing and partial melting, and probably resulted from pre-0.56 Ga hydrothermal alteration. 1 INTRODUCTION The role of fluids in the development of medium- to high-temperature (>650"C) high-P (eclogite- or subeclogite facies) rocks is contentious. Although earlier believed to result from "dry" metamorphism, a number of studies have shown that synmetamorphic fluid infiltration, coupled with deformation, is necessary to trigger the development of such high-P assemblages from magmatic precursors (eg. Rubie 1990). Alternatively, in variably hydrothermally altered igneous rocks the subsequent development of high-pressure assemblages may occur preferentially in those rocks whose protoliths were most hydrated prior to metamorphism (Barnicoat & Cartwright, 1997). These two different scenarios require very different extents of syn-metamorphic fluid-rock interaction. The wet solidi for granites or quartz-bearing tholeites have temperature minima in the interval -620-680°C at -10-15 kbar (Johannes & Holtz 1996), suggesting that high-Phigh-T terrains may undergo wet-melting if sufficient fluid is available. Water-saturated quartzofeldspathic melts have high dissolved H20 contents (16-21 wt % at 10-20 kbar; Johannes & Holz 1996). Therefore, water-saturated partial melting in high-pressure terrains requires high time-integrated fluid fluxes. However, sources of fluids, and how they reach extreme crustal depths, is commonly unclear. In the absence of syn-high pressure metamorphic fluid flow, anatexis will occur

under fluid-absent conditions, and will be restricted to those rocks that are most fertile (have a higher abundance of hydrous minerals). In this study we investigate relationships between partial melting, fluid flow and oxygen isotopic resetting during high-pressure intraplate metamorphism of latest Neoproterozoic/ Cambrian age in the Petermann Orogen, central Australia.

2 GENERAL GEOLOGY During the late Neoproterozoic to early Palaeozoic (-0.56 Ga) Petermann Orogeny, basement rocks of the Musgrave Inlier were thrust northwards over the southern margin of the Amadeus Basin during dextral intraplate transpression involving at least 125 km of shortening (Scrimgeour & Close 1999). The deepest crustal levels of the Petermann Orogen (immediately to the north of the Mann Fault; Fig. 1) occur in the Mann Terrane (Fig. l), which comprises -1.3-1.2 Ga (MI-DI) orthogneiss-dominated granulites intruded by a range of late syn- to postorogenic igneous rocks (-1.18 Ga, clino- and orthopyroxene-bearing I-type granites; -1.07 Ga mafic dykes, and -0.8 Ga mafic dykes: Sun et al. 1996; White et al. 1999). At these crustal levels, the rocks contain a variably-developed, but generally pervasive, D2a protomylonitic fabric formed under transitional garnet granulite- to eclogite-facies conditions (-1 1-13 kbar, -750 "C), and related to 677

Figure 1. Regional geological setting of the Palaeo- to Mesoproterozoic Musgrave and Arunta Inliers, and the intracratonic Neoproterozoic to Palaeozoic Amadeus and Officer Basins in central Australia. High-pressure anatexis occurs in DZb shear zones within the Mann Terrane (inset box; modified after Scrimgeour & Close 1999).

north-vergent D2a Petermann-age deformation (Scrimgeour & Close 1999). At higher crustal levels, further to the north, the Dza foliation in the same rock types can be traced into discrete D 2 b shear zones that contain abundant migmatite. 3 PETROLOGY Rocks with the highest-pressure D2a protomylonitic fabrics have recrystallised to transitional highpressure granulite- to eclogite-facies assemblages. Sheared and recrystallised mafic dykes contain garnet, rutile, clinopyroxene, kyanite, clinozoisite, and locally sodic plagioclase, whereas recrystallised metagranites contain variably recrystallised igneous quartz, plagioclase, K-feldspar, clinopyroxene f orthopyroxene and magnetite, with growth of finegrained garnet, hornblende, biotite and titanite. In particular, K-feldspar occurs as incipiently recrystallised porphyroclasts with grain diameters that commonly exceed 2-3cm. Metagranites at intermediate crustal levels somewhat contain similar Dza assemblages as those at the deepest crustal levels. In addition, within D 2 b shear zones they contain variably deformed leucosomes of quartz, K-feldspar, plagioclase, hornblende, biotite, garnet, titanite, allanite and apatite. Garnet and hornblende occur as coarsegrained poikiloblasts that contain abundant inclusions of titanite and allanite. Relict igneous Kfelspar persists as centimetre-diameter porphyroclasts in many of the leucosomes. Migmatitic amphibolites (ex-mafic dykes) contain leucosomes of plagioclase and quartz within a matrix of hornblende, plagioclase, garnet, clinopyroxene, plagioclase, titanite and ilmenite.

Pressures and temperatures of shearing ranged from 11-13 kbar and -700 "C between the Mann Fault and Woodroffe Thrust (Fig. 1) to -7 kbar and -650°C immediately to the north of the Woodroffe Thrust (Scrimgeour & Close 1999) ie under highpressure amphibolite-facies conditions. Migmatites are found only within the D2b shear zones and can be traced outside into the deformed, unmigmatised, and less hydrous metagranites and mafic dykes. Migmatitic leucosomes within the D 2 b shear zones are themselves variably overprinted by mylonitic fabrics developed during continued D 2 b deformation. 4 OXYGEN ISOTOPE GEOCHEMISTRY Metamorphosed granites and mafic dykes show a wide range of whole rock 6l80 values (metagranites: -4 to +8.0 %o, n = 38, with one exceptionally low value of -16.1%o; recrystallised mafic d kes: +4.4 to +8.6 %o, n = 19; Fig. 2). The highest 6' l7O(WR) values generally occur in the southemmost (structurally deepest levels) portion of the Mann Terrane, where the rocks are unmelted and contain patchilydeveloped, transitional eclogite-facies mineral assemblages (metagranites: +7.3 to +8.0 %o; mafic dykes: +5.9 to +8.6 %o, typically +6 to +7 %o). The 6 l 8 0 ( w ~values ) of the deepest-level metabasites and metagranites are similar to the range expected for igneous precursors (continental intraplate basalts and I-type granites, respectively; Fig. 2; Taylor & Sheppard 1986; Hoefs 1997). The lowest ~"O(WR) values occur at the slightly higher crustal levels that underwent extensive partial melting in discrete D2b shear zones. For both metabasites and metagranites the 6 1 8 0 ( ~values ~ ) at these crustal levels are significantly lower than those 678

shear zones commonly have lower 6 I 8 0 ( w ~ values ) than adjacent recrystallised mafic dykes; and 4) nonmigmatitic metagranites well away from DZb shear zones locally have 6 l 8 0 ( w ~values ) as low as +2 to +3%0 (Fig. 2). In contrast to metagranites, garnet amphibolites in the same D2b shear zones show quite ~ ~ ) (+4 to +5%0) that are homogeneous 6 1 8 0 ( values typically only -2%0 lower than incipiently recrystallised mafic dykes. This is surprising given that garnet amphibolites would be expected to have lower 6 1 8 0 ( ~values ~ ) than interlayered migmatitic metagranite if precursors (granite and mafic dykes) had equilibrated with a common 10w-'~Ofluid at moderate to high temperatures. This may reflect initial differences in permeability, such that most fluid flow occurred through the granites.

5 DISCUSSION AND CONCLUSIONS The 6 l 8 0 ( w ~ values ) of high-pressure amphibolitefacies metagranites and, to a lesser extent, recrystallised mafic dykes are abnormally low. Metagranites with 6 l 8 0 ( w ~ 1000 t Ag) is located in the Xionger terrain of the Huaxiong block, was formed in Mesozoic collisional orogenesis between the North China and South China paleocontinents. Therefore, it is an ideal example to study collisional petrogenesis, metallogenesis and fluidization (simplified as CPMF). The authors have studied the geology and geochemistry of this deposit in detail. In this paper, we report the geochemistry of its ore-forming fluid.

3 FLUID INCLUSION GEOCHEMISTRY

3.1 Morphology and thermometry of inclusions

2 GEOLOGIC SETTING & ORE GEOLOGY During Mesozoic collision between the North China and South China continents, the Machaoying fault, the southern boundary fault of Xionger terrain, trending EW 200 km in length and 34 38 km in depth, acted as a northward intracontinental subduction zone. In its north side, i.e. the Xionger terrain, occur in sequence: form south to north the lode AdAg ore zone, including more than 10 significant fault-bound deposits such as Tieluping deposit; the Yanshanian batholithic granite zone, and the porphyry-breccia and associated ore zone from south to north. All these orebodies and Yanshanian granitoids occur in the crystalline basement, or in the

-

Heating-stage microscope study shows that the fluid inclusions within minerals mainly include gas-liquid and liquid inclusions. The size of inclusions increases from early to late. Inclusions of stage I11 are less than 1 pm; those of stage I are ranging in 2 6 pm. The inclusions of stage I are so small that we can get only one temperature of 373 "C (Table l), which is the highest. Considering that the mineralizing temperatures of stage I for the other gold or silver deposits in the Xionger terrain range 300 450 "C, we think the value of 373 "C is reliable. The samples of stage 111 yields 10 values of temperature in range from 135 "C to 203 "C, with the average of 158 "C. The samples of stage I1 yield 7 values in range of 193 249 "C, with the average of 223 "C. Sample TS13 yields 7 values,

-

-

-

689

Table 1. Gaseous ( 10-6moVg)and aqueous (pg/g) composition in fluids and isotopes of the Tieluping Ag deposit.

Stage Mineral H20 CO2 CH4 CO N2 H20lC02 CO2lCH4 CH4lCO c021c0 Na' Ki Ca" Mg2' F-

c1so42-

CMo* CM' * EM- * 6l3cCo2(%o) 6D (%o) 6 l 8 0 M (%o)

paw** (4 0 T "C

I: rangelaverage (4) quartz

I1 (TS17) quartz

11.046 124.469 1 68.429 6.147 9.443 17.822 0.013 - 0.028 10.018 0.040 - 0.123 I 0.080 0 - 0.092 10.023 4.7 - 14.0 I 11.5 337 610 1435 0.19 - 0.35 I 0.25 66- 1941 117 2.28 - 10.83 I8.11 10.08 - 25.00 I 19.00 0.06 1.08 10.48 0.14 0.50 I 0.28 0.17 1.0 10.51 7.98 21.39 115.63 12.12 - 26.77 I 19.41 19.047 - 134.063 176.371 21.06 34.62 127.88 21.11 48.58 135.64 0.3 4.0 12.0 -96 -84 1-90 13.72 15.60 I 14.76 7.90 9.78 18.94 3 73

164.608 10.648 0.073 0.107 0 15.5 146 0.68 99 15.50 36.25 0.08 0.17 0.55 27.60 35.35 175.436 52.00 63.50

II+III (TS13) quartz 84.279 4.647 0.023 0.048 0 18.1 202 0.48 97 6.88 28.25 0.43 0.06 0.94 27.26 6.82 88.997 35.62 35.02

-109 13.00 1.79 233

0.1 -73 10.10 -4.19 185

-

-

-

-

-

111: rangelaverage (4) carbonate 30.228 - 69.715 149.336 3.484 4.741 14.127 0 0.015 10.009 0.028 0.057 10.042 0.031 0.04 31 0.035 8.7 14.7 I 11.8 282 431 I 373 (3) 0 0.39 10.25 61 169 I 109 2.50 7.00 13.64 4.59 10.75 17.86 unrneasured unrneasured 0.35 - 0.55 10.42 3.86 8.66 15.97 2.85 - 5.51 13.83 33.800 - 74.528 153.549

-

-

-

-

9.29 12.34 I10.22 -2.0 - -0.3 I -1.3 -88 - 60 I -73 10.50- 11.801 11.14 -1.58 - -0.18 I -0.92 158

* CM-, CM-and CMO are sums of contents of cations, anions (pg'g) and molecules (IO-6moVg)respectively. **

10001nc(~.~~= 3.57

x

106T'- 2.73, 10001nac,I,~~v = 2.78 x 10GT-?-2.89 (from Zhang 1985).

highest ionic concentration, showing that it contributed most to mineralization among the three stages, which is consistent with the ore-exploration experience. The contents of H20, COz, CO and CH4 in the minerals of stage I1 are obviously more than those of stage I and 111, which proves that the stageI1 minerals captured more fluid when they crystallized. The capture of a great deal of inclusions reflects that the crystallization is rapid, which could be further evidenced by facts such as complicated mineral assemblage (e.g. polymetallic sulfides), fine-grain size (e.g. ash-like pyrite), poor idiomorphic texture, intense isomorphic substitution causing apparent deviation of lattice parameter from theoretical value, etc. For example, stage- I1 galena has = 5.9391 rt 0.0004 x 10-l' m, higher than the theoretical a0 = 5.9360 rt 0.0005 x IO-" m, while stage-I galena has a0 = 5.9349 f 0.0004 x 10-l' m, lower than theoretical one. Usually, rapid crystallization was caused by the rapid change in nature of the fluid. Then, what is the factor that leads to a drastic change of fluid? Fluid hybridization and boiling are the two main mechanisms for rapid change in fluid nature. The

five of them < 203 "C and two of them > 203 "C, showing superimposition of stage 111 fluidization onto the stage I1 quartz. The stage 111 inclusions distribute linearly, and regarded as secondary ones.

3.2 Composition oxfluid inclusion Gaseous and aqueous compositions (Table 1) are anal yzed with gas-liquid chromatography for fluid inclusions extracted from 6 quartz samples and 4 carbonate samples by decrepitation at temperatures of 120 500 "C. The major cations are K' and Na' with minor Ca2' and Mg2+, while major anions are C1- and SO*: and minor F- The ratios of KNlMC=(K+Na)l(Mg+Ca) and K+/Na+ are high in all the samples (Table 1). In addition, their values of F-ICl- range from 0.019 to 0.052, far higher than those of typical mantle fluid or crustal abundances. This result may indicate that the fluid source is shallow, mainly composed of sediments and continental crustal materials. The stage- I fluid has higher ionic concentration than that of stage 111, showing that it is more capable of mineralizing; while the stage-I1 fluid has the

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690

H2O/C02 values of the stage-I1 fluid are obviously higher than those of stage-I and stage-111, indicating that a large amount of CO2 fled away, which means that the ore fluid boiled in stage 11. The loss of CO:! should decrease the values of HC03-, C032-and Eh, and increase the p H and ion concentration, breaking up metal-bearing complexes. Therefore, a large amount of sulfides were deposited, and Ag+ and Au+ were reduced into native gold and silver. The above considerations are in good agreement with the stage I1 mineralization characteristics, the lowest value of C02/CH4 and the sharp increase in contents of Na+, K+, CM+, Cl-, SO,:' etc. in the stage-I1 fluid. If boiling leads to intense mineralization and fast change in fluid nature, the contents of Ca2+and Mg2+ in stage-I1 fluid should be the highest among the three stages. However, the stage-I1 fluid has the lowest content of Ca2' and Mg2+,and much higher KN/MC ratios than stage-I. We have to consider the input of shallow-derived low salinity fluid. Coincidently, N:! is found occasionally in samples of stage I and 11, but is common in samples of stage 111, showing that N2 content in fluid increased gradually, and the fluid became more and more open and oxidizing. The high N2 content in stage-I11 fluid might indicate the input of meteoric water because N2 could be derived from atmosphere. Additionally, the abundant Ca2' and Mg" (too high to measure) and low H20/C02 ratio suggest the feedback of atmospheric CO2 with meteoric water, resulting in a characteristic carbonatization of stage 111. In summary, the fluid boiling and mixing made the stage-I1 play a decisive role in the formation of the Tieluping Ag deposit.

A larger difference exists between the fluid of stage I and the inclusion fluid of quartzes from migmatite and pegmatite within the Early Precambrian metamorphic basement. 6D of stage-I fluid (-84 -96 %o) is far lower than the latter (6D= -24.5 -27.6 %o); while 8'0 (ca. 7.90 9.78 %o) is much higher than the latter (d80= 5.8 - 6.5 %o). Hence the fluid of stage I could not come from the metamorphic basement or the host-rock. With 8'0 ranging -1.58 -0.18 %o, averaging -0.92 %o, the stage-111 fluid could be dominantly composed of meteoric water. In fact, in the 8'0 6D plot, it lies near the meteoric water line. The stage I1 fluid, with 8'0 = 1.79 %o, just between the stage I and the stage 111, was likely the mixture of meteoric and metamorphic fluids.

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4.2 Carbon isotope composition of orefluid The fluids of various stages of the Tieluping deposit have different d3CCo2values, indicating significant discrepancy in fluid nature and source. In stage I, 8 3 C ~ ranges ~2 from 0.3 %O to 4.0 %o, obviously higher than those of igneous material, organic matter and atmospheric C02. Hence the fluid should derive from marine carbonate formations. d3Cc02 in stage 111 fluid ranges -0.3 -2.0 %o, clearly lower than that of the stage-I, implying a contribution of atmospheric C02. Only one Lf3CCo2value for the sample formed during stages I1 and 111 is about 0.1 %O (TS13). This may suggest a hybridization of meteoric solution and metamorphic fluid, or the transition from the stage I metamorphic fluid to the stage I11 meteoric solution.

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4 H-0-C ISOTOPE FEATURES OF ORE FLUID 4.I Evolution of H - 0 isotope composition

5 DISCUSSION: USE OF THE CPMF MODEL

As shown in Table 1, the variations in 6D and 8'0 of fluids of various metallogenic stages reflect the differences in source and nature of fluids. 8'0, of 9.78 %O with an stage I varies between 7.90 average of 8.94%0,indicating that the fluid was not of meteoric origin, but a deep-seated fluid. In the 8'0 - 6D diagram, samples plot in the domain of metamorphic fluid. Accordingly, the fluid was probably of metamorphic origin. Although the 6D and 8'0 of stage I are close to those of the fluid equilibrated at 600 "C with the quartz in Mesozoic granite in the Xionger terrain (8'0, = 6.6 - 8.5 %o; 6D = - 64.7 - 68.7 %o), its 8'0 is higher and 6D is lower than the latter respectively. Hence the fluid has not originated from Mesozoic magmatism.

As described above, the fluid characters of the Tieluping Ag deposit are similar to those of the orogenic gold deposits (Groves et al. 1998, Kerrich et al. 2000). The deposit was formed in the Mesozoic collision between the North China and South China continental plates. According to the CPMF model (Chen 1990, 1998), in the early stage of the collision, a series of imbricate intracontinental subductions occurred within the orogens. In the process of intracontinental subduction along the Machaoying fault, the down-going slab was subjected reworking, metamorphism and partial melting; and the hydrothermal deposits, batholithic granites and porphyries zonationally appeared in sequence in the overlying slabs. Therefore, the stage-I ore fluid of the Tieluping deposit has the

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characters of deep-seated metamorphic fluid. In the middle stage, the tectonic setting changed from compression to extension, and the geothermal gradient increased continuously. The deep-seated continental crust would partially melt in large scale because of decompression- pyrogenation (heating). Then a large amount of fluid and granitic magma was formed. The uprising of fluid and magma would lead to mineralization and provided heat energy for shallow fluid circulation. Shallow structures such as fractures were expanded and provided suitable channels for fluid circulation. Hence, in stage 11, the shallow-derived meteoric fluidization was most intense. Mixing of the deep seated and shallow derived fluids was responsible of rapid mineral precipitation, leading to polymetallic sulfide assemblage that contributed most to metallogenesis of the Tieluping deposit. In the extensional coolingexhumation stage, the deep-seated fluid decreased, only the shallow meteoric fluid circulated (< 200 "C in general) and caused the stage-I11carbonation. The water-rock interaction of the Tieluping Ag ore includes three stages with mineral assemblages of quartz-pyrite, polymetallic sulfides and carbonates respectively. The stage-I1 minerals captured most fluid inclusions with highest values of H20/C02 and ionic concentration, indicating crystallization was rapid in the multiplex process of fluid mixing and boiling under decompression setting. In general, it can be said that the mineralization and fluidization of the Tieluping Ag deposit are in agreement with the CPMF model. D-0-C isotope-systematic study confirms that the stage-I1 fluid was indeed a hybrid of shallowderived and deep-seated fluids, and the stage-I and stage-I11 fluids came from the reworkingmetamorphism of carbonate-bearing strata and meteoric water respectively, in agreement with the CPMF model. The D-0 isotope system of the stage-I ore fluid is close to that of Mesozoic magmatic fluid; but its 8'0 is apparently higher than the latter; its 6D is lower than the latter; and its 8 3 C ~ is~ very 2 high. All these features indicate the ore fluid could not derive from differentiation of the granitic magma. The d3Cc0z of stage-I fluid, ranging 0.3 4.0 %o, strongly excludes the possibility that the ore fluid came from the main rocks in the Xionger terrain, and suggests that it should derive only from metamorphism of carbonate strata. It is known that a large volume of CSAF of Guandaokou and Luanchuan Groups occurs south to the Machaoying fault, and abundant fragments of carbonate strata from the Guandaokou and Luanchuan Groups have been identified within the fault belt. Accordingly, it can be supposed that the slab made up of the

Figure 1. The CPMF model (from Chen 1990, 1998).

Guandaokou and Luanchan Groups underthrusted along the Machaoying fault northward into the Xionger terrain. Reworking, metamorphism and partial melting of the subducting slab produced the stage-I fluid of the Tieluping silver deposit, and induced large scale mineralization and accumulation of ore elements in the Xionger terrain. Hence the fault-bound AdAg deposit zone, the batholithic granite zone and the porphyry and porphyryassociated Au-MO deposit zone occurred in sequence north of the Machaoying fault (Fig. 1). REFERENCES Chen, Y.J. 1990. Gold Mineralization in West Henan. Ph.D Thesis, Nanjing University Chen, Y.J. 1998. Fluidization model for continental collision in special reference to study ore-forming fluid of gold deposits in the eastern Qinling Mountains, China. Progress in Natural Science, 8: 385-393. Kerrich, R., Goldfarb, R.J. & D.I. Groves 2000. The characteristics, origins and geodynamic settings of supergiant gold metallogenic provinces. Science in China Series D, 43(Sup.): 1-68. Zhang, L.G. 1985. The Application of the Stable Isotope to Geology (in Chinese). Xian: Shaanxi Science & Technology Publishing House, 267.

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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Water-rock interaction in genesis of perlite at Monte Arci volcanic complex (West Sardinia, Italy) R .Cioni, G .Macciotta & M .Marchi Dipartimento di Scienze della Terra, Universitd di Cagliari, Italy

G .Padalino & R.Simeone Dipartimento di Geoingegneria e Tecnologie Ambientali, Universita di Cagliari, Italy

M .Palomba Centru Studi Geominerari e Mirzeralurgici, CNR, Cagliari, Italy

ABSTRACT: At Monte Arci (West Sardinia) Plio-Pleistocene rhyolitic lavas with hyaline texture outcrop. A detailed investigation has been conducted to define the geochemical characteristics of the nonhydrated (obsidian) and hydrated (perlite) volcanic glasses recognized to date. XRF analysis was used to determine major and trace elements in whole rock, and glass inhomogeneity was detected using EPMA methodology. 6D and 6"O isotopes and FTIR analyses were also performed. Perlite may have originated from later hydration by meteoric water giving rise mainly to Na20 removal. A negative correlation between Na2O and H20, lower 6D and 6I80values and higher OH-/H20 ratios in obsidian seem to support this hypothesis. 1 INTRODUCTION

2.2 The Monte Arci volcano

The varying water content of the rhyolitic glassy rocks outcropping at Monte Arci (from less than 1% for obsidian to more than 1-2% for perlite) may be the result of at least three different processes: - secondary hydration of obsidian by meteoric water percolating through each cooling unit of glassy lava flows; - varying degassing of rhyolitic lavas; - post-emplacement late-stage hydrothermal circulation within lava flows.

The Monte Arci volcanic complex is located in central-western Sardinia, east of the Campidano N-S trending and westwards dipping rift. it includes a large number of lava flows and minor pyroclastics. The basement is composed of Miocene sedimentary rocks with associated calcalkaline submarine lavas. For the most part the Pliocene complex (about 150 Km2) is composed of a subalkaline sequence evolving from subalkaline basalt to rhyolite. A small amount of transitional basalt also occurs, related to low silica-oversaturated alkali-trachyte. Alkaline basalt is also found sporadically. General Pliocene stratigraphy (Assorgia et al. 1976; Beccaluva et al. 1974) can be summarized as follows, from bottom to top: (1) rhyolitic lava flows, (2) dacitic and alkali-trachytic lava flows and (3) basaltic and andesitic lava flows.

2 GEOLOGICAL BACKGROUND 2.1 Regional setting In the time span from Early Pliocene to Upper Pleistocene, within-plate volcanic activity took place on the island of Sardinia, spreading chiefly basic lavas of alkaline to subalkaline affinity. Minor amounts of intermediate and differentiated products also occurred. The first volcanic products of this period (Beccaluva et al. 1985, and references therein) located at Cap0 Ferrato (5 Ma), later extending to other areas of the island (Monte Arci 3.8-2.8 Ma, Montiferro 3.9-1.6 Ma, Orosei-Dorgali 2.9-2.0 Ma, central plateaux 3.7-1.4 Ma, southern plateaux 3.8-1.7 Ma). The most recent activity, aged 0.9 to 7: llacite lava, 8: Black orebody, 9: High-grade siliceous orebody, 10: Iaw-gradd3 siliceous orebod), I I : G) psiiiii orebociy

ceous orebody is funnel-shaped, the Au-bearing bedded siliceous orebody is dish-shaped, and the massive base-metal orebody is thinly lenticular in shape (Fig. 3; Yamada et al. 1988).

Au-bearing stockwork and bedded orebodies consist of Au-bearing and Pb-Zn-bearing siliceous ores. Massive base-metal orebodies consist of compact and brecciated black ores. Au-bearing siliceous ores are cut by a stockwork of Pb-Zn-bearing siliceous ores. The Au mineralization is earlier than the PbZn mineralization associated with the formation of Pb-Zn-bearing siliceous ores and compact and brecciated black ores. Au-bearing siliceous ores consist of major amounts of pyrite, chalcopyrite and quartz, and lesser amounts of electrum, sphalerite, galena, hematite and kaolin and sericite. Pb-Zn-bearing siliceous ores include major amounts of quartz, sphalerite and galena, and small amounts of chalcopyrite, pyrite and sericite. The main constituent minerals of compact and brecciated black ores are sphalerite, galena, pyrite, chalcopyrite and barite, with lesser amounts of tetrahedrite, pearceite, pyrargyrite and sericite and rarely electrum and bornite. In the black orebody composed of brecciated black ores, there is alternation of brecciated black ores, a layer rich in barite and a layer composed of mixture of kaolin and sericitehmectite mixed layer minerals. The clay-bearing layer shows a banded texture, thus texture suggests that the clay minerals were precipitated directly from a hydrothermal solution. Based on the fact that the mineral assemblage of kaolin and sericite is in altered rocks of Au-bearing siliceous ores and clay layers in brecciated black ores, there is a possibility that the pH of hydrothermal solution forming Au-bearing siliceous ores and black ores is more acidic than the pH of hydrothermal solution forming typical Kuroko deposits.

3 MODE OF OCCURRENCE OF ORES The ores of No. 5 orebody are divided into four types: Au-bearing siliceous ore, Pb-Zn-bearing siliceous ore, massive black ore and brecciated black ore.

Figure 3. Cross section of the No. 5 orebody of the Nurukuwa deposit (Yainada 1988). 1: Tuff breccia (Tobe I:.), 2: Pumice tuff & inudstone (U. Hayasemori F.), 3: Dacite lava, 4: brecciated and compact bedded black ores. 5: Au-bearing bedded siliceous ore, 6: Au-bearing network siliceous ore, 7: fault

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Eigiire 1.Salinities of versus homogenization temperatures of fluid inclusions from No. 5 orebody.

4 FLUID INCLUSION STUDIES Homogenization temperatures and salinities of fluid inclusions in quartz of Au-bearing siliceous ores and in sphalerite and barite of brecciated black ores were found to be 253 to 286°C and 3.3 to 5.3 wt % NaCl eq. for quartz, 210 to 252°C and 2.7 to 4.1 wt % NaCl eq. for sphalerite, and 145 to 262°C and 1.5 to 2.7 wt %, NaCl eq. for barite, respectively (Fig. 4). Homogenization temperatures and salinities of fluid inclusions in the Au-bearing siliceous ores are higher than those of fluid inclusions in the brecciated black ores. The salinity of fluid inclusions in the brecciated black ores is similar to the salinity of seawater.

Figure 5. Histogram of 6D values of clay minerals from No. 5 ore body. Data of 61) values of clay minerals froin Kurokotype &posits, vein-type &posits and modem gcothermal area are from Matsubaya and Marumo (1986).

ated black ores were estimated on the basis of the hydrogen isotopic ratios of sericite, the formation temperature, and the fractionation factor of sericite-water by Marumo et al. (1980). The estimated hydrogen isotopic ratios range from - 14 to - 10 per mil. Hydrogen isotopic ratios of kaolin of Au-bearing siliceous ores range from -62 to -53 per mil. The hydrogen isotopic ratios of kaolin are about 3 0 per mil lower than those of kaolin that is associated with typical Kuroko deposits in Japan (Fig. 5). The hydrogen isotopic ratios are also different from those of kaolinite equilibrated with meteoric water from veintype deposits and modern geothermal area in Japan. Hydrogen isotopic ratios for hydrothermal solution

5 STABLE ISOTOPE STUDIES The oxygen isotopic ratios of quartz in Au-bearing siliceous ores and Pb-Zn-bearing siliceous ores range from +9.2 to +10.2 and +9.0 to +10.0 per mil, respectively. The oxygen isotopic ratio of quartz crystals in druse of brecciated black ores is + 10.5 per rnil (Yamada et al. 1988). There is no significant difference among these oxygen isotopic ratios. The oxygen isotopic ratios of fluid responsible for the formation of Au-bearing siliceous and black ores were estimated on the basis of oxygen isotopic ratios of quartz presented above, the formation temperatures and fractionation factor of quartz-water by Matsuhisa et al. (1979). The ranges of the calculated oxygen isotopic ratios of hydrothermal solution responsible for Au-bearing siliceous ores and brecciated black ores are +0.4 to +2.8 per mil and -0.5 to + 1.7 per mil, respectively. The hydrogen isotopic ratio of sericite in brecciated black ores ranges from -37 to -41 per mil. The hydrogen isotopic ratio is very close to the hydrogen isotopic ratio of sericite associated with other Kuroko deposits in Japan (Fig. 5). The hydrogen isotopic ratios for hydrothermal solution responsible for brecci-

Figure 6. The hydrogen and oxygen isotopic ratios ofore fluid responsible fbr the fbmiation of Au-bearing siliceous and black ores gom No. 5 ore body. Those ofseawater, magmatic fluid (Taylor 1979) and meteoric water are also shown in the diagram. Water/Rock ratios offfuids of the Au-bearing siliceous ores (thin solid lines) and the black ores (broken lines) were calculated using hctionation &tors of kaolinite-water (Sheppard & Gilg 1996), sericite-water (Marumo et al. 1980) and quartz-water (Matsuhisa et al. 1979).

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responsible for Au-bearing siliceous ores were also calculated using the hydrogen isotopic ratios of kaolin, the formation temperatures, and the fractionation factor of kaolinite-water by Sheppard and Gilg (1996). The calculated hydrogen isotopic ratios range from -47 to -33 per mil. The hydrogen isotopic ratios of hy drothermal solution directly extracted from fluid inclusions in quartz for Au-bearing siliceous ores range from -55 to -45 per mil. The hydrogen isotopic ratios of hydrothermal solution forming the Au-bearing siliceous ores is distinctly smaller than those of hydrothermal solution forming the brecciated black ores (Fig. 6), although there is a difference of about 10 per mil between the calculated hydrogen isotopic ratios of fluid and those measured in fluid extracted from fluid inclusions in quartz. 6 FORMATION ENVIRONMENTS OF THE NURUKAWA DEPOSIT Considering the higher salinity of fluid inclusions of Au-bearing siliceous ores (Fig. 4) and the relationships of hydrogen and oxygen isotopic ratios of hydrothermal solution (Fig. 6), it is thought that the Au-bearing siliceous ores at No. 5 orebody of the Nurukawa deposit were formed by hydrothermal solution containing fluid of magmatic origin. The fluid responsible for the formation of black ores, on the other hand, is dominantly of seawater origin. Based on the geological relationships around the Nurukawa deposit (Fig. 2), there is a possibility that the generation of fluid of magmatic origin was caused by the emplacement of dacitic cry pto-lava-domes under the Nurukawa deposit. When the Au-bearing siliceous ores were formed, the contribution of fluid of magmatic origin to hydrothermal solution of seawater origin would have been large. In the case of the formation of black ores, the hydrothermal solution is thought to be mainly a solution of seawater origin, and the fluid of magmatic origin would have occasionally mixed with the hy drothermal solution. The style of circulation of hydrothermal solution changes from the hydrothermal system associated with great contribution of magmatic water to the seawater dominant hy drothermal system according to the decline in activity of dacitic magmatism.

teranatinal Syniposiuni on Water-Rock Interaction Extended Abstract: 382-385. Matsuhisa, Y., Goldsmith, J. R. & R. N. Clayton 1979. Oxygen isotopic fractionation in the system quartz-albiteanorthite-water. Geochini. el Cosnzochim. Acta 43: I I3 11140. Metal Mining Agency of Japan (MMAJ) 1988. Report of detailed geological survey. Hokuroku-Kita (north) area 1987 fiscal year: 102p. Nishitani, Y., Tanimura, S., Konishi, N., Yamah, R. & M. Sato 1986. Exploration for the Nurukawa deposits - A summary of prospecting to the discovery of ores, the geology and the ore deposits-. Mining Geology 36: 149- 161. Sheppard, S. M. F. & H. A. Gilg 1996. Stable isotope geochemistry of clay minerals. Clay Minerals 3 1 : 1-24. ‘l’animura, S., Date, J., Takahashi, T. & H. Ohmoto 1983. Geologic setting of the Kuroko deposits, Japan. Part 11. Stratigraphy and structure of the Hokuroku district. Econ. Geol. Monograph 5: 24-38. Taylor, H. P. Jr. 1979. Oxygen and hydrogen isotope relationships in hydrothermal mineral deposits. In H. L. Banes (ed), Geocheniistrj of’ Hydothermal Ore Deposits, 2nd 236-277. New York: John Wiley & Sons Inc. Yamada, R. 1988. Geology of the Au-Ag-Rich Kuroko Deposit at Nurukawa, Aomon Prefecture. Society of’ Mining Geologists of’Japan, Guidebook (Kuroko Deposits and Geofhemial Fields in Northern Honsliu) 3: 26-38. Yamada, R., Nishitani, Y., Tanimura, S . & N. Konishi 1988. Recent development and geologic characteristics of the Nurukawa Kuroko deposit. Mining Geology 38: 309-322.

REFERENCES Marumo, K., Nagasawa, K. & Y. Kuroh 1980. Minedogy and hydrogen isotope geochemistry of clay minerals i n the Ohnuma geothermal area, northeastern Japan. Emh Planet. Sci. Let. 47: 255-262. Matsubaya, 0. & K. Marumo 1986. Hydrogen isotopic evidence for conservation of water in clay minerals. Fifth f n -

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2

Sea water-basalt interaction in the Kerguelen Plateau, Indian Ocean V.B .Kurnosov, B .P.Zolotarev & A .V.Artamonov Geological Institute, Moscow, Russia

ABSTRACT: All studied basalts from Hole 747C (ODP Leg 120) and Holes 1136A, 1137A, 1138A, and 1140A (ODP Leg 183) of the Kerguelen Plateau are altered to various degree in low-temperature hydrothermal environment. Temperature of alteration reached 150°C and probably higher at alteration of basalts from Hole 1138A. Obtained results on the mobility of chemical elements due to alteration of basalts from the Kerguelen Plateau at sea water-rock interaction revealed that in ((non-oxidative))environment of alteration loss of most elements (Si, Fe, Mn, Ca, K, Cu, Zr, Y, Sc, Nb, Rb, BayThySr, and REE) prevails. On the contrary, in the oxidative environment basalts accumulate K, Fe, P, Sr, Bayand REE. Alteration of effusive basalts is related mainly to horizontal fluid flows within highly permeable contact zones between lava flows.

1 INTRODUCTION Basalts processes from aseismic structures, as well as their alteration, are poorly studied compared to those from mid-ocean ridges. Aseismic ridges and plateaus represent giant structures of oceanic floor. These structures demonstrate characteristic features of magmatic rock compositions, environments of effusion (from shallow-water and subaerial to deepwater environments), thermal history, and circulation of both sea water and fluids (Kurnosov et al. 1995, Kurnosov & Murdmaa 1996). Combination of these peculiarities for aseismic ridges and plateaus is somewhat specific in comparison to mid-ocean ridges. So, alteration of basalts at water-rock interaction in these structures have its own characteristic features which need further investigation. The Kerguelen Plateau is a typical large igneous province. We studied altered basalts recovered during ODP Leg 120 (Hole 747C) and ODP Leg 183 (Holes 1136A, 1137A, 1138A, and 1140A) on the Kerguelen Plateau (Fig. 1). The Kerguelen Plateau is located in the south-eastern sector of the Indian Ocean. Site 1136 is located in the southern Kerguelen Plateau. Hole 1136A penetreted 161.4 meters below sea floor (mbsf) - 128.1 m of sediment overlies 33.3 m basalts. Lower 48.19 m of sedimentary section is Cenomanian to early Albian in age. Basalts (three inflated pahoehoe flows) were recovered below 128.8 to 161.4 mbsf These basalts probably reflect subaerial eruption (Coffin et al. 2000). Site 1137 lies on the eastern portion of the crest of 721

Elan Bank in the western margin of the Kerguelen Plateau. Hole 1137A was rotary cored to a depth of 371.2 mbsf. The basalts were recovered from the lower 150 m (from 219.5 to 371.2 mbsf). Basalts (seven basaltic lava flows) intercalate with three sedimentary and volcaniclastic layers. The flows are 7 to 27 m thick. Six basaltic flows erupted subaerial,

Figure 1. Location map of the Kerguelen Plateau showing ODP sites.

one erupted probably under shallow environment (Coffin et al. 2000). Sites 747 and 1138 are located in the central part of the Kerguelen Plateau. The cored basement at Hole 747C consists of approximately 12 lava flows separated by brecciated basalts. The total of 53.9 m of basement were penetrated. Hole 1138A penetrated 842.7 mbsf approximately 144 m of basalt basement and 698 m of overlying sedimentary section are Pleistocene to Upper Cretaceous in age. Igneous basement was divided in 22 units. Upper two units were interpreted as subaerial pyroclastic flow deposits. Other units represent subaerial basaltic inflated pahoehoe to classic aa lavas (Coffin et al. 2000). Site 1140 lies in the northernmost part of the Kerguelen Plateau. Hole 1140A penetrated 32 1.9 mbsf (including about 87 m of deep marine pillow basalts). The latter are middle Miocene in age. From south to north the age of recovered basalts decreases from approximately 105 Ma to 35 Ma (Coffin et al. 2000). 2 BASALT GEOCHEMISTRY Basalts from various parts of the Kerguelen Plateau are derived from tholeiitic melts. Compared to MORB, they are enriched in Ti, P, Zr, Sr, and largeion REE. Along the Kerguelen Plateau basalts formed from compositionally similar initial melts, which generated at approximately similar depths but greater than parental melts of basalts in mid-ocean ridges. Only in the northern part of the Kerguelen Plateau (Hole 1140A) basalts formed probably under the influence of two sources located at different depths. Variable concentrations of Ti mark this difference. We recognized both shallow, low-Ti basalts similar to basalts from mid-ocean ridges, and highto moderate-Ti basalts which are considered to derive from initial melts generated deeper. 3 ALTERATION In basalts we examined secondary minerals in groundmass, vesicles, and veins. Also we studied mobility of chemical elements in relation of alteration of basalts under both oxidative and ((nonoxidative)) environments (Bass 1976, Bass et al. 1973). 3.1 Minerulogy Thin sections show that the basalts froin Hole 1136A are weakly to moderately altered. According to wholerock chemical analyses, altered basalts contain 0.90 - 1.58 wt% of H20'. The degree of rock

oxidation is feeble. Both olivine and interstitial glass are completely replaced by smectite-chlorite aggregate. Plagioclase is partly replaced by smectitechlorite aggregate. Vesicles are filled mainly with smectite. Smectites contain significant concentration of Fe. Veins are filled with calcite. Secondary minerals identified in basalts from Hole 1136A indicate low temperature environment at water-rock interaction. Analysis of thin sections reveals that basalts from Hole 1137A are more altered than basalts from Hole 1136A. We estimate the degree of basalts alteration from 10-15% to 60%. According to wholerock chemical analyses, the altered basalts contain 0.19 3.03 wt% of H20+.The ratio Fe203/Fe0varies from 0.88 to 3.57, and also indicate various degree of basalt oxidation. Tuff is strongly altered (80% - estimation based on the thin section analysis) and strongly oxidized (Fe203/Fe0= 10.14). Thin sections show that alteration of phenocrysts and groundmass causes femic minerals to be completely replaced with chlorite and chlorite-smectite aggregates. Plagioclase phenocrysts in porphyritic basalt are replaced with K-feldspar (adular?), smectite-chlorite aggregates, and carbonate. Plagioclase in the matrix is also replaced with IS-feldspar. The degree of alteration of rock is high - about 40% in thin section. Highly altered tuff contain pseudomorphs of chlorite and carbonate upon prismatic mineral, biotite is chloritized. Smectite-chlorite aggregates replace interstitial glass. In the upper part of basalt section, smectite is the most common mineral filling vesicles. Often, it occurs together with some admixture of clinoptilolite and heulandite. Besides, we have observed an admixture of 7A-mineral (probably dickite) and calcite. In the lower part of basalt section besides smectite, chlorite, mixed-layer smectite-chlorite minerals are present in vesicles. Chlorite - a predominant mineral - is present of lower depth. Clinoptilolite, hydromica, and quartz are recognized as admixture to smectite and chlorite minerals. The appearance of chlorite phases is probably related to the location of basalts at significant depth that may have caused their alteration under high temperature. Veinlets in basalts contain smectite, calcite and clinoptilolite and quartz in trace amounts. The entire complex of secondary minerals revealed in basalts from Hole 1137A indicates low temperature conditions of their alteration. Also, the appearance of chlorite minerals downward in the hole may have been caused by increased temperature. Hole 1137A is located probably close to an ancient center of eruption. The presence of dickite in trace amounts in the upper part of basalt section provide evidence of aerial weathering. The degree of basalts alteration from Hole 1138A according to study of thin sections varies from slight

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Zeolites were not found. This is the principal difference in alteration of basalts froin Hole 1140A (sub-marine environment of basalt effusion) in coinparison with alteration of basalts from Holes I 136A, 1137A, and especially Hole 1138A (subaerial environment of lavas effusion). Thus, basalts of pillow lavas from Hole 1140A show low degree of alteration and mostly are not oxidized. Secondary minerals indicate a lowtemperature environment. Circulation of hotter fluids probably occurred along cracks filled now with smectite, defective chlorite, serpentine(?) and quartz

to intensive (from 10% to 50%). According to chemical analyses, the altered basalts contain 0.36 4.96 wt % of H20+. Tuff is strongly oxidized (Fe203/Fe0ratio is 25.64) and altered (about 60% in thin section). Thin sections show that olivine is completely replaced with light green chlorite. Plagioclase is partly replaced by chlorite. The central part of plagioclase is replaced by micaceous mineral. Interstitial glass is completely replaced with chlorite and smectitechlorite aggregates. The study of vesicles filled with secondary minerals in basalt section of Hole 1138A revealed typical characteristics among other holes of Leg 183. Of particular note feature is the abundance of zeolites: heulandite, clinoptilolite, mordenite, stilbite, analcime, and natrolite. Thomsonite is present occasionally. No vertical zonation in zeolite distribution in Hole 1138A was observed. Also, clay minerals in basalt section of Hole 1138A do not show vertical zonation in their distribution. Smectite occurs in all parts of basalt section. Veins are filled mainly with smectite and zeolite. XRD show that among zeolites, heulandite, transitional heulandite-clinoptilolite, stilbite, analcime, and inordenite are present. All secondary minerals (clay and non-clay minerals) studied in vesicles and veins, as well as in basalt groundmass, do not show vertical zonation in their distribution throughout the basalt section of Hole 1 138A. Hole 747C is located near Hole 1138A. The studied basalts froin Hole 747C show various degree of alteration (H20+varies from 0.69 to 5.14 wt%). In altered basalts dominant secondary minerals are represented by smectite, or sinectites with chlorite or swelling chlorite, and chlorite. Zeolites were determined in amygdules, in the groundinass, and altered plagioclase phenocrysts. Zeolites are represented by chabazite, natrolite, thomsonite, mesolite, stilbite, and heulandite (Sevigny et al. 1992). The temperature of formation of similar, present-day zeolite from Iceland is about 120'C (Kristmannsdottir & Tomasson 1978). Calcite occurs in vesicles, veins and in the groundmass. Quartz occurs in small amounts in some vesicles. Two types of alteration - oxidative and nonoxidative ones - are recognized in Hole 747C. Oxidative zones are marked by goethite, Fe-hydroxides, calcite, and celadonite(?). This secondary mineral assemblage forms at low-temperature water-basalt interaction. Basalts of pillow lavas from Hole 1140A show alteration weaker than basalts from other holes of Leg 183. Thin sections reveal alteration of basalts from 5-10% to 15-20%. Basalts are fresh or poorly oxidized. The ratio Fe203/Fe0 is low and varies from 0.53 to 1.07.

3.2 Mobility of chemical elements In bulk samples of basalts we analyzed major elements by ((classical)) wet-chemistry. To determine minor elements we used both atomic emission spectroscopy and XRF. REE was analyzed by neutron activation analysis. Rocks from Hole 1I36A consist of non-oxidized or, more freguently, slightly oxidized basalts and oxidized basalts. Weakly altered ((non-oxidized))basalt (H20+1.25 wt %) reveals that alteration leads to a decrease in Si, Fe, Mn, Ca, Cu, Zn, and Th, while the concentration of most REE slightly increase. A marked increase in Rb, Sr, and Ba has been observed and occurs even under slightly oxidative environment. It is probable that slightly oxidative processes affect gaidloss of elements. The increase in oxidation degree of basalt (with Fe203/Fe0 ratio of 2.07), contrary to gaidloss of elements in ((non-oxidized))basalts, leads to a decrease in P, REE, Nb, and Zr. Basalts from Hole 1140A are similar in the degree of alteration. In spite of their low degree of alteration, these samples can be used in the study of chemical element mobility in ((pure)) non-oxidized environment. A decreasing trend under nonoxidative alteration in basalt (H20+ 1.40 wt %, Fe203/Fe0 0.69) is evident for elements such as REE, Cu, Zr,Y, Rb, and Sr. Other elements (Ni, V, CO, Zn, and Ba) show weak increasing trend. The oxidized basalt (H20+ 1.65 wt %, Fe203/Fe0 2.41) revealed that many elements follow another trend than that identified for non-oxidative environment of basalt alteration. Under oxidative type of alteration, basalt mostly accumulates all minor elements, including REE. Petrogenetic elements (Si, total Fe, Mn, Ca, and Na) demonstrate a weak decreasing trend. On the contrary, K and P accumulate in basalt, while AI and Mg mostly remain constant. Under non-oxidative type of basalt alteration (HzO+ 2.03 wt %, Fe203/Fe0 2.15) from Hole 1137A, the dominant trend is a loss of majority of elements, except Mn, K, Rb, and Ba. Oxidative type of alteration mostly leads to accumulation in basalts 723

of many elements and primarily (Hole 747C) minor elements, including REE. This is especially evident for Fe, Sr, and Ba. Highly altered basalts from Hole 747C contain less Sc, V, Cu, Y, Nb, Sr, and Ba but more Rb. Schlich et al. (1989) have proposed that Nay K, Bay and Rb were mobile and redistributed during the waterbasalt interaction.

Iceland (Walker 1960; Tomasson & Kristmannsdottir 1972; Kristmannsdottir & Tomasson 1978). Results obtained for the mobility of chemical elements during alteration of basalts from the Kerguelen Plateau revealed that in non-oxidative environment of alteration loss of most elements prevails. On the contrary, in oxidative environment basalts accumulate elements.

4 CONCLUSIONS

ACKNOWLEDGMENTS

All studied samples from Holes 1136A, 1137A, 1138A (subaerial and aerial effusions), and 1140A (deep marine environment), as well as tuffs cored in various parts of the Kerguelen Plateau, are at various degree altered in low-temperature hydrothermal environment. Also, dickite(?) in trace amounts in the upper part of basalt section of Hole 1137A is indicative of aerial weathering. Basalts of pillow lavas recovered in Hole 1140A in north part of the Kerguelen Plateau show minimal alteration, as they are present in the lowest temperature environment. Smectite prevails among secondary minerals, while vesicles and veins contain no zeolites, contrary to other holes of Leg 183. It is probable that basalts of Hole 1138A are located in the vicinity of the paleo center of eruption. A wide distribution of a broad spectrum of zeolites, together with significant variations in chemical composition of each of them, result from rock alteration in hydrothermal environment. Temperature of alteration was about 150°C and probably higher. We have observed an absence of vertical zonation throughout the basalt section cored by Hole 1138A. No zonation in distribution was documented for all secondary minerals (clay minerals and zeolites) from vesicles and veins, as well as in groundmass of basalts. Also, band-like distribution of analcime, stilbite, and mordenite, as well as indications of zonality within basalt flows, provide evidence that alteration of effusive basalts is related mostly to mineral and chemical peculiarities of basalt flows, and to horizontal fluid flow along contacts of lava flows. This picture characterizes the alteration environment of basalts recovered from Holes 1136A and 1140A. It holds also for the ccchloritic)) lower part of basalt section of Hole 1137A. Zonation in distribution of secondary minerals within lava flows, and the absence of vertical zonation in basalt section, were observed during a previous study of altered basalts from Suiko Guyot in the Emperor Seamount Chain, Hole 4336, Leg 55 (Kurnosov 1986) and from Allison Guyot in the West Pacific Guyots, Hole 865A, Legs 143 and 144 (Kurnosov et al. 1995). The influence of heated waters of interlayer-fissure circulation on the formation of subhorizontal zeolite zones in basalts was studied in

We thank ODP for providing the samples. The study was supported by Russian Foundation for Fundamental Research: grants 98-05-64856 and 99-05-65462.

REFERENCES Bass, M.N. 1976. Secondary minerals in oceanic basalt, with special reference to Leg 34, Deep Sea Drilling Project. In Yeats, R.S., Hart, S.R., et al., Init. Repts. DSDP, 34: 393432, Washington (US Govt. Printing Office). Bass, M.N., Moberly, R., Rhodes, J.M., Shih, C.S. & S.E. Churh 1973. Volcanic rocks cored in the central Pacific, Leg 17, Deep Sea Drilling Project. In Winterer, E.L., Ewing, J.I., et al., Init. Repts. DSDP, 17: 429-503, Washington (US Govt. Printing Office). Coffin, M.F., Frey, F.A., Wallace, P.J., et al. 2000. Proc. ODP, hit. Repts., 183 [CD-ROM]. Available from: Ocean Drilling Program, Texas A&M University, College Station, TX 77845-9547, USA. Kristmannsdottir, H. & J. Tomasson 1978. Zeolite zones in geothermal areas in Iceland. Natural zeolites, Occurrence, Properties, Use. Oxford and New York: Pergamon Press, 277-284. Kurnosov, V.B. 1986. HydrothernialAlteration of Basalt in the Pactfic Ocean and Metal-bearing Deposits, Using Data of Deep-sea Drilling: Moscow (Nauka). Kurnosov, V.B., Zolotarev, B.P., Eroshchev-Shak, V.A., Artamonov, A.V., Kashinzev, G. & 1.0. Murdmaa 1995. Alteration of basalts from the West Pacific Guyots, Legs 143 and 144. In Haggerty, J.A., Premoli Silva, I., Rack, F., and McNutt, M.K. (Eds), Proc. ODP, Sci. Results, 144: 475491, College Station, TX (Ocean Drilling Program). Kurnosov, V.B. & 1.0. Murdmaa 1996. Hydrothermal and cold-water circulation within the intraplate seamounts: effects on rock alteration. The oceanic lithosphere & scientrJic drilling into the 21" Century, Woods Hole, MA, USA, 87-88. Sevigny, J.H., Whitechuech, H., Storey, M. & V.J.M. Salters 1992. Zeolite-facies metamorphism of Central Kerguelen Plateau basalts. In Wise, S.W., J.R., Schlich, R., et al., 1992. Proc. ODP, Sci. Results, 120: 63-69, College Station, TX (Ocean Drilling Program). Schlich, R., Wise, S.W., J.R., et al. 1989. Proc. ODP, Init. Repts., 120: College Station, TX (Ocean Drilling Program). Tomasson, J. & H. Kristmannsdottir 1972. High temperature alteration minerals and geothermal brine, Reykjanes, Iceland. Contr. Mineral. Petrol., v. 36, N 2: 123-134. Walker, G.P.L. 1960. Zeolite zones and dyke distribution in relation to the structure of the basalts in Eastern Iceland. J. Geol. 68: 515-525.

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Magmatic versus hydrothermal processes in the formation of raw ceramic material deposits in southern Tuscany P.Lattanzi Dipartinzento di Scienze della Terra, Universith di Cagliari, Italy

M .Benvenuti, PCostagliola, C .Maineri, I .Mascaro & G.Tanelli Dipartimento di Scienze della Terra, Universitci di Firenze, Italy

A .Dini CNR-IGGI, Ghezzano (PI), Italy

G .Ruggieri CNR-IIRG, Ghezzano (PI), Italy

ABSTRACT: Ceramic raw material in Tuscany is currently mined in four main deposits (Marciana and La Crocetta in Elba Island, Botro ai Marmi near Campiglia Marittima, and Piloni di Torniella near Roccastrada). The deposits are associated with magmatic rocks of the so-called “Tuscan Magmatic Province”, and reflect various degrees of modification of original magmatic features by interaction with hydrothermal fluids. At Marciana, the mined material is an essentially unaltered K- and Na- rich porphyritic aplite; at Botro ai Marmi, a sienogranite intrusion (possibly already K-rich because of phenomena developed at the magmatic stage) was enriched in K through deposition of a second generation of K-feldspar by high-temperature, highly saline fluids, of dominantly magmatic nature; at La Crocetta, a porphyritic aplite similar to that mined at Marciana was transformed into a sericitized, K-rich Na-poor rock by moderate temperature, moderate salinity fluids of magmatic(?)-meteoric origin; at Piloni, rhyolitic lavas were kaolinitized in a shallow epithermal environment, perhaps by steam-heated meteoric waters 1 INTRODUCTION Tuscany has been for millennia a major mining region of Italy, with production of a large variety of commodities (Cipriani & Tanelli 1983, Tanelli 1983, Tanelli & Lattanzi 1986, Lattanzi et al. 1994). Traditionally, the economic, as well as the scientific, interest has been focused on the pyrite and metalliferous deposits, whereas the relevant wealth of industrial minerals and rocks received comparatively little attention. However, with the progressive closure of all metalliferous deposits, industrial minerals and rocks (together with geothermal fluids) are currently the only mineral resources of the region. There is, therefore, a new surge of interest around these deposits, both because of the economic importance, and because of the increasing awareness of maintaining compatibility of mining activities with the environment (e.g., Tanelli 1993). Among industrial minerals of Tuscany, raw materials for the ceramic industry are especially important (ca. 600,000 t/yr, about one third of total Italian production). Currently, the main activity is localized in four deposits (Marciana and La Crocetta in Elba Island, Botro ai Marmi near Campiglia Marittima, and Piloni di Torniella near Roccastrada; Fig. 1). All these occurrences are associated with magmatic rocks Of the so-ca11ed “Tuscan Magmatic Province” (TMP; Marinelli 1959, 1961). In this C0mmunication, we summarize the main features of these four

deposits, and we discuss the importance of magmatic and hydrothermal processes in determining the economic quality of the ore.

Figure 1. Map of the northern Tyrrhenian area, showing outcrops of Neogene magmatic rocks, and the location of the four deposits of this study.

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Table 1. Representative chemical analyses of mined rocks at Marciana, LA Crocetta, and Botro ai Marmi (average values of N analyses) Marciana N= 3 73.23 0.02 16.20 0.18 0.19 0.09 0.65 3.96 4.27 0.02 1.15 99.98

Crocetta N=12

B. Marmi N=3

70.63 0.00 17.39 0.24 0.26 0.28 0.64 7.92 0.13 0.01 2.26 99.75

70.77 0.28 15.48 0.13 0.16 1.73 0.60 7.40 2.36 0.14 0.83 99.89

whereas a later high-angle fault system is associated with carbonate + sulfide veins, and has an adverse effect on ore quality. Fluid inclusion studies suggest that sericitization was associated with a lowtemperature ( 500°C) where F-bearing minerals coexist with F-rich fluids. Alkaline igneous rocks and skarn systems are examples of such environments. REFERENCES Banks, D., Yardley, B., Campbell, A. & K. Jarvis 1994. REE composition of an aqueous magmatic fluid: A fluid inclusion study from the Capitan Pluton, New Mexico, U.S.A. Cliem. Geol. 113: 259-272. Chao, E., Back, J., Minkin, J., Tatsumoto, M., Junwen, W., Conrad, J., McKee, E., Zonglin, H., Quingrun, M. & H. Shengguang 1997. The sedimentary carbonate-hosted giant Bayan Obo REE-Fe-Nb ore deposit of Inner Mongolia, China: A cornerstone example for giant polymetallic ore deposits of hydrothermal origin. USGS Bulletin 2 143. Giere, R. 1989. Hydrothermal mobility of Ti, Zr and REE: examples from the Bergell and Adamello contact aureoles (Italy). Term Nova 2: 60-67. Lipin B. & G. McKey 1989. Geochemistry and mineralogy of rare earth elements. Reviews in Mineralogy 21, Min. Soc. Am. Moine, B., Rakotondratsima, C. & M. Cuney 1985. Les pyroxenites a urano-thorianite du Sud-Est de Madagascar: conditions physico-chimiques de la metasomatose. Bull. Mindral. 108: 325-340. Moine B., Ramambazafy, A., Rakotondrazafy, M., Ravolomiandrinarivo, B., Cuney, M. & P. de Parseval 1998. The role of fluorine-rich fluids in the formation of the thorianite and sapphire deposits of S.E. Madagascar. Mineral. Mng. 62A: 999-1000. Olivo, G. & A.E. Williams-Jones 1999. Hydrothermal REErich eudialyte from the Pinalesberg Complex, South Africa. Can. Min. 37: 653-663. Pearson, R.G. 1963. Hard and soft acids and bases: J . Am. Cliem. Soc. 85: 3533-3539. Salvi, S. & A.E. Williams-Jones 1996. The role of hydrothermal processes in concentrating HFSE in the Strange Lake peralkaline complex, northeastern Canada. Geochim. Cosnzochinz. Acfn. 60: 1917-1932. Salvi, S., Fontan, F., Monchoux, P., Willianls-Jones, A.E. & B. Moine 2000. Hydrothermal mobilization of HFSE in alkaline igneous systems: Evidence from the Tamazeght Complex (Morocco). Econ. Geol. 95: 559-576. Tagirov, B., Schott J., Harrichourry, J.-C. & S.Salvi 2001. Experimental study of aluminum speciation in chloride- and fluoride-rich supercritical fluids. Ceochim. Cosnzochim. Actn (in review). Vard, E. & A.E. Williams-Jones 1993. A fluid inclusion shidy of vug minerals in dawsonite-altered phonolite sills, Montreal, Quebec: Implications for HFSE mobility. Contrib. Mineral. Petrol. 113: 410-423. Zhu, C. & D. Sverjenskii 1992. F-C1-OH partitioning between biotite and apatite: Geochim. Cosnzochim. Actn. 56: 34353467.

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Zaraisky, G. 1994. The influence of acidic fluoride and chloride solutions on the geochemical behavior of Al, Si and W. In K. Shmulovich, B. Yardley, and G. Gonchar (eds), Fluids in the CrList; eqililibrirtn1 arid transport properties; 139161. London: Chapman & Hall.

Water-Rock Interaction 2001, Cidu (ed.), 02007 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Time-depth-temperature relations for igneous, metamorphic and hydrothermal processes: Visualized. through simplified-model numerical simulations Hiroshi Shigeno Geological Survey of Japan, Tsukuba, Ibaraki, Japan

ABSTRACT: Time-space temperature distributions of igneous-metamorphic-hydrothermalsystems have been studied in a semi-quantitative manner by numerical simulations based on simplified one-dimensional transient 'extended thermal conduction' models. A series of results obtained by changing systematically top depth and thickness of static geothermal reservoirs and magma chamber were visualized, and their effects on large variations of time-space temperature distributions were discussed. Evolution history of the Kakkonda igneousmetamorphic-hydrothermal system, where temperatures over 500°C were observed in a late Quaternary granitic body by deep drilling, was explained through simulation with a dynamic environment model: temporal changes of dominant heat transport mechanism from thermal conduction to convection at the present reservoir depth.

1 INTRODUCTION

Htot = Nu Hcond Hcond = Krn (TL - T U ) I L

Thermal aspects of igneous, metamorphic and hydrothermal processes are important but very complicated due to various kinds of controlling factors. Based on simplified one-dimensional transient macroscopic models applying the 'extended thermal conductivity' to convective hydrothermal systems (Shigeno 1999), simulation studies have been conducted on the subject, especially for the purpose of improving exploration and evaluation of deep geothermal resources (ca. 2.0 to 4.0 km depth) in Japan (e.g. Shigeno 1995, 2000~). In the present paper, general results based on simple hypothetical static models will be shown. More complex results based on concrete dynamic models will be explained for the thermal history data obtained from the Kakkonda geothermal area, Japan. 2 MODELING AND SIMULATION METHODS Shigeno (1999) proposed a simplified vertical onedimensional transient thermal conduction model for better understanding through visualizing the diversities of macroscopic features of magma (igneous)hydrothermal systems. In the model, the 'extended thermal conductivity' (Kext) was defined by the following equations (1) to (4), based on a onedimensional steady -state thermal convection layer with constant top and bottom boundary temperatures:

Htot Kext

(1) (2)

(TL - T U ) / L Km

= Kext = Nu

where Htot, Nu, Hcond, Km, TL, TU, and L were total heat flow by conduction and convection (W), Nusset Number (-), heat flow by conduction (W), thermal conductivity of the layer (W/m-K), bottom and top boundary temperatures (OK), and thickness of the layer (m) ,respectively. Figure 1 shows the general model for the present study. Major points are as follows (Shigeno 1999, 2000a, 2000b): 1. It is a vertical one-dimensional transient thermal conduction model. Simulation depth and time are from the surface to 20 km depth, and fiom instant magma chamber emplacement to 400 ky later, respectively. 2. Distribution of averaged units (magma chamber, geothermal reservoir, or non-magma-and-nonreservoir) is assumed at each depth. Fluid phases are included in them. Depths of the top and bottom ofthe magma chamber and reservoir are essential parameters. 3 . The reservoir is assumed to have a high Kext value, ten times higher than that of other units (2.5 Wlm-K). Heat capacity and density ofall the units are assumed to be constants: 1.O Wkg-K and 2700 kglm3, respectively. 4. Boundary and initial conditions for temperature distributions, and latent heat of magma consolidation are assumed as shown in Figure 1.

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Based on the above model, forward simulations using explicit difference equations were conducted.

3 RESULTS AND DISCUSSION FOR HYPOTHETICAL MODEL STUDY

Figure 1. Simplified one-dimensional transient thermal conduction model with 'extended thermal conductivity' for macroscopic . indicate magma (ignmus>.hydrothe-ma] systems. The * parameters whose values were changed fbr kiur hypothetical models and models of the Kakkonda a m , respectively.

For better understanding macroscopic time-dep thtemperature relations for igneous, metamorphic and hydrothermal processes, a systematic simulation study has been conducted using four hypothetical reservoir distribution models, A, B, C and D, based on the above general model (see Table 1). For each model. the top depth and thickness of the magma chamber were changed from 3 to 6 km, and from 1 to 8 km, respectively (Shigeno 2000b). Figure 2 summarizes the simulation results for the temperature distributions at 3 km depth changingwith time. ~i~~~~ 3 shows examples of the simulated temperature distributions On time-depth planes. The km depth temperatures tend to be r e d a t e d mainly by the top depth of magma chamber ifthe time

Figure 2. Summary ofsystematic simulation results kir temperature distributions at 3 km depth. Four rows, and hur columns correspond to four hypothetical models, and the time after magma chamber emplacement, respectively. Each small figure shows temperature contours (50°C intervals) at 3 km depth on the plane of the top depth and thickness of magma chamber.

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Table 1. Four hypothetical reservoir distribution models 6 r simulating various igneous-metamorphic-hydrothermal systems. Numbers of Depth of 'Kext' (W/m-K) reservoir(s) reservoir(s) value of reservoir(s) Model A 0 .-_-----...-____________ ____ __ Model B 2 1.0-2.0 km and 2.5-3.5 km depth. 25.0 Model C I 1 .O km-magma chamber top depth. 25.0 Model D 1 2.5 km-magma chamber top depth. 25.0 -.___

Figure 3. Examples ofsimulation results for hypothetical model study. Four rows and three columns correspond to four reservoir models, A, B, C and D, and three cases of magma chamber distributions, respectively. Each small figure shows temperature contours (50°C intervals) on the depth and time plane br the indicated reservoir model and magma chamber distribution.

after its emplacement is less than 150 ky, but mainly by its thickness if more than 150 ky . The reservoir attached directly to the magma chamber top and covered by a thick cap rock of Model D (possibly of concealed nature) tends to keep high to very high temperatures (250" to 500°C) at 3 km depth regardless of the magma chamber conditions. However, a similar reservoir covered by a thin cap rock of Model

C is much colder, and can hardly keep the 3 km depth temperatures higher than 250°C provided that the thickness of the magma chamber is less than 2 km. The deep reservoir of Model B can hardly keep the 3 km depth temperatures higher than 250°C provided that the distance between the reservoir bottom and magma chamber top is more than 1.5 km regardless of the thickness of the magma chamber. The conductive systems of Model A corresponding to contact metamorphic environments can keep high temperature distributions at depth for long periods. The above simulation results show that many possibilities exist for hy drothermal-sy st em development at depth even if thermal conduction dominates at shallow levels. The above time-space-temperature relations could be useful guidelines for exploration and evaluation of deep geothermal resources that will be exploited in the future (Shigeno 2000b, 2000~). 4 RESULTS AND DISCUSSION FOR KAKKONDA AREA STUDY At the Kakkonda geothermal area, a large late Quaternary granitic body is distributed below ca. 2.5 km depth from the surface associating with past hightemperature contact metamorphic zones over 1 km

Figure 4. Comparisons of measured (A) and simulated (B) temperature profiles br the past and present time fbr the WD-la well, Kakkonda geothermal (igneous-metamorphic-hydrothermal) area, Japan. Figure (C) shows an optimal simulation scenario, time-space distribution changes of the 'Kext' values, which produced the temperature profiles (B).

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thick in the overlying pre-Neogene and Neogene formations. A meteoric-water-dominated type hydrothermal system developed at the present time has been exploited for power generation by two stations with a total capacity of 80 M We. During the 'DeepSeated Geothermal Resources Survey' in this area, exploration well WD-Ia was drilled to 3729 m depth in 1995, and a thermal conduction zone with amaximum temperature over 500°C was observed at the deepest part of the well (Fig. 4 A) (e.g. Uchida et al. 1996, Sasaki et al., 1998). In this case, applied simulation models, based also on the general model, were more complex using various Kext values and dynamic conditions for reservoirs and magma chamber in order to solve the complex thermal history (Shigeno 2000a). The magma chamber was assumed to be thick (8 km) to cope with the very rapid cooling of the reservoirs covered by a thin cap rock. The time-space temperature distribution data obtained from WD- 1a can be explained fairly well by the following dynamic environment model (Fig. 4 B & C): (1) At first, the thick contact metamorphic zones (ca. 350"-650°C at 1.5 to 2.8 km depth) were produced through heat storage by thermal conduction from the magma chamber (top depth 2.8 km deep) emplaced probably ca. 200 ka; (2) Later (e.g. 30 ka), reservoir formation caused the present convective temperature profile (ca 220"-380°C at 0.5 to 3.1 km depth) with the high conductive temperature gradient below 3.1 km depth in the cooling Quaternary granitic body. Adjusting the model for contributions of magmatic fluid and magma convection by applying a high Kext value (7.5 W/m-K), which enhanced heat transport from the magma chamber to the overlying formations, for the early stage improved temperature profile fitting for the contact metamorphism. 5 SUMMARY In the present paper, time-depth-temperature relations for igneous-metamorphic-hydrothermal systems were visualized and discussed in a semiquantitative manner based on simulation results using simplified one-dimensional transient 'extended thermal conduction' models (Fig. 1). 1. A series of the results based on the four hypothetical geothermal reservoir distribution models, A, B, C and D, changing systematically the top depth and thickness of magma chamber-igpeous body were visualized (Figs. 2 & 3). Effects ofthe distributions of various types of hydrothermal systems, and the depth and amount of initially stored heat energy on large variations of time-space temperature distributions, especially at 3 km depth, were summarized. 2. A case study result for the Kakkonda geothermal area, where temperatures over 500°C were observed in a late Quaternary ganitic body by deep drilling was reported. Simulation analysis revealed that the pres-

ent temperature profile and wide distributions of the past high-temperature contact metamorphic zones can be explained fairly well by a dynamic model for the dominant heat transport mechanism that changed from thermal conduction to convection at the present reservoir depth (Fig. 4). The method of this study might seem too simple, but could be very usefbl because of its high flexibility for modeling and simulating complex real conditions. The above results encourage applications of the method to other igneous-metamorp hic-hy drothermal systems for improving macroscopic and quantitative understanding. ACKNOWLEDGMENTS

I thank the group members of the 'Data Analyses and Evaluations for the Deep-seated Geothermal Resources Survey' at GSJ for their cooperation and discussion. The present study was conducted with financial support by the office of the New Sunshine Program', AIST, MITI, and through cooperation with NEDO, JMC, GERD and WJEC. REFERENCES Sasaki, M., K. Fujimoto, T . Sawaki, H. Tsukamoto, H. Muraoka, M. Sasada, T . Ohtani, M. Yagi, M. Kurosawa, N. Doi, 0. Kato, K. Kasai, R. Komatsu& Y. Muramatsu 1998. Characterization of a magmatidmeteoric transition mne at the Kakkonda geothermal system, northeast Japan. In G.B. Arehart & J.R. Hulston (eds), Proc. Water-Rock Interaction 9: 483-486. Rotterdam: Balkema. Shigeno, H. 1995. Estimating deep environments of Japanese hydrothermal systems based on geochemical data @om geothermal power plants. Proc. World GeothermalCongress95, Florence 1995: 101 9- 1024. Shigeno, H. 1999. Fundamental study k r diversities of magma-hydrothermal system environments based on simplified-model numerical simulations. Bull. G a l . Surv. Japan 50: 725-741 (*J). Shigeno, H. 2000a Evolution history of the Kakkonda magma-hydrothermal system, Japan, estimated through simplified-model numerical simulations. Proc. 25th Workshop on Geothermal Reservoir Engineering, Stanforci Univ. 2000: 135- 142. Shigeno, H. 2000b. Systematic tbrward and preliminary inversion analyses fa magma-hydrothermal system environments based on simplified-model numerical simulations. Bull. Geol. Sum. Japan 5 1 : 63 1-648 (*J). Shigeno, H. 2000c. Toward hture progress of exploration and exploitation fbr deep geothermal resources in Japan. Rept. Geol. Surv. Japan, no. 284: 3 15-338 (*J). Uchida, T., K. Akaku, M. Sasaki, H. Kamenosono, N.Doi & S . Miyazaki 1996. Recent progress of NEDOs "Deep-seated Geothermal Resources Survey" project. Geothermal Resources Council Transactions 20: 643-648. (*J: in Japanese with English abstract).

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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Mass Transfer, Oxygen Isotopic Variation and Gold Precipitation in Epithermal Systems: A Case Study of the Hishikari Deposit, Southern Kyushu, Japan Naotatsu Shikazono, Noriyuki Yonekawa & Takahumi Karakizawa Department of Applied Chemistry, Faculty of Science and Technology,Keio University

ABSTRACT: Geochemical and mineralogical data (bulk composition, hydrothermal alteration, oxygen isotopes), thermodynamic considerations and dissolution kinetics-fluid flow coupling model indicate that mixing between hydrothermal solutions and acidic groundwater plays an important role in hydrothermal alteration and gold deposition at the Hishikari epitermal gold mining district, southern Kyushu, Japan. 1 THE HISHIKAIU DEPOSIT The Hishikari gold mining district is located at southern part of Kyushu, Japan. The district is composed of sedimentary rocks of the pre-Neogene Shimanto Supergroup (dominantly shale and sandstone) and Quaternary andesitic and dacitic volcanic rocks. The deposit consists of three vein systems, the Honko, Sanjin and Yamada veins which are hosted by Quaternary andesites and Shimanto sedimentary rocks.

2 ANALYTICAL PROCEDURE AND RESULTS XRD (X-ray diffraction) analyses of the collected samples (about 140 samples) indicate that hydrothermal alteration zones are divided into zone I (cristobalite-smectite zone), I1 (quartz-smectite zone), I11 (quartz-mixed layer zone) and IV (quartzchlorite zone). Bulk compositions of altered andesite and shale were obtained by XRF (X-ray fluorescence analysis). The XRF data indicate the following features of compositional variation of altered rocks. (1) K20 and MgO contents of altered andesite decreases away from the vein. (2) Na2O and CaO contents of altered andesite increase with the distance from the vein. ( 3 ) Analytical data are plotted on (Na20+CaO)-K20 diagram. This shows that @a20 + CaO) content inversely correlates with K20 content. (4) Si02 content tends to irregularly vary near the veins. These results are consistent with XRD results. The amounts of K-feldspar, K-mica and chlorite are

higher closer to the veins and Ca-zeolites and smectite decrease in amounts towards periphery of the alteration zones. 3 KINETICS-FLUID FLOW-MIXING MODEL

Hydrothermal solution containing appreciable amounts of H&04 migrates through andesitic volcanic rocks, accompanying SiO2 (quartz and cristobalite) precipitation. Decrease in temperature due to heat conduction alone cannot explain the distributions in quartz and cristobalite; Temperature at which heat conduction trend crosses cristobalite saturation curve is ca. 200°C which is higher than 100°C corresponding to cristobalite/quartz boundary in active geothermal system. The curve for mixing of hydrothermal solution and groundwater lie always below cristobalite saturation curve. Therefore, the heat conduction and mixing of fluids alone are considered to be not main cause for the precipitation of quartz and cristobalite. Therefore, in order to know the change of H4Si04 concentration and temperature during the precipitation of quartz and cristobalite and mixing of fluids, the following equation was used.

in which C: concentration of H&04 of output fluid (mol/kgH20), t: time (sec), k: rate constant (sec-'), CO:saturation concentration of H4Si04 (mol/kgHzO), C1: concentration of H4SiO4 of input hydrothermal solution (mol/kgH20), C2: concentration of H4Si04 of input groundwater (mol/kgH20 , 91: volume flow rate of hydrothermal solution (m /sec), q2: volume

3

753

5 GOLD PRECIPITATION DUE TO MIXING OF FLUIDS IN EPITHERMAL SYSTEMS

flow rate of groundwater (m3/sec), M: mass of aqueous solution in a system (kg), V: volume of aqueous solution in a system (m3), A: surface area of rocks which contacts with aqueous solution (m2). Assuming the ranges of A/M, flow rate of mixed fluid, porosity and giving the precipitaiton rate of SiO2 minerals, the relationship between H4Si04 concentration of mixed fluid and tempearture was obtained. The results of calculation indicate that the flow rate, 10-4.2m/sec,is the best estimate for A/M=O.lO. The flow rate of geothermal water in the reservoir is generally 1o - -~ 1o4rn3/sec. In the present model, the area considered is (2kmx0.5km) and flow rate is 1044.2m/sec. When porosity is 2%, total mass flow rate is estimated as 10-4.2mx106m2x2/100=1.4x106g/sec. This is very similar to that in the Wairakei geothermal area in New Zealand which is 1.3x 106g/sec. (Elder 1966).

Shikazono and Nagayama (1993) favored the two fluids mixing as a cause for the gold deposition, and the presence of acid alteration zone overlying the Hishikari gold-quartz veins. For example, alunite occurs in higher elevations in the eastern part of the Hishikari district. Shikazono (1985) suggested based on the studies of wall rock alterations associated with Japanese epithermal-type Au-Ag deposits that acid sulfate fluids coexist with near-neutral chloriderich fluids at relatively shallow zone from the surface (less than lkm) at the time of epithermal-type Au-Ag ore formation. Reed and Spycher (1985) made a computation on the gold precipitates e E ciently from the mixed fluids. These results support that gold deposition occurred in the Hishikari veins by the mixing of two fluids (hydrothermal solution and acid sulfate groundwater). However, it is uncertain that the alteration minerals (K-feldspar, K-mica and kaolinite) formed in the same stage of hydrothermal alteration. It may be possible that descending acid low temperature solution (groundwater) formed kaolinite and cristobalite at different stage of gold-quartz mineralization associated with K-feldspar precipitation.

4 INTERPRETATION OF OXYGEN ISOTOPIC VARIATION IN HYDROTHERMAL ALTERATION ZONE BY TWO FLUIDS (HYDROTHERMAL SOLUTION AND GROUNDWATER) MIXING MODEL

6 SUMMARY AND CONCLUSION

Naito et al. (1993) showed that oxygen isotopic composition (6"O) of volcanic rocks in the Hishikari mining area vary systematically; +5.9 to +15.9%0(zone I), +7.1 to +12.4 %O (zone II), +2.8 to +l 1.7 %O (zone 111), and +2.1 to +8.2 %O (zone IV). They calculated the change in 6l80 values of hydrothermally altered volcanic rocks as a function of water-to-rock ratio by weight, assuming isotopic equilibrium in a closed system with several different temperatures and demonstrated that the increase in 6"O values from the veins towards peripheral zones can be interpreted as a decrease in temperature from the vein system. In their calculations, the effect of mixing of hydrothermal solution with groundwater was not taken into account. Here we interpret 6l'O zonation based on hydrothermal solution-groundwater mixing model. Initial 6l'O value of hydrothermal solution (O%O) was estimated from 6l'O values of K-feldspar and quartz in the veins. Initial 6"O value of groundwater (-7%o) was estimated from meteoric water value of the south Kyushu district (-7%0). Mixing ratio of hydrothermal solution and groundwater was calculated based on the tem erature of each reservoir. Using the estimated 6l sp0 of fluid and oxygen isotopic fractionation between water and mineral, we estimated 6l'O of altered rocks. Results show a fairly good agreement between the model calculation and analytical data on 6l'O except the data for the most peripheral zone.

(1) Hydrothermal alteration zoning in the Hishikari gold mining district from the Au-Ag veins towards peripheral and shallower part is as follows; quartz-chlorite zone (zone IV), quartz-mixed layer zone (zone 111), quartz-smectite zone (zone 11), cristobalite-smectite zone (zone I). (2) XRF analyses of samples from cores of two drill holes and from underground indicate a relative enrichment of K and Mg and depletion of Ca and Na near the Au-Ag vein, no systematic but irregular variation of Si02, and relatively constant variations of Al, Ti, Mn, and Fe. (3) The variations in K, Ca, and Na contents roughly correlate with the variation in 6"O of altered rocks. (4) The variations in K, Na, and Ca contents and alteration zoning from K-minerals (K-feldspar, Kmica) near the veins to kaolinite towards marginal part are explained by the temperature decrease due to the mixing of hydrothermal solution containing high K' concentration, low Na and Ca concentrations with acid groundwater and destruction of feldspar by K-rich hydrothermal solution. (5) The variation in Si02 content of altered rocks and zoning of Si02 minerals cannot be explained by the thermodynamic equilibrium between hydrothermal solution and Si02 minerals. 754

(6) Precipitation kinetics-fluid flow-mixing model was applied to explain the variation in mineralogical zoning of SiOl minerals (quartz at center and deeper and cristobalite at peripheral and shallower parts) by putting the values of AIM and flow rate of hydrothermal solution which satisfy the SiOz mineral zoning and temperature of cristobalite/quartz boundary. (7) 6l80 of altered rocks can be also explained by the mixing of two fluids of hydrothermal solution and acid groundwater. (8) The mixing of two fluids can cause also efficient deposition of gold in epithermal systems. Consequently, the most important conclusion of this paper is the role of mixing processes involving hydrothermal solutions and acid low-temperature groundwater causes hydrothermal alteration zoning, variations in chemical composition and 6l80 of altered rocks, and deposition of gold in epithermal systems. REFERENCES Elder, J. W. 1966. Heat and Mass Transfer in the Earth: Hydrothermal Systems. B. NZ. Dep. Sci. Ind. R. 169: 1-1 15. Naito, K., Matsuhisa, Y., Izawa, E. & H. Takaoka 1993. Oxygen isotopic zonation of hydrothermally altered rocks in the Hishikari gold deposits, southern Kyushu, Japan. Its implication for mineral prospecting. Resource Geol. Spec.Issue, N0.14: 71-84. Reed, M. & N.F. Spycher 1985. Boiling, cooling, and oxidation in epithermal system: A numerical modelling approach. In Geology and Geochemistry of Epithermal System. Rev. Econ. Geol. 2: 249-26 1. Shikazono, N. 1985. Mineralogical and fluid inclusion features of rock alterations in the Seigoshi gold-silver mining district, Izu Peninsula, Japan. Chem. Geol, 49: 213-236. Shikazono, N.& T. Nagayama 1993. Origin and depositional mechanism of the Hishikari gold-quartz-adularia mineralization. Resource Geol. Sec. Issue No. 14: 47-56.

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Wafer-RockInteraction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2

Overprinted Cenozoic hydrothermal activities at the Toyoha Ag-Pb-Zn deposit, Japan T.Shimizu & A.Aoki Mineral and Fuel Resources Department, Geological Survey of Japan, Japan

ABSTRACT: This paper presents the results of mineral alteration, WAr ages, fluid inclusion and stable isotope studies in order to reconstruct a hydrothermal history of the Toyoha-Muine geothermal system near the giant Toyoha Ag-Pb-Zn deposit in Hokkaido, Japan. Alteration by neutral-pH water is pervasive and abundant in the studied area. The K-Ar ages show that three hydrothermal regimes (24.6 Ma, 8.6-8.4 Ma and 2.90.4 Ma) are juxtaposed in a single area. Fluid inclusion studies show that ore minerals precipitated at around 212-290°C from diluted fluids (0.7 to 3.4 wt% NaCl equiv.). Oxygen and sulfur isotope studies indicate that metals were transported from magma, and precipitated as ore minerals from the mixtures of magmatic hydrothermal fluids and meteoric groundwater. During periods of 8.6-8.4 Ma and 2.9-0.4 Ma, some base metals could be remobilized from the older deposits. The Toyoha-Muine Ag-Pb-Zn mineralization could be partially remobilized from older Ag-Pb-Zn mineralizations. 1 INTRODUCTION

The Toyoha-Mt. Muine area, located in southwestern Hokkaido, Japan, has a good potential for the discovery of new Ag-Pb-Zn ore deposits as well as for geothermal exploration. The Toyoha deposit has been one of the most productive Ag-Pb-Zn deposit in Japan. The deposit consists of more than 50 veins. Total Production from the deposit is 2350 t Ag, 413,000 t Pb, and 1,230,000 t Zn (Watanabe & Ohta 1995). Discovery of various rare elements such as Sn, In and W (Yajima & Ohta 1979; Ohta 1980) as well as a Cu-rich zone (Yoshie et al. 1986) has revealed the polymetallic character of the deposit. A few previous studies have been conducted on the hydrothermal alteration accompanied with basemetal mineralization in the region. The distribution of alteration zones and K-Ar ages around the Toyoha deposit have been documented by Sawai (1984; 1986), Marumo & Sawai (1986) and Sawai et al. ( 1989); however, there have been no detailed studies on the hydrothermal alteration combined with fluid inclusion and stable isotopes relative to the evolution of the hydrothermal system. This work mostly focuses on the more than 7UUm of hydrothermally altered rocks cored in drill holes 9MATY-1 and 9MATY-2 in the southeastern extension of the Toyoha Ag-Pb-Zn deposit. In this paper, alteration mineral assemblage, K-Ar dating, fluid inclusion homogenization temperatures and salinities, and stable isotopes are used to document the evolution of the Toyoha-Muine geothermal field.

2 GENERAL GEOLOGY Toyoha-Muine area is closely located at the conjunction of Northeast Japan Arc and Kril Arc (Fig. 1). The host rocks consist of early to middle Miocene sediments, pyroclastic rocks, and lavas (Fig. 2). They are divided into three formations: Koyanagizawa, Motoyama, and Nagato Formations. The Koyanagizawa Formation is composed of altered basalt and andesite lavas. The Motoyama Formation consists of a basal conglomerate, sandstone and mudstone. The Nagato Formation is characterized by a large amount of pyroclastics and lavas of altered andesite, dacite and ryolite (Kuwahara et al. 1983; New Energy and Industrial Technology Development Organization 1988). 3 HYDROTHERMAL ALTERATION Drill holes 9MATY-1 and 9MATY-2 dip 30" and 25" south, and extend 1051 m and 1201 m in the southeastern extension of the Toyoha deposit, respectively (Fig. 1 and 2). Rocks in both cores generally display moderate to strong hydrothermal alteration. Alteration is divided into three zones on the basis of its characteristic mineral assemblage. Zone I extends from 300 to approximately 800 m and is characterized by the presence of quartz, albite, calcite, laumontite, chlorite, and interstratified sericite/smectite. It is also characterized by veins con7c7

6 STABLE ISOTOPE STUDIES (0,S)

/

Ag-l'b-%n vein Landslide scars

Oxygen isotope values were obtained on quartz from veins by the CO2 laser microprobe analytical technique (Sharp 1990). Quartz values are converted to fluid values using an equation of quartz-water isotopic equilibrium (Matsuhisa et al. 1979). The calculated &'gofluidvalues are - 4.2 to - 4.7 per mil from the 9MATY-2 core. Sulfur isotope values were obtained on pyrite, sphalerite and galena from 9MATY-I and 9MATY-2 cores by a conventional method (Sasaki et al. 1979; Robinson & Kusakabe 1975). The 634Svalues from ore minerals range from 2.3 to 8.1 per mil.

I'utnarole

7 DISCUSSION Figure 1. Location of Toyoha-Muine geothermal field.

taining variable proportions of calcite and quartz. Zone I1 is dominantly present around Ag-Pb-Zn veins from 800 m to 1200 m. Main alteration minerals are quartz, chlorite, sericite, pyrite and calcite. Zone 111 is locally present between 747 and 767 m in 9MATY- 1 where Ag-Pb-Zn minerals are rare. This zone is characterized by quartz, pyrophyllite, sericite and pyrite. 4 K-Ar AGES The K-Ar ages of samples from hydrothermally altered rock around the veins from drill holes 9MATY-I and 9MATY-2 were obtained (Table 1). The localities of the sample used for age determination studies are shown in Figure 2. The specimens were crushed and clay particles were separated by hydraulic elutriation. Above 800 m, where zone I is dominant, the ages are 1.87-0.5 Ma, whereas at greater depths, where zone I1 and 111 are dominant, the ages show 8.6-8.4 Ma and 24.6 Ma. 5 FLUID INCLUSION STUDIES Fluid inclusion homogenization temperature and salinity measurements were made on doubly polished thin plates of sphalerite and quartz from 9MATY-I and 9MATY-2 cores, using a Fluid Inc.adapted USGS-type heating/freezing system calibrated with synthetic fluid inclusions. The accuracy is estimated to be 4 0.1"C at 0°C and k 2°C at 374°C. Homogenization temperatures of primary fluid inclusions range between 212 and 290°C in sphalerite, and between 212 and 265°C in quartz. Salinity of the fluid inclusions varies from 0.7 to 3.4 wt percent NaCI equiv. in sphalerite, and from 0.2 to 1.9 wt percent NaCl equiv. in quartz.

Previous studies show that the hydrothermal events, which were responsible for Ag-Pb-Zn mineralization at the Toyoha deposit, occurred at 2.9 to 0.4 Ma (Marumo & Sawai 1986; Sawai et al. 1989). Aoki et al. (1999) revealed that the Pleistocene (3.7-2.2 Ma) acidic alteration zone around Mt. Muine overlapped with the middle to late Miocene (1 1.7-9.0 Ma) neutral-pH alteration zones which was responsible for the Ag-Pb-Zn mineralization. The acidic alteration was triggered by Muine volcanism (3.OMa; Watanabe 1990). Combined with the previous data, we conclude that at least three hydrothermal regimes have occurred in the Toyoha-Mt. Muine area; the earliest hydrothermal activity began in the early Miocene (24.6 Ma) followed by an intermediate thermal event in the late Miocene (1 1.7-8.4 Ma). The latest (2.9-0.4 Ma) thermal regime was associated with Pleistocene to recent hydrothermal activity. Based on the data of the mineral alteration zoning and WAr ages, the youngest Ag-Pb-Zn deposits including the Toyoha overprinted the other older deposits. Some metals could be remobilized from the older deposition during the evolution of hydrothermal activities. Fluid inclusion data indicates that the ore veins were formed at 212-290°C from dilute fluids (0.73.4 wt%) in both the late Miocene and Pleistocene mineralizations. These values show a similar variation in both temperature and salinity for the Toyoha deposit (1 5O-30O0C, 0.2-3.4 wt%; Yajima and Ohta 1979) except for values from the most southeastern vein, Shinano (20O-32O0C,0.5-8.5 wt%; Sanga et al. 1992). The oxygen isotope values (- 4.2 to - 4.7 %o) are similar to the values (- 3.2 to -6.5%0;Matsuhisa et al. 1979) from the late stage of the Toyoha Ag-Pb-Zn deposit. As in the case of the Toyoha deposit, we conclude that the source of the fluid for the Ag-PbZn mineralization in our studied area was derived from the mixture of meteoric groundwater and deep hydrothermal fluids. Sulfur isotope values from ore 758

Figure 2. Cross section of geo-logy and K-Ar ages in the southeast extension of the Toyoha AgPb-Zn deposit (Modified after Shimizu et al. 1997). A.S.L refers to above sea level. minerals close to 5 per mil are similar to the values consistent with the sulfur of magnetite-series granitoids magma in Japan (Sasaki and Ishihara 1979). The oxygen and sulfur isotopes indicate that metals were originally transported by deep hydrothermal

fluids from magnetite-series granitoids magma. In both period between 8.6-8.4 Ma and 2.9-0.4 Ma, mixing of deep fluids and meteoric ground water caused the dilution of fluids, resulting in the precipitation of ore minerals.

Table 1. K-Ar ages from 9MATY- 1 and 9MATY-2 cores in the southeast extension of the Toyoha Ag-Pb-Zn deposit Sample name

Mineral

Ar x IO-~scc/gm

% 4oAr

355.3 m

Sericite

0.04 I 0.039

28.4 25.5

7.13 7.10

1.45k0.09

419.85 m

Sericitekmectite

0.01 I 0.0 I2

4.4 6.2

3.48 3.47

0.9k0.3

649.8 m

Sericite

0.032 0.033

13.7 13.5

4.46 4.46

1.87f0.25

1003.6 m

Sericite

0. I89 0. I86

46.8 46.4

5.76 5.75

?o'

K

Age (Ma)

9MATY-I

.............................

--------- ---

8.4k0.4

-__---------

9MATY-2 383.9 m

Seric i tekrnect i te

0.0039 0.0036

3.8 6.5

I .75 I .74

0.520.2

91 1.6 m

Sericite

0.408 0.409

59.6 65.3

4.24 4.25

24.6f1.2

1186.55 m

Sericite

0.225 0.207

64. I 66.3

6.49 6.47

8.6k0.4

The measurement were carried out by Teledyne Isotope Co., Ltd. Constants used for the age calculation are; hp = 4.96*1O-lO/yr, he:= 0.581*10-lo/yrand 40WK = 1.167*10-2(atom.%). 759

8 CONCLUSIONS The K-Ar measurements demonstrate an evolution pattern of initial hydrothermal activity (24.6 Ma) overprinted by the younger activities (8.6-8.4 Ma and 2.9-0.4 Ma). Deep hydrothermal fluid transported metals from magnetite-series granitoids magma. In the 8.6-8.4 Ma and 2.9-0.4 Ma hydrothermal events, some metals may have been leached from the older deposits. Ore minerals precipitated at 212-290°C by mixing of deep hydrothermal fluid and meteoric groundwater. REFERENCES

Sawai, O., Okada, T. & T. Itaya 1989. K-Ar ages of sericite in hydrothermally altered rocks around the Toyoha deposits, Hokkaido, Japan. Mining Geol. 39: 191-204. Yajima, J. & E. Ohta 1979. Two-stage mineralization and formation process of the Toyoha deposits, Hokkaido, Japan. Mining Geol. 29: 291-306. Yoshie, T. Narui, E. & K. Kato 1986. On the process of mineralization and distribution of minor elements in the Toyoha deposit, Hokkaido. Mining Geol. 36: 179-193 (in Japanese with English abstract). Watanabe, Y. 1990. Pull-apart vein system of the Toyoha deposit, the most productive Ag-Pb-Zn vein-type deposit in Japan. Mining Geol. 40: 269-278. Watanabe, Y. & E. Ohta 1995. The relation of two stage mineralization at the Ag-Pb-Zn Toyoha deposit, southwest Hokkaido, to subduction of the Pacific Plate. Resource Geol. Special Issue. 18: 197-202.

Aoki, M., Ohta, E., Shimizu, T., Watanabe, Y., Yoshie, T. & T. Kabashima 1999. The potentiality of ore deposits at the Muine-Toyoha hydrothermal system. Report on the regional geologicalIy structural survey in the south Hokkaido, 1998fiscalyea~Tokyo, Japan. 145 p. (in Japanese). Kuwahara, T., Miyazaki, T., Tani T. & K. Iida 1983. A characterization of the vein mineralizations at the Motoyama deposit, Toyoha mine from the viewpoint of their tectonic setting and ore assays. Mining Geol. 33: 115-129 (in Japanese with English abstract). Marumo, K. & 0. Sawai 1986. K-Ar ages of some vein-type deposits in the southwestern Hokkaido, Japan. Mining Geol. 36: 21-26 (in Japanese with English abstract). Matsuhisa, Y., Goldsmith, J.R. & R.N. Clayton 1979. Oxygen isotopic fractionation in the system quartz-albite-anorthitewater. Geochimica et Cosmochimica Acta 43: 1 131- 1140. Matsuhisa, Y., Yajima, J. & E. Ohta 1986. Oxygen isotope and fluid inclusion studies of quartz from Toyoha Ag-Pb-Zn deposit (abs.). Japanese Association of Mineralogists, Petrologists and Economic Geologists, the Mineralogical Society of Japan, and the Society of Resource Geology, Joint Annual Meeting, Abstracts with Programs: 26 (in Japanese). New Energy and Industrial Technology Development Organization. 1988. Report on geothermal energy exploration and development of the Toyoha area. 1156.p (in Japanese). Ohta, E. 1980. Mineralization of Izumo and Sorachi veins of the Toyoha mine, Hokkaido, Japan. Bull. Geol. Sum. Japan. 3 1 : 585-597. (in Japanese with English abstract). Robinson, W.B. & M. Kusakabe 1975. Quantitative preparation of sulfir dioxide 34S/32Sanalyses from sulfides by combustion with cuprous oxide. Anal. Chem. 47: 11791181. Sasaki, A. & S. Ishihara 1979. Sulfur isotopic composition of the magnetite-series and ilmenite-series granitoids in Japan. Contrib. Mineral. Petrol. 68: 107-115. Sasaki, A., Arikawa, Y., & E.R. Folinsbee 1979. Kiba reagent method of sulfur extraction applied to isotopic work. Bull. Geol. Surv. Japan, 30: 241-245. Sharp, Z.D. 1990. A laser-based microanalytical method for the in-situ determination of oxygen isotope ratios of silicates and oxides. Geochimica et Cosmochimica Acta, 54: 13531357. Sawai, 0. 1984. Wall rock alteration around the Motoyama deposits, Toyoha mine, Hokkaido, Japan. Mining Geol. 34: 173-186 (in Japanese with English abstract). Sawai, 0. 1986. The distribution of alteration zones in the eastem area of the Toyoha mine, Hokkaido, Japan. Mining Geol. 36: 273-288 (in Japanese with English abstract).

760

Water-Rock Inferaction 2001, Cidu (ed.). 0 2001 Swets & Zeitlinger, Lisse, lSBN 90 2651 824 2

Natural decay series studies of the Kujieertai uranium deposit, N W China Sun Zhanxue. Liu Jinhui, Li Xueli & Shi Weijun East China Geological Institute, Fuzhou, Jiangxi 344000, China

ABSTRACT Natural decay series disequilibrium techniques were used to provide information on the redox zoning and ore-forming processes of the Kujieertai uranium deposit, NW-China. The activity ratios of 234U/238U, 230Th/234U and 230Th/238U for about 26 samples were measured in the deposit. The characteristics of uranium and thorium isotopes for different redox zones are significantly different, which can serve as indicators for locating ore-bodies of the sandstone-type uranium deposit. 1 INTRODUCTION Natural uranium has three primary isotopes, 238U, 235Uand 234U,which occur at the present time in the proportion 99.28%238U/0.71 %235U/0.006%234U. The half lives of the isotopes are 4.5 1x 109, 7.1 x 108, and 2.44~ 10' years respectively. Because 238Uand 23sU both are parent elements and their physiochemical properties are similar to each other, the separation between the two isotopes is insignificant. For this reason, uranium isotopic ratio usually refers to 234U/238U. Thorium has six isotopes (234Th,230Th, 23'Th, 227Th,232Th,228Th),of which 230Thand 232Th are the most important ones. Radioactive disequilibrium between the long-lived members of the uranium natural decay series: 238U +234U+ 230Thcan readily be established during water-rock interactions as a consequence of (a) the pronounced insolubility of U4+,and (b) the preferential loss o f 2 3 4relative ~ to 2 3 8 from ~ minerals to solution in groundwater (Ivanovich & Harmon 1982, Shi 1990, Scott et al. 1992). Measurement of such radioactive disequilibrium has become a well established technique in the study of the formation mechanism of sandstone-type uranium deposits (Osmond et al. 1983, Liu 1999). In order to identify and characterize mineralization process, the natural decay series study of the redox zoning at the Kujieertai uranium deposit were carried out. The Kujieertai uranium deposit is a Jurassic sandstone-type mineral deposit, which is located in the southern margin of the Yili Basin, Xinjing Province, NW-China. In the deposit, there are vertically three

uranium-containing layers, Layer I, Layer I1 and Layer V. The oxidized zone of the lower layer, Layer V extends vertically over about 280 m. 2 SAMPLING, ANALYSIS AND RESULTS

26 samples collected from Layer 1-11 of the Kujieertai uranium deposit have been analyzed for U-series nuclides such as 238U,234Uand 230Th.The analytical methodology is briefly described as follows: After dissolution of the powdered samples, anion exchange method was used for chemical separation. Spike tracers were added prior to chemical separation, that is, 236Uor 232Ufor uranium and "'Th for thorium. The strong alkaline anion exchange resins (AG 1X8) were used to separate uranium fraction and thorium fraction from its matrix. TBP-xylene extraction was used for the further purification of uranium fraction. The uranium-containing solution was then preconcentrated by evaporation. Then uranium source disk and thorium source disk were prepared using electrodeposition techniques. Finally, the source disks were measured by alpha spectrometer. The analyzed results are reported in Table 1 and plotted in Figure 1. 3 DISCUSSIONS 3. I The strongly oxidized zone

As shown in Table 1 and Figure 1, most samples in the strongly oxidized zone lie in Sector A in Figure 1.

761

Table 1. Radioactive isotopic ratios of samples from Layer 1-11, the Kujieertai uranium deposit. Sample Code

Location (Borehole No.)

Depth Redox zoning

?34u,’38u

m

YL-33 YL 4 2 YL45 YL46 Y L-5 1 YL-53 YL-55

ZK4624 ZK4604 ZK4608 ZK4608 ZK4600 ZK4620 ZK4620

160.5 218.0 205.7 200.0 23 1 .O 181.6 170.8

0.179 0.415 0.132 0.061 0.147 0.120 0.118

0.166 0.466 0.143 0.069 0.286 0.121 0.123

0.158 0.491 0.177 0.074 0.393 0.128 0.127

0.927 1.123 I .083 1.131 1.946 1.008 1.042

0.952 1.054 1.238 1.072 I .374 1.059 I .033

0.883 1.183 1.341 1.213 2.673 1.067 1.076

YL-65 YL41 YL-56 YL 4 4 Y L-T6 YL-T 2 Y L-T 3 Y 1,-T 9 YL-T 10 YL-3 1 YL47 YL48 YL-50 YL-58 YL-39 YL-49 YL-52 Y L-54 YL-57

ZK4605 ZK4604 ZK4620 ZK4604 ZK3004 ZK3002 ZK3002 ZK3006 ZK3006 ZK4626 ZK4608 ZK4608 ZK4600 ZK46 12 ZK4628 ZK4608 ZK4600 ZK4620 ZK4620

175.8 234.0 158.5 236.8 172.0 188.0

0.033 1.047 0.102 0.239

0.027 1.036 0.113 0.264

0.033 1.061 0.135 0.269

I .222 1.024 1.195 1.019

2.57 1.298 3.31 1.928 1.838 0.101 0.030 0.119 0.351 0.022

2.420 1.208 3.1 10 1.652 1.73 0.102 0.036 0.116 0.353 0.023

2.25 0.744 3.26 1.706 1.57 0.141 0.037 0.117 0.360 0.026

0.818 0.989 1.108 1.105 1.73 1.26 1.15 1.15 I .02 0.942 0.93 1 0.940 0.857 0.941 1.010 1.200 0.974 1.006

I .ooo 1.013 1.324 1.126 0.17 0.07 0.03 0.03 0.11 0.875 0.573 0.985 0.885 0.854 1.396 1.233 0.983 1.026 1.182

2.00

185.0

196.0 202.0 2 17.5 202.0 221.5 230.4 199.3 227.0 23 1 .0 236.5 180.0 151.0

Weakly Oxidized Zone

Transitional Zone

Reduced Zone

3It suggests that significant increase of the 234U/238U ratios occurred during the oxidation process. The phenomenon was probably caused by adsorption of 234Uin groundwaters with characteristics of rich 234U that came from mountain areas by limoiiite and hematite of rocks in the strongly oxidized zone. In fact, the two minerals are quite rich in the zone (Liu 2000). The 234TW238U ratios of nearly all samples except for sample YL-33 are greater than unity, which implies that dissolution and removal of uranium within the last 3x105 years as speculated by MacKenizier et al. (1 992).

1.75 1.so

3 1.25 00

cu m

\ 1.00

3 d

0.930 0.619 1.048 1.003 0.908 1.382 1.028 I .009 1.020 0. I30

Strongly Oxidized Zone

0.75

3.2 The weakly oxidized zone 0.00

0.25 0.50 0.75 230

1.00 1.25 1.50 1.75 2.00

nl/ 2 3 8 ~

Figure 1 Plot of 234U/238U versus r30Th/23SUfor rock samples in the Kujieertai uranium deposit. SOZ = the strongly oxidized zone, WOZ = the weakly oxidized zone, TZ = the transitional zone, RZ = the reduced zone; I is the strongly oxidized zone, I1 is the weakly oxidized zone, 111 is the U mineralization zone, IV is the U deposition front, V is the reduced zone, and A is the complex sector.

Rocks in the weakly oxidized zone are inhomogeneous, and their radioactive isotopic ratios of samples are quite complex. The 234U/238Uactivity ratios of samples YL-65 and YL-41 are less than unity, which suggests that preferential loss of 234Ufrom rocks occurred within the last 106years and the 234U has remained in solution in groundwater for a sufficient length of time to be removed from the section of rock being studied. In addition, the230Th/234U ra762

tios of all samples in the zone being greater than unity shows that dissolution and removal of uranium within the last 3 x 1OS years. However the234U/238U ratios of samples YL-56 and YL-44 in the zone are greater than unity, which indicates that the rock has experienced deposition of uranium from groundwater within the last 106 years, and the deposition may have been continuous to the present time as pointed out by MacKenizier et al. (1 992). This may be explained by rolling dynamic mineralization process of sand-type uranium deposit. On one hand, uranium deposition formed from the oxygen and uranium-containing groundwaters from recharge area. On the other hand, leaching and migration Of PreviouslY formed Ore bodies occurred. Due to preferential dissolution of 2 3 8 rela~ tive to 230Th from rocks to groundwaters, the 230Th/238U ratios of rocks increased.

3.3 The transitional zone In the transitional zone, the 230Th/234Uratios for yL-31, yL-47, and yL-58 are less than unity, which suggests that it has experienced uranium deposition from groundwater within 3 x 1OS years and the deposition may have been continuous to the present time. The ratios for samples YL-48 and YL-50 are close to 1, which implies that equilibria between solid phase and groundwater are attained. Of the 10 samples taken from the transitional zone, the 234U/238 U ratios of 5 samples are less than 1, and of the rest 5 samples are greater than I , which reflects that uranium mineralization process is a rolling and multiple superimposed process. The old ore bodies were destroyed by oxygen-containing groundwater, and new ore bodies were formed in the lower course of the groundwater system. The residual old ore bodies lie in Sector 111, and newly formed ones lie in Sector IV in Figure 1. 3.4 The reduced zone

For almost all samples in the reduced zone, the 230Th/234Uratios are greater than unity, which suggests that they have experienced dissolution and removal of uranium within 3 x 10’ years. However the 234U/238Uratios are greater than unity, which indicates that uranium deposition have been occurred and continuous to the present time from about one million years ago. Obviously the water-rock interactions in this zone are quite complicated. Most samples taken from the reduced zone lie in the sector A

instead of falling in the expected sector V in Figure 1, which implies such complexity. The 234U/238U ratios being less than expected indicates that partial oxidation happened in the reduced zone probably due to oxygen-containing groundwater conducted by some fractures moving into the zone. The partial oxidation results in preferential loss of 234Ufrom rocks and decrease o f 2 3 4 ~ / 2ratios. 38~ 4 CONCLUSIONS

Based on the above discussions. The conclusions can be reached as follows: 1) The strongly oxidized zone: Of the most rock samples, the 234u/238u activity ratio are greater than unity instead of being less than unity due to the adsorption of by hematite and limonite rich in the zone, and the 230Th/234Uratio is greater than unity, which implies that dissolution and removal of uranium has occurred within 3 x 10’ years. 2) The weakly oxidized zone: Of almost all rock 3 4 ~ samples, the ratios of 2 3 4 u j 2 3 8 u and 2 3 0 ~ ~ 2 are greater than unity, which reveals that both dissolution and deposition coexist in the zone, and that uraniuln ore-formation and destruction is a process. In the zone, Some previously formed uranium mineralization may be destroyed by late corning oxygen-containing groundwater, and some new uranium deposition may occur at the same time. 3) The transitional zone (mineralization zone): The ratios of 230Th/234Ufor all rock samples being less than unity suggests that they have experienced deposition of uranium from groundwater within 3 x 105years, and the deposition may have continuous to the present time, and the 234U/238Uratios being partly than unity and partly less than unity indicates that uranium ore-formation is a rolling and multiple superimposed process. 4) The reduced zone: The 234U/238U ratios of most rock samples being greater than unity shows that the rocks have experienced uranium deposition for the last 106 years, the ratios of 230Th/234U for most samples being less than unity indicates uranium dissolution may have occurred for the last 3 x 10’ years, and the 234U/238U ratios being less than expected suggests that partial oxidation may have happened in the zone due to some oxygen-containing groundwater’s intrusion through some fractures.

The research was supported by China foundation of Nuclear sciences at Project No. Y7190R1801. 763

REFERENCES Ivanovich, M. & Harmon, R.S. 1982. Uranium series disequilibrium: Application to environmental problems. Oxford: Oxiford University Press. Liu, J.H. 2000. Studies of ore hydrogeochemistry of sandstonetype uranium deposits in the Yili Basin and the ZhungeerBasin, NW-China. Dissertation of China University of Geosciences: 1-97. MacKenzie, A.B., Scott, R.D., Linsalata, P. & Miekeley, N. 1992. Natural decay series studies of the redox front system in the Pocos de Caldas uranium mineralization. J. Geochem. EXplol: 4.5: 289-322. Osmond. J.K. Cowart, J.B. & Ivanovich, M. 1983. Uranium isotopic disequilibrium in groundwater as an indicator of anomalies. In(. J. Appl. Radial. hot. 34: 283-308. Scott, R.D., MacKenzie, A.B. & Alexander, W.R. 1992. The interpretation of 2’sU-234U-230Th--’26Ra disequilibria produced by rock-water interactions. J. Geochem. Explol: 45: 323-343. Shi, W.J. 1990. Hydrogeochetnistry of uranium. Beijing: Atomic Energy Press.

764

Wafer-Rocklnferacfion2001, Cidu (ed.,),02001 Swets & Zeitfinger,Lisse, ISBN 90 2651 824 2

Thermodynamic framework of the contact metamorphism around the Kakkonda granite in a active geothermal field, northeast Japan N.Takeno, H.Muraoka, TSawaki & M.Sasaki ~ u t i o n uInstitute l of Advanced I n d ~ s t r ~ Science al and ~ e c ~ n ~Japan l ~ g ~ ,

ABSTRACT: Contact metamorphism, up to the andalusite isograd, around a Quaternary granite complex was examined by chemical equilibrium calculation for rock samples from the WD-la well in the Kakkonda geothermal area, northeast Japan. We have calculated metamorphic mineral assemblages for the bulk chemical composition of five core samples from 300" to 800°C and from 521 to 851 bar using the Gibbs free energy minimization method. Stability temperature ranges of the metamorphic minerals and their possible reactions for isograd are obtained, and the calculated results are consistent with those observed in the well. The estimated peak metamorphic temperature is about 200°C higher than the present temperature.

1 INTRODUCTION Metamorphic minerals have been found in the deeper parts of some geothermal areas in the world, such as Larderello (Bertini et al. 1995) and Salton Sea (Helgeson 1968). The Kakkonda geothermal area in northeast Japan is one of these areas (Fig. 1). A contact metamorphic aureole was found around the Kakkonda granite during the company's exploration of the deep geothermal reservoir for the geothermal power plant (Kato et al. 1993). The New Energy and Industrial Technology Development Organization has carried out the drilling of the WD-1a exploration well as part of the Deep Geothermal Resource Survey Program of the Kakkonda geothermal area. This well has intersected the contact metamorphic aureole and Quaternary granite complex in which the bottom hole temperature reached 500°C. We have tried to estimate the thermodynamic framework for this contact metamorphism by equilibrium calculation using the Gibbs free energy minimization method, and discuss the evolution of the thermal structure of this geothermal area.

Below this contact, granodiorite or tonalite (Kakkonda granite) is encountered. The Kakkonda granite is a deep-seated Quaternary granite (Kanisawa et al. 1994) and does not outcrop. The upper part of the Tertiary formation mainly consists of dacite tuff and shale, while the lower part consists of andesite, andesite tuff and shale. The Pre-Tertiary formation consists of andesite, slate and sandstone. Contact metamorphism and hydrothermal (geothermal) alteration are recognized in the well. Biotite, cordierite, anthophyllite, and andalusite isograds are recognized (Kato et al. 1993, Fig. 2). Doi et al. (1998) summarized the metamorphic or hydrothermal alteration minerals in this geothermal system: 1) metamorphic minerals such as biotite, cordierite, anthophyllite, andalusite, hastingsite, -

,

'

II

m:a iQMatsul

wa Mt.

'BKakkonda

2 GEOLOGIC SETTINGS

ocs,

The Kakkonda geothermal area is located in anticlinal Tertiary formations, outcropping along the Kakkonda river, and surrounded by Quaternary volcanic rocks. Tertiary Miocene formations are distributed from the surface to 2660 m depth in WDla. Below this depth, Pre-Tertiary formations are found before the granite contact at 2860 m depth.

,

AMt. Komagatake

5

,

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'hlorioka

~

\'

Figure I. Locality map. Open circle: hot spring, open square: geothermal power plant, open triangle: active voicano.

765

clinopyroxene, curnmingtonite, garnet, orthopyroxene, and spinel, 2) common minerals for lowgrade metamorphism and hydrothermal alteration like chlorite, illite (muscovite), epidote, tremolite, and magnetite, 3) hydrothermal alteration minerals such as zeolite, smectite, prehnite, anhydrite, gypsum, calcite, hematite, and sulfide minerals.

Table 1. Sample depth and estimated pressures.

A B D E

depth (m) 1447 1696 2276 2688 2843

estimated pressure (bar) 52 1 580 718 815 85 1

3 METHOD Cuttings and core samples recovered from WD-la were subjected to chemical analysis. We used five core samples as a starting composition for the chemical equilibrium calculation. The chemical composition of these samples as well as cuttings are shown in Figure 3. These five samples cover the compositional variation of the cuttings from 1000 to 2860 m depth (granite contact). For the system composition given by rock and excess water (Na-KCa-Mg-Fe-Al-Si-H-0 system), we calculated the equilibrium mineral composition in 1°C steps from

Figure 3. Chemical composition of rock samples and cuttings from WD- 1a.

300" to 800°C using the Gibbs free energy minimization method after Eriksson (1975). We ignored the possibility of partial melting in the high temperature region. Some cuttings contain carbonaceous matter, so we also carried out the same calculations for the systems with graphite (Na-K-Ca-Mg-Fe-AlSi-C-H-0 system). Pressure is estimated from the measured density data and overburden erosion (0.9 km; Muramatsu 1987) (Table 1). To examine the pressure effect, we carried out calculations assuming various pressure values: minimum pressure (521 bar; estimated pressure of A), estimated pressure, and maximum pressure (85 1 bar; estimated pressure of E). Thermodynamic data for the minerals and gases are from Holland & Powell (1998). Fugacity data for the gases are tabulated values calculated from SUPERFLUID (Belonoshko et al. 1992). The mixing properties for solid-solution minerals are after Holland & Powell (1998). Ideal mixing for gases is assumed.

4 RESULTS AND DISCUSSION Figure 2. Geology and logging data of WD-la (Doi et al. 1998 and Muraoka et al. 1998) and sample points (from A to E). Horizontal shaded bars in the right column are minimum isograd temperatures.

One Of the is presented in Figure as an example. Over twenty-six results, the Same type was obtained for five samples under two or three pres-

766

sures and with/without graphite, These results are compiled in Figures 5 and 6. Horizontal bars in Figure 5 show the stability temperature range of major metamorphic minerals for samples A to E between 300" and 700°C. Compare the stable minerals and their prograde metamorphic sequences with the observed minerals in Figure 2. All minerals in Figure 2 agree with the calculated results in Figure 5. Although talc was not studied in previous work, the calculation shows that its existence is reasonable from the thermodynamic point of view. Garnet, tremolite and orthopyroxene were found in other company wells (Kato et al. 1993) and they are also confirmed in these calculations. It is also noticed that the biotite stability temperature range is dependent on the bulk rock composition. This requires caution in interpreting the biotite isograd, because many kinds of reaction are involved. In the case of graphite coexistence, iron oxide minerals and talc disappear (Fig. 6). This disagreement with the field occurrences indicates that graphite does not equilibrate with the fluid. Figure 6 also shows that the low fugacity of water caused by the mixture of CO, and CH, considerably lowers the stability range of biotite, anthophyllite and orthopyroxene compared with Figure 5. Sawaki et al. (1998) found a significant CO, content in the fluid inclusions of some minerals in the deeper part of WD-la. This may account for the rather shallow distribution (lower temperature) of anthophyllite in Figure 2 compared with the calculated result in Fig-

3 3

Temperature ('C) 590 690

490

790

epidotel garnet I'

II A E

biotite

c li

t remolite

A

R:l

:

.,

I/

1

1 - "

cordierite

AI

anthophyllite

B

andalusite

E

J l

I

c Is

l

orthopyroxene

Figure 5. Temperature-stability relation of major metamorphic minerals without graphite. Horizontal bars indicate the stability temperature range of minerals for pressures: 521 bar (dotted line), estimated pressure (thin line), 85 1 bar (broken line).

ure 5. If it is assumed that the low end of the stability range in Figure 5 indicates the temperature at which the mineral begins to appear in the geothermal

WD-1 a 2843 m 800 I

750

I

__ __

c:. A ::

I

I

t remolite

700

650

02 2

600

U

E

cordierite

550

E 500

anthophyllik

450

400

+--

BI

c

wi1

or t hopyroxene D'

350

forsterite

300

E

Figure 6. Temperature-stability relation of major metamorphic minerals with graphite. Horizontal bars indicate the stability temperature range of minerals for pressures: 52 1 bar (dotted line), estimated pressure (thin line), 85 1 bar (broken line). Epidote, garnet ,talc, andalusite and iron oxides are absent.

Figure 4. One example of the results. Sample E for 851 bar without graphite. The crowded part of inset is enlarged. A broken line indicates end-member composition of solid-solution mineral.

767

Muramatsu, Y. 1987. Distributions, paragenesis and fluid inclusions of hydrothermal minerals in the Kakkonda geothermal field, Northeast Japan.* J. Japan. Assoc. Petrol. Min. Econ. Geol. 82: 216-229. Muraoka, H., Uchida, T., Sasada, M., Yagi, M., Akaku, K., Sasaki, M., Yasukawa, K., Miyazaki, S., Doi, N., Saito, S., Sato, K. & S. Tanaka 1998. Deep geothermal resources survey program: igneous metamorphic and hydrothermal processes in a well encountering 500°C at 3729 m depth, Kakkonda, Japan. Geothermics 27(5/6): 507-534. Sawaki, T., Sasaki, M., Fujimoto, K., Takeno, N., Sanada, K. & T. Maeda 1998. Corundum and zincian spine1 from the Kakkonda geothermal system, Iwate Prefecture.* Proc. of annual meeting of Mineralogical Society of Japan: 173. Fukuoka, Japan. * in Japanese with English abstract

field, we can estimate the lowest isograd temperature. For this purpose, we used biotite, cordierite, anthophyllite and andalusite for the estimated pressure, because pressure variation does not lead to serious error. However, the influence of the H,O fraction in the fluid phase is important for the stability of some minerals such as anthophyllite, and therefore it is taken into account as a horizontal bar in Figure 2. A thermal gradient curve drawn through these bars would indicate a 200°C temperature decrease at 2800 m depth from the time of the metamorphic peak to the present day.

5 CONCLUSIONS Our thermodynamic framework of contact metamorphism based on the Gibbs free energy minimization method satisfactorily reproduces the metamorphic mineral assemblage and their prograde metamorphic sequences. The biotite isograd should be carefully interpreted for its dependence on the bulk rock chemical composition. The stability of anthophyllite is highly dependent on the H,O ratio of the fluid phase. The present temperature at 2800 m depth is at least 200°C lower than the metamorphic peak. REFERENCES Belonoshko, A.B., Shi, P. & S.K. Saxena 1992. SUPERFLUID: A FORTRAN-77 program for calculation of Gibbs free energy and volume of C-H-0-N-S-Ar mixtures. Computers & Geosciences 18: 1267-1269. Bertini, G., Cappetti, G., Dini, I. & F. Lovari 1995. Deep drilling results and updating of geothermal knowledge on the Monte Amiata area. In Proc. World Geothermal Congress '95: 1283-1286. Florence, Italy. Doi, N., Kato, O., Ikeuchi, K., Komatsu, R., Miyazaki, S., Akaku, K. & T. Uchida 1998. Genesis of the plutonic-hydrothermal system around Quaternary granite in the Kakkonda geothermal system, Japan. Geothermics 27(5/6): 663-690. Eriksson, G. 1975. Thermodynamic studies of high ternperature equilibria. XI 1. SOLGASMIX, a computer program for calculation of equilibrium composition in multiphase systems. Chemica Scriptu 8: 100-103. Helgeson, H.C. 1968. Geologic and thermodynamic characteristics of the Salton Sea geotheremal system. Am. Jour. Sci. 266: 129- 166. Holland, T.J.B. & R. Powell 1998. An internally consistent thermodynamic data set for phases of petrological interest. J. Metamorphic Geol. 16: 309-343. Kanisawa, S., Doi, N., Kato, 0. & K. Ishikawa 1994. Quaternary Kakkonda granite underlying the Kakkonda geothermal field, Northeast Japan.* J. Japan. Assoc. Petrol. Min. Econ. Geol. 89: 390-407. Kato, O., Doi, N.& Y. Muramatsu 1993. Neo-granitic pluton and geothermal reservoir at the Kakkonda geothermal field, Iwate prefecture, Japan.* Jour. Geotherm. Res. Soc. Japan 15: 41-57.

768

Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitiinger, Lisse, ISBN 90 2651 824 2

Interaction of fluid inclusions with dislocations in quartz N .A .Tchepkaia Far East Geological Institute Rus. Acad of Sci., Vladivostok, Russia

Z.A.Kotelnikova Institute of Lithosphere Rus. Acad. of Sci., Moscow, Russia

ABSTRACT: The relationship between fluid inclusions and the dislocational structure of quartz has been determined using transmitted light and electron microscopy methods under selective etching and chemical enhancement. The results testify that all primary inclusions of 1-5 pm size have, at least, an internal link to one dislocation. Inclusions of 10-20 pm are characterized by a bunch of parallel or fan-like dislocations dispersing fiom the cavity. Experimental modeling of the hydrothermal influence on fluid inclusions indicates that their chemical composition and density can be modified. The dislocation form is also changed during plastic deformation of the mineral-host. 1 INTRODUCTION

Aqueous fluid inclusions have been used to interpret the volumetric data (PVT) of fluids related to various geological processes. Previous studies (Bodnar & Sterner 1984, Sterner & Bodnar 1985) have proven that the fluid in inclusions has the same composition and density as the parent solutions. However, fluid inclusions in minerals are often connected with dislocations, the contents of which provide very important information about crystal growth and its following evolution. Furthermore, dislocations represent the leaking-out channel for fluid in inclusions. Although the occurrence of the association of fluid inclusions and dislocations is well known (Ser et al. 1980, Wilkins, Bird 1980), we presently lack a thorough understanding of the character of this relationship. In this paper we describe in detail the connection between fluid inclusions and dislocations in synthetic and natural quartz. Transmitted light and electron microscopy methods under selective etching and chemical enhancement were used in this study. In order to understand the behavior of the fluid inclusions which are joined together with dislocations in a high temperature, high pressure environment, we assessed the hydrothermal influence on the mineral-host and investigated the fluid inclusion properties. In addition, the dislocation form was studied before and after mineral-host deformation.

2 EXPERIMENTAL PROCEDURE For trapping fluid inclusions we used the experimental procedures described in detail by

Sterner & Bodnar (1984). Therefore, we only give a summary here. Quartz cores approximately 2x3 x 15 mm were cut from non-inclusioned natural quartz (from the mineralized fissure of Pamir, Aldan and Anabar shield, Polar Ural) and fractured by a thermal-shock procedure. Synthesis experiments were carried out by the thermogradient method at 350-400 OC in NaCI, KCI, NaN03 and NaOH solutions of different compositions. The experiments studied the influence of elevated temperature and pressure upon fluid inclusions under hydrothermal conditions. The samples of quartz with fluid inclusions were loaded into an autoclave and pure water or a solution of NaCl (10 or 20 wt.%) was added. The experiments were run at 5OO0C and 1600 bar pressure. In order to avoid dissolution of the quartz samples, 0.03-0.07 g amorphic SiO2 was added to the autoclave. Fluid inclusions were examined in quartz discs, polished on both sides, using microthermometry techniques. Microthennometric measurements have an estimated accuracy f 0.5OC. In order to reveal dislocations we used a solution of HF, and a mixture of HF and HNO3 as etchants (Ball & White, 1977). The procedure suggested by Kotelnikova & Sonushkin (1994) was applied for chemical enhancement of the dislocation. The quartz samples were placed into an autoclave and treated in a solution containing 3-5 parts of HF and 93-97 parts of distilled water at 2OOOC for approximately 2-6 hours. The samples investigated by transmitted light and electron microscopy methods were prepared using ion thinning. Electron microscopy was employed only for studying very minor fluid inclusions ranging less 769

than 0.5 pm. The method resolution is 10" cm-2. The disadvantage of this method is the difficulty of preparing samples. Chemical enhancement was the major method for studying fluid inclusions. This method allowed the observation of dislocations near fluid inclusion cavities larger than 10 pm in diameter, giving reasonably good results. The resolution for the chemical enhancement method is 107 cm-'. The resolution for selective etching is 108 cm-2. Only very well polished quartz samples can be utilized for investigation using this method. A special etching procedure is necessary for every sample.

3 RESULTS The data obtained indicate that all primary fluid inclusions located in the overgrowths on the original quartz zone are connected to at least one dislocation. We found no fluid inclusions, which were disconnected from dislocations. The primary fluid inclusions and dislocations associated with them are shown in figure 1. The fluid inclusions are located in the places where the dislocations cross. Arrows on figure 1 mark these places. The etching procedure was selected to ensure that etching occurred only along dislocation lines. Many primary inclusions associated with subgrain boundaries in the crystals are connected by plane defects, i.e. screw dislocation boundaries. Moreover the number of dislocations contacting inclusions strongly depends on the fluid inclusion size. Small inclusions, 1 to 5 pm in diameter, are associated with one dislocation, while inclusions ranging from 10-20 pm are characterized by a cluster of parallel or fan-like dislocations dispersing fron

Figure 1. The relationship between the dislocation lines and fluid inclusion cavities in a quartz crystal from a mineralized fissure (the Polar Ural). Data was obtained by using the diffraction electron microscopy method. Fluid inclusion cavities are marked by arrows.

Figure2. Secondary fluid inclusions in quartz from Pamir with a solution-bearing channel branching off the cavity. Data was obtained by selective etching method.

the cavity. The dislocation density near large fluid inclusions (50 to 100 pm in diameter) reaches 5" cmm2.The density of dislocations on the inclusion wall is higher then on the rest of the crystal. The analysis of secondary fluid inclusions formed by healing fractures indicates that, in this case, practically all the fluid inclusion cavities are linked with rows or screens of dislocations. According to the chemical enhancement data, the dislocation density in secondary fluid inclusion cavities is increased, reaching 107cm-'. The results of our study agree satisfactorily with the observations of Kotelnikova & Sonushkin (1994). Dislocations very often connect many secondary fluid inclusions with each other. Our data indicate that the last generation of secondary fluid inclusions is often connected by thin channels which are sometimes filled with parent solution. We observed vapor bubbles and salt crystals, which were identified as halite, within such channels. An example of a secondary vapor-rich inclusion associated with a solution-bearing channel is shown in figure 2. The formation of these channels is likely to occur during etching of the dislocation and trapping of the solution, or during dendrite crystal growth. Our data agree closely with those of Wilkins (1986). In order to demonstrate the hermetical property of fluid inclusions connected with dislocations, a number of experimental runs modeling of the hydrothermal influence upon the mineral-host were conducted. Results are presented here of the runs performed a quartz crystal which was grown in 4 wt.% NaCl solution. The temperature of homogenization and the composition of the fluid in inclusions were determined before and after runs. The fluid inclusions had the following parameters 770

Table 1. Results microthermometry of fluid inclusions after experiments. Run number

1 2 3 4 5 6

Run conditions Temperature, Pressure, Solution in run "C bar 500 1600 Pure water 500 1600 Pure water 500 1600 10wt.%NaCl 500 1600 10wt.%NaC1 500 1600 20wt.%NaCl 500 1600 20wt.%NaC1

Run duration, days 14 32 14 32 14 32

Termometry Th (L-P), "C Salinity Number of measured (average) (average) inclusions 375 4.1 5 375,5 4.9 7 3SO 5 13 380 695 17 377,5 692 13 316,5 12,45 10

homogenisation temperature.

before the experiments: the average temperature of homogenization (Th) was 375.5 *C and the salinity of the fluid was 4 U$.% NaC1. In all quartz samples newly formed (during the run) secondary fluid inclusions were found. Their density and composition exactly matched the fluid parameters of the run. Table 1 summarizes the results of the microthermometry studies of the fluid inclusions after the runs. The microthermometry results of newly formed inclusions are not represented in this table. Most of the fluid inclusions investigated did not maintain their original characteristics, the exception being the run with pure water (numbers 1 and 2). After the other runs the temperature of homogenization of all measured fluid inclusions was increased. However, the deviations of the homogenization temperature of fluid inclusions after runs number 5 and 6 were small, not exceeding 2 "C. This characteristic changing of homogenisation temperature indicates that there was not any decrepitation of fluid inclusions during the runs. If this had been the case, the homogenisation temperature would have been decreased. It is likely that fluid leaked via dislocation or microtracks. According to the composition of fluid inclusions after the runs, we can distinguish two groups of inclusions: those maintaining their initial composition and those that have changed. The maximum concentration of change we observed was always connected with dislocations in needle-shaped inclusions. The calculated interior pressure during heating to homogenisation in fluid inclusions before runs was 820 bar. During runs interior pressure in the inclusions was about 1350 bar and pressure in the autoclave was 1600 bar; therefore, all inclusions were maintained without decrepitation. The dislocation forms were also investigated before and after crystal deformation. The form of the investigated dislocation is shown in Figure 3: 3A indicates the form of dislocations before quartz

plastic deformation and 3B - the dislocation form after quartz plastic deformation. The data obtained show that a11 dislocations in undistorted quartz were found in a straight line (Fig. 3A). After impacting high temperature and pressure on the quartz, the form of dislocations was changed; some bent into loop-like shapes. The dislocation loops connected with plastic deformations imposed upon the fluid inclusion cavities are indicated in Figure 3B. Sometimes, we observed a dislocation rosette. The maximum density of dislocations was 21° cm-*. Thus, the dislocation form allows us to identiij the nature of the mineral containing fluid inclusions and, therefore, defines the fluid inclusion applicability for PVT determination. 4 DISCUSSION

It is likely that the relationship of the fluid inclusion cavities to the dislocation, as well as to the increase

Figure 3. The form of dislocations outgoing from the fluid inclusion cavities. Data obtained by the chemical enhancement method. A - form of dislocations before plastic deformation of quartz. B - form of dislocations after plastic deformation of quartz.

771

in density dislocation at inclusion walls, is very important during the influence of elevated temperature and pressure on the mineral-host. The increase of temperature causes a rise of pressure in fluid inclusions. The structure of the mineral-host is also changed. First of all, a coordination of lattice defects occurs via the transformation of dislocations as the most mobile of elements of the structure of mineral. If the dislocation density in the inclusion wall is sufficiently high (approximately 107-1010 cmm2),the fluid inclusion does not re-equilibrate or leak during heating and over-pressure. If the density of dislocations in fluid cavities does not exceed 104 cm-2, microcleavage structure is likely to appear to neutralize tension. Increased stress causes microcrack formation. Like microcleavage, microcracks are channels for re-equilibration, stretching or leaking the fluid inclusions. In summary, the results of this study are in reasonable agreement with previous experimental data (Wilkins, 1986; Kotelnikova & Sonushkin, 1994) indicating that the re-equilibration of fluid inclusions is a result of host mineral plastic deformation.

geological processes are the primary fluid inclusions associated with single dislocation.

Ball A. & White S. 1977. An etching technique for revealing dislocation structures in deformed quartz grains Tectonophysics, 37: 9-14. Bodnar R.J. & Sterner S.M. 1985. Synthetic fluid inclusions in natural quartz. 11. Application to PVT studies. Geochim. Cosmochim. Acta, 49: 1855-1859. Sterner S.M. & R.J. Bodnar. 1984. Synthetic fluid inclusions in natural quartz I: Compositional types synthesized and applications in experimental geochemistry. Geochim. Cosmochim. Acta, 48: 2659-2668. Kotelnikova Z.A. & V.E. Sonjushkin. 1994. Dislocational structure of quartz and the problem of hermetical tightness of the fluid inclusion cavities. Zapiski Rus. MineKSociety, 3: pp.9-18. Ser F.P., Bideau J.P., Clastre J. & A.Zarka. 1980. Etude des defauts de croissance dans des monocristaux naturels de quartz. J.App1. C y s t . 13: 50-57. Wilkins R. W.T. 1986. The relationship between dislocations and fluid inclusions in minerals. Res.Rev. CSIRO div. Minel:Geochem, Camberra. 1 13-134. Wilkins R.W.T. & J.R.Bird 1980. Characterisation of fracture surfaces in fluorite by etching and proton irradiation. Lithos. 13: 11-18.

CONCLUSIONS Using transmitted light and electron microscopy methods with selective etching and chemical enhancement, the character of the relationship between fluid inclusions and dislocations in quartz was determined. The data obtained indicate that primary inclusions located in the growth zones of crystals are always connected to at least one dislocation. Moreover, larger fluid inclusions are associated with higher amounts of dislocation then smaller ones. Very often dislocations connect the fluid inclusion cavities with subgrain boundaries of the crystal. The dislocation are etched and enlarged in volume under the influence of elevated temperature and pressure upon the mineral-host. Therefore the dislocations are the channels by which fluid leaves the inclusions. The density of dislocations is increased near the fluid inclusion cavities. The amount of dislocation at the fluid inclusions’ wall is strongly influenced on the behavior of fluid inclusions during the impacting of high temperature and pressure on the mineral-host. The experimental results reported here show that the most hermetical inclusions and, therefore, the most suitable inclusions for accurate interpretation of volumetric (PVT) data for fluid in different 772