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
'Mudrock' is a general term for clastic fine-grained sediments and sedimentary rocks that contain mixtures of fine sand-, silt-, and clay-size grains, with >35% of the particles being <63 pm diameter (Grainger 1984). Mudrocks have varying degrees of compaction, consolidation and induration, ranging from soft plastic layered clays to hard cemented argillaceous rocks. Harder mudrocks may fracture along bedding planes andor in irregular joints. Mechanical strength depends on the proportions of clays, silt- and sand-size particles, the layering and sorting of the grains, and the degree of cementation by diagenetic calcite, quartz and clays. Plastic mudrocks accommodate stress by deforming and, if they do fracture, by self-healing over time. For engineering purposes, mudrocks can be classified according to compressive strength: non-durable 0.63.6 MPa and durable B3.6 MPa (Grainger, 1984). Diagenetic mudrocks are distinguished from metamorphic equivalents by an upper limit on strength of 100 MPa. Slaty fabric may already be a compaction feature below this limit. An extensive review of theory and experimental evidence for the hydraulic and solute transport properties of mudrocks is contained in Horseman et al. (1996). Consolidated mudrocks have permeabilities that are very low: 10-" to
3
different thermodynamic properties from those of the ‘free’ pore water.
There has also been laboratory research and limited field investigation of the interactions of mudrocks and pure clays (e.g. bentonite) with trace metals, much of it investigating the sorption of radionuclides. This important topic is outside the scope of this paper, but it is worth noting that quite detailed models for exchange of trace cations in competition with major cations have been developed, discriminating between different types of exchange sites on clay minerals. For example, a model for exchange of Cs has been constructed and has been tested against data from laboratory experiments with some of the mudrock formations that are described below (Bradbury & Baeyens 2000). The second part of this paper introduces a set of papers on studies of the Jurassic Opalinus Clay formation from the international cooperation project based at the Mont Terri Underground Research Laboratory (URL) in Switzerland. The Mont Terri Project is also investigating other relevant mudrock characteristics including in situ pore water pressures, osmotic pressures, and diffusion rates by tracer testing. An illustrated overview of the Mont Terri project and of the partners involved can be found at http://www.mont-terri.ch.
Figure 1. Sketch diagram of the structure of pores and pore water in mudrocks. Water-rock reactions in mudrocks are of potential importance because mudrocks occur over substantial parts of the landmass. They constitute, for example, about 10-20% of the surface rocks of Europe. Mudrocks account for significant proportions of the total groundwater in the first few hundreds of metres of the subsurface although that pore water only exchanges very slowly with adjacent groundwaters. At greater depths, pore waters being expelled from compacting shales have substantial influences on the compositions of formation waters in adjacent higher permeability formations. The general pattern found in overpressured immature sedimentary basins is that the expelled pore waters have a diluting effect on saline formation waters, though the detail is often more complex and probably is dependent on the depositional facies, stage of compaction and many other factors (Chilingarian et al. 1994). This paper reviews previous geochemical studies of pore waters in relatively shallow mudrocks (i.e. -600 m depth). Some of these studies have been carried out to assess the feasibility of safe long-term containment of radioactive wastes in repositories constructed in mudrocks. In this context, water-rock reactions are important in three ways: firstly, they provide the boundary conditions for the behaviour of repository backfill and buffer materials and for the degradation of the waste form and waste-package components; secondly, they control the geochemical speciation of leaking waste solutes and thus their mobilities and sorptive retardation; thirdly, they modify the mudrocks themselves over long periods and thus influence the evolution of physical and geochemical properties.
2 PREvIouswoRK 2.1 Boom Clay at Mol, Belgium The first investigation of mudrocks by in situ measurements was the ‘Hades’ project at the URZ, constructed in 1983 at Mol in Belgium (Beaufays et al. 1994). This is located at 223 m depth in the Tertiary Boom Clay, a plastic clay-silt formation about 100 m thick with total porosity -0.4 and hydraulic conductivity -3x1 0-l2 ms-I. The pore water composition in the Boom Clay is Na-HC03 with 18 mg/l chloride. The exchangeable cation population of the clays is dominated by Mg and Na, and is close to that predicted from the pore water composition, thus confirming that the exchange reaction is at equilibrium (Baeyens et al. 1985). The EC-fimded ‘Projet Archimkde-argiles’ developed methods for characterising pore water and minerals in the Boom Clay and used geochemical modelling to understand the processes that regulate the in situ chemical conditions (Griffault et al. 1996). pH of water that had seeped into a piezometer in the clay was measured with an optical sensor and by conventional electrode, giving consistent values at 8.2W0.05 (Beaucaire et al. 2000). Water in the underlying aquifer is also fresh at this location but becomes brackish elsewhere as 4
the aquifer deepens. Although Br/Cl in the aquifer has a marine value, it is higher in the clay pore waters due to an increase of Br apparently leached from the formation. Calcium and magnesium are both depleted in the pore waters relative to their marine ratios to sodium, indicating that Ca and Mg have displaced Na from exchange sites on the clay as pore water compositions have transformed from marine to fresh. A geochemical model of the evolution and present state of water-rock reaction was constructed in which fresh ‘clay’ water was mixed with brackish water and then re-equilibrated with clay (Beaucaire et al. 2000). It suggested that most of the major elements are controlled by diagenetic or secondary mineral phases (calcite, dolomite, kaolinite and other clay minerals, Mg-aluminosilicate) and superimposed on this is the rapid re-equilibration of the exchange of Na, K, Ca, Mg, etc between pore water and clays. Apparent supersaturation of the carbonate phases was observed though whether this was real or was an artefact of CO2 loss was not clear. The geochemical model also showed that the evolution of pore water pH is dependent on the concentration of Cl; this relationship is evident in the underlying aquifer in which pH decreases with increasing salinity (Beaucaire et al. 2000).
and oxidation of pyrite are thought to be significant factors in these mudrocks (Bath 2001). The Tertiary mudrock is at a coastal site and shows the effect of sea level changes on chloride concentrations. Chloride and isotope variations were consistent with a diffusion model with timevarying boundary concentrations(Falck et al. 1990).
2.3 Cretaceous marl at Wellenberg, Switzerland Geochemical investigations of the slightly metamorphosed marl of the Palfris Formation (Cretaceous) at Wellenberg in Switzerland were reported by Pearson & Scholtis (1995). Pore water compositions were inferred from water samples taken from boreholes. Chloride concentrations increase with depth to about 12,000 mg/l at >400 m (Nagra 1997). Concentrations of reactive solutes have greater uncertainties due to contamination by drilling fluids. Geochemical modelling was used to get the best estimate for in situ water compositions in the rock matrix. Cation and anion concentrations were constrained by equilibria with calcite, dolomite and fluorite and by ion exchange equilibria with clays for Na, K, Mg and Sr. The chloride concentration constrained the Pc02 at about 10-’ bar. pH is uncertain due to the absence of reliable alkalinity data, and sulphate measurements were considered unrepresentative of in situ pore water because of the suspected effect of pyrite oxidation.
2.2 Jurassic and Tertiary mudrocks, southern and eastern England Cored research boreholes were drilled to more than 300 m depth in the Cretaceous-Jurassic sedimentary sequence at Harwell in southern England in the early-1980s. Hydrogeological testing and sampling was carried out and pore waters were extracted from drillcore samples by squeezing (Brightman et al. 1985). Chemical and isotopic analyses of pore waters produced vertical profiles of salinity and stable isotopic compositions. These were interpreted in terms of the history of the basinal sequence and diffusive mixing between pore waters and groundwaters in adjacent aquifers (Bath 200 1). Pore water profiles were subsequently obtained from relatively shallow mudrocks of Middle Jurassic and Tertiary ages at sites in southern and eastern England that were being investigated as potential sites for storage of low-level radioactive wastes. Chloride concentrations and stable isotope ratios show different degrees of mixing between syndepositional or very old pore water and younger water from the surface or adjacent aquifers (Bath et al. 1988, Bath, 1994, 1995). Various factors have influenced pore water compositions: interbedded aquifers, geological history, erosion and faulting,
2.4 Indurated Jurassic mudrocks at Tournemire, southern France Boreholes were drilled from a disused railway tunnel through indurated Jurassic (Toarcian) mudrocks at Tournemire in southern France where 200 m of mudrocks is overlain by up to 270 m of limestones. Tectonisation has created a network of fissures and fractures in the mudrock through which there are low water flows. Pore water contents (3-5 wt%) were insufficient for extraction by squeezing, and a typical composition of pore waters has been estimated from a combination of leaching experiments and geochemical modelling @e Windt et al. 1998). This result was compared with the composition of waters flowing into packered intervals in boreholes, indicating that these flows have the chemical character of pore water in the mudrock rather than of groundwaters in the adjacent limestone aquifers. This suggests that fracture water and pore water exchange by difision and that their composition is buffered by reactions with the rock matrix.
5
Inflow waters were in equilibrium with calcite, dolomite, siderite, fluorite and chalcedony. Typical selectivity coefficients for Na, Ca and Mg exchange between waters and cation populations of clays were also measured. Results suggested that both faster (e.g. exchange) and slower (e.g. feldspar dissolution) equilibria have r6les in controlling the in situ hydrochemistry of mudrocks (De Windt et al. 1999).
Compositions of pore waters in Plio-Pleistocene clay-rich sediments from various sites throughout Italy, mostly at the onshore margins of marine basins, were reported by Fontanive et al. (1993). Clay minerals were dominantly kaolinite and illite, with subsidiary smectite in some samples; water contents ranged from 11 to 35 wto/o. Pore waters were extracted by squeezing. pH values were in the range 6.8-8.8, and salinities varied with C1 up to 13500 mg/l, showing that the original pore waters in marine sediments have been diluted. Dilute pore waters in lacustrine sediments have Ca-Mg-SO4 compositions with up to 4000 mg/l Sod. NdC1 and Mg/Cl ratios have increased relative to seawater ratios in some of the diluted marine pore waters due to cation exchange during the dilutiodmixing process.
of the Holocene about 10 ka ago and has produced shallow pore waters with Na-Mg-SO4 compositions to 15 m depth, below which SO4 decreases and pore waters become reducing. Below a sandy layer at 13 m, the pore waters are thought to be essentially syndepositional with the till. The stable isotope (S2H) composition of pore water in the clay changes to an isotopically-lighter value that is characteristic of Pleistocene recharge. Chloride concentration changes are more complex, showing a pronounced maximum at 13 m and decreasing to minimum values in Pleistocene-age water below 36 m (Hendry et al. 2000). Chlorine-36 isotope data show that this Cl does not originate from underlying Cretaceous shales and that it is syndepositional with the till. Pore water evolution is interpreted by a diffusive mixing model and has been driven by concentration changes in the upper few metres due to oxidation (i.e. pyrite oxidation, carbonate dissolution, Ca/Na ion exchange). Pore waters are oversaturated with respect to calcite and dolomite (saturation index values >0.1), with calculated P C Ovalues ~ up to 10-’.7 bar; the oversaturation is attributed to the effect of dissolved organic carbon as an inhibitor of calcite nucleation. Reactions with other minerals in the unoxidised till-water system are thought to be so slow that they are not detected in the shape of the profiles (Hendry & Wassenaar 2000).
2.6 Clay aquitards in Flanders, Belgium and Saskatchewan, Canada
2.7 Experiments and modelling in compacted bentonite
Water-rock interactions in recharge through a Tertiary marine clay, 10-20 m thick and hydraulic conductivity 10-9ms-I, to a sandy aquifer in Belgium have led to a progressive evolution (‘freshening’) in the composition of the aquifer (Walraevens et al. 1998). The original Na-dominated clay has undergone exchange with infiltrating water, so that recharge composition has evolved over time from Na-Mg-HC03 to Ca-HC03. Geochemical modelling indicates that about 400 flushing cycles have occurred, probably exploiting preferential flow paths through the clay, removing Na+ from infiltration water and also raising the pH by H+ exchange. In this case, advective mass transport has occurred through a relatively unconsolidated mudrock, in which a predictable ‘chromatographic’ sequence of ion exchange reactions have occurred (cf. Appelo & Postma 1993). The movement of solutes and contaminants in mudrock has been investigated to 45 m depth in a clay-rich (predominantly smectite) glacial till aquitard in Saskatchewan, Canada (Hendry & Wassenaar 2000). Oxidising water is believed to have been infiltrating from the surface since the start
There is also quite a large body of experimental and modelling research into the reactions of bentonites with groundwaters, originating from the proposed use of bentonite to surround and seal packages of radioactive waste in repositories. There have been two main areas of research - transport and interactions of dissolved radionuclides in bentonite, and alteration of bentonite in the thermal and geochemical conditions prevailing around heatemitting high-level radioactive wastes (Alexander & McKinley, 1999). Reactions with trace mineral phases in bentonite (quartz, calcite, siderite, anhydrite and pyrite) could control the chemistry of pore waters. Geochemical modelling of water-bentonite reactions is therefore quite analogous to predictions of water-rock reactions in mudrocks. Modelling shows that, as groundwater solutes diffuse into the bentonite at normal temperature, reaction with calcite is the predominant buffer on both pH and the distribution of exchangeable cations (Bruno et al. 1999). If adjacent groundwater has abundant Ca and neutral pH, then these conditions are maintained in the bentonite and Ca progressively displaces Na from
2.5 Plio-Pleistocene clay-rich sediments, Italy
6
probably remained buried under marine conditions until regional domal uplift in two events (Cretaceous and mid-Tertiary) and subsequent thrusting and folding during the mid- to late-Tertiary period, culminating in erosion and the start of meteoric water ingress to the Jurassic rock sequence some time between 10 and 2 million years ago. The OPA that is intersected by the Mont Terri tunnel is contained within an overthrust-anticline so that the mudrock beds dip at between 30-55"; the tunnel section of the OPA is about 240 metres long and the maximum overburden above the tunnel is about 300 m. The mudrock formation is about 160 metres thick normal to bedding, and has been subdivided into five lithostratigraphic sub-units ('shaly', 'sandy' and 'carbonate-rich sandy'; see Fig. 3); it has typical uniaxial compressive strengths of 718 MPa (parallel to bedding) and 23-28 MPa (normal to bedding). The mineral composition of OPA is dominated by clay minerals (up to 70%) which are kaolinite and illite with lesser amounts of mixed layer illitehmectite and chlorite. O f the nonclays, calcite and quartz predominate with small proportions of K-feldspar, dolomite, albite, siderite and pyrite. Hydraulic tests of borehole intervals more than 5 m ftom the tunnel wall, i.e. beyond the excavation-damaged zone, gave hydraulic conductivities with a mean hydraulic conductivity of 5 x lO'I3 ms-'. A single fault zone has been mapped within the tunnel section. Hydrochemical and hydrogeological evidence suggests however that this zone does not have any anomalous groundwater movements associated with it. Tests over the fault zone gave hydraulic conductivity values within *3 x lO'I3 ms-' of the average value, i.e. there is no significant contrast. Total physical porosity of the OPA, as measured by water loss on heating, is mostly within the range 0.08-0.19.
clay exchange sites. On the other hand, if ambient groundwater is relatively depleted in Ca, then modelling shows that calcite dissolves and pH rises to rather high values around 10. Modelling also suggests that redox is controlled by pyrite and other sources of Fe" (siderite, Fe-montmorillonite). 3 MONT TERRI UNDERGROUND RESEARCH LABORATORY 3.1 General description Further investigations of many of the issues raised above have been carried out in the international collaborative research programme at the Mont Terri Underground Research Laboratory (Thury & Bossart 1999). A group of papers in this volume introduces some of these projects and summarises the results so far (Arcos et al., 2001; Blanc et al. 2001, Fernhdez et al. 2001, Gaboriau et al., 2001, Gaucher et al., 200 1, Pearson and Waber, 200 1, Rubel and Sonntag, 2001, Waber et al., 2001). The following section provides some background to the Mont Terri URL and a brief synthesis of these studies. Jurassic mudrocks at Mont Terri in the Jura Canton of Switzerland were first studied in 1989 jointly by the Swiss National Hydrological and Geological Survey and Nagra (the Swiss National Cooperative for the Disposal of Radioactive Waste), during construction of a reconnaissance tunnel for a motorway. Hydrogeological and geochemical studies (Gautschi et al. 1993) confirmed the potential of that location for an underground research laboratory in which a broad-based geoscientific study of mudrocks could be carried out, including investigations by in situ sampling, pore water extraction, and physico-chemical modelling of the mechanisms by which solutes move and of the water-rock interactions that occur in undisturbed and perturbed conditions. The Mont Terri Rock Laboratory is situated in a 3900 m long motorway tunnel in the Jura mountains of north-western Switzerland (Fig. 2). Construction commenced in 1989 with the excavation of a reconnaissance tunnel, within which the Rock Laboratory is situated (Thury & Bossart, 1999). The Opalinus Clay formation (OPA) is of Aalenian age (Middle Jurassic) and lies between Lower Dogger sandy limestones of Bajocian age and the Jurensis Mar1 formation of Toarcian (Lias) age. The OPA was deposited about 180 million years ago in shallow coastal marine conditions and was buried by further limestones and marls/clays over the following 50 million years or more, through to the early Cretaceous period when it was covered by about 800 metres of accumulated overburden. It
3.2 Characterisation of pore waters The first attempt to characterise the compositions of pore waters in the OPA was based on a combination of physicochemical testing and geochemical modelling (Bradbury & Baeyens 1998). The testing comprised aqueous leaching and selective extraction of exchangeable cations. Aqueous leaching provided C1 and SO4 data and selective extraction provided cation occupancies of exchange sites on clays. It was inferred from the leaching experiments that the prevailing mineral equilibria are with calcite, dolomite, fluorite and chalcedony. The measured data and dissolution and ion exchange equilibria provided suflicient constraints to allow an iterative estimation of likely pore water 7
Figure 2. Geological cross-section through Mont Terri, showing the motorway tunnel through the overthrust of the folded Jura and the location of the Underground Research Laboratory (from Thury & Bossart 1999). compositions. In situ pH was estimated as being in the range 6.8-7.0, corresponding to an assumed value for p C 0 ~of 1Oq2 bar. The experimental approach has various uncertainties, for example how much of the measured sulphate is an artefact of pyrite oxidation, and whether the relatively high water:rock ratios in leaching experiments give a reliable indication of mineral equilibria that control the in situ system at very low water:rock ratio. The distribution of pore water in the OPA between that bound on mineral surfaces and free water is a key factor in understanding water-rock reactions and in interpreting analytical data for the compositions of pore waters. The geochemical porosity has been deduced by a systematic comparison of chloride concentrations of squeezed pore waters with those calculated from leaching experiments and total porosity data (Pearson 1999). Pearson found that the geochemical porosity of OPA was 0.3-0.7 times the total porosity. This gives an idea of the proportion of the total pore water content in which the reactions of water and ionic solutes are reasonably close to those of bulk solutions. The paper by Gaboriau et al. (2001) describes the physical characterisation of the ‘bound’ water and confirms that it constitutes a large proportion, up to 6O%, of the total water in the OPA. Obtaining reliable data for pore water compositions has been a major objective of the Mont Terri project. Water samples have been collected in a number of purpose-drilled boreholes into which water flows very slowly (in the order of 10-1000 ml per month) from the OPA due to the high hydraulic gradients in the vicinity of the sampling intervals. It
has also been possible to squeeze pore water from drillcore samples: various laboratories have used different compaction loads in the range 60-500 MPa, yielding between 2 and 34% of the total water content (these proportions are larger if related to the ‘free’ water fraction only). Squeezed water data are discussed by Fernhdez et al. (2001), and are shown to be reasonably comparable with compositions of borehole water samples. Squeezing at loads up to 200 MPa has not substantially altered the major ions. However pH, alkalinity and Pc02 have been altered and cannot be readjusted reliably to in situ conditions. Therefore squeezed pore water data on their own are not adequate for modelling water-rock reactions that are sensitive to pH or alkalinity. 3.3 Origins and movement of solutes
The origins of non-reactive solutes and the time required for water and solutes to exchange between the OPA and the adjacent higher permeability zones are interpreted in the papers by Pearson & Waber (2001) and Riibel & Sonntag (2001). Pearson and Waber point out that the free or mobile solutes, i.e. C1 and possibly SO4 and C02, have a key role in the evolution of the water-rock system because they are ftxed independently of internal reactions. Chloride and sulphate have evolved by difisive exchange of pre-existing pore water and solutes of marine origin with adjacent fresher groundwaters since exposure of the sequence to meteoric water ingress. The profile of Cl concentrations through the OPA is the shape expected for such a diffusive evolution 8
Figure 3. Horizontal profile of chloride concentrations through the Opalinus Clay and adjacent formations at the Mont Terri Underground Research Laboratory. Water samples were collected in packered borehole intervals in the mudrock, in groundwater seepages in adjacent formations, and as pore waters squeezed from drillcore samples. (Fig. 3). A similar profile has been measured for 6'*0 and 6'H although there is not a 1:l relationship between C1 and 6'*0, suggesting that they have not diffised at the same rate. Evidence that diffusion has controlled water and solute movements in the OPA over at least several million years is strengthened by measurements of radiogenic helium and argon (Rubel & Sonntag 2001). The sampling and analytical method has minimised losses and contamination of the noble gases dissolved in pore waters. Dissolved 4He is in a steady state balance between in situ production within the pores and diffusive loss to younger groundwaters at the boundaries of the diffusion-controlled system. Modelling of the profile gives an in situ diffusion coefficient for He (3.5 x lO-" m2s-').
the natural processes and the sampling perturbations that control CO;! are adequately understood. Innovative experimental work to attempt to constrain the uncertainties in CO2 is in progress within the Mont Terri Project. An experimental approach to understanding claypore water interactions is necessary as a calibration of model simulations and as a test of the validity of model results. The limitations are, firstly, the necessity to use higher water:rock ratios and, secondly, the different timescales of experiments compared with those of in situ evolution. Results of batch experiments are presented by Blanc et al. (2001) in which OPA was reacted with a representative solution composition with Pc02 being controlled at 10-2.4bar. Observed results mostly confirm the model predictions, but there are noteworthy deviations for K and Fe. The experimental results challenge previous assumptions about the mineral phases that regulate the geochemical environment and the adequacy of thermodynamic data for them. Of particular interest is the conclusion that a Fe-containing phase other than siderite appears to control Fe concentration. Although ion exchange of cationic solutes is clearly one of the most important interactions between mudrocks and pore waters, the process itself has generally been represented rather
3.4 Geochemical modelling and experimental studies
Arcos et al. (2001) have used a 1D reactive transport model to show that, in addition to C1 and SO4, the distributions of Na, K, Ca, Mg and Sr through the OPA profile can be reasonably well simulated by including cation exchange and calcite, dolomite and celestite equilibria in the model, As with other geochemical modelling exercises, calculated values for pH and alkalinity differ significantly from measurements. This poses the question of whether 9
Organisation of water and solutes within a heterogeneous mudrock in which pores have a broad size distribution and different types of mineral surfaces exist; how this influences reactions and achievement of equilibrium at different scales, and how this can be addressed by water sampling strategies. How to measure or estimate the labile (faster reacting) chemical species that are easily disturbed by sampling: for example, Pco~,pH, carbonate alkalinity, redox-sensitive species Fe"/Fe'". What are the water-rock reactions, in addition to simple reversible sorption, that control the abundances and distributions of trace elements including heavy metals, radionuclides, other cationic and anionic species, and organic contaminants in mudrocks; what approaches are required for geochemical modelling? The nature and r6le of trace solid and dissolved organic compounds in chemical equilibria, and the existence and the potential movement of colloidal and microbiological species into and out of mudrock pores. In what circumstances does osmosis, caused by chemical gradients across mudrocks acting as semipermeable membranes, become a significant potential for water movement or overpressuring? Water and solute transport and mass budgets in mudrocks that are at an advanced state of diagenesis where pores are substantially cemented by mineral precipitates and fracturing becomes permanent and conductive so that a dual porosity medium evolves.
simplistically as involving a single type of exchange site. Gaucher et al. (2001) have measured cation exchange isotherms on clays from the OPA and have used these to test a multi-site ion exchange model. Comparisons between model predictions and observed distributions are significantly improved with a 2-site model for Na' versus K' and with a 4site model for Na' versus H+.
4 IMPLICATIONS FOR MUDROCK PROPERTIES AND DIAGENESIS Water-rock interactions in mudrocks and similar low permeability intergranular material are inherently difficult to study because of (i) low water:rock ratios, (ii) practical limitations on water sampling, (iii) the rapidity with which some of the chemical equilibria shiR in response to perturbations, and (iv) the complexity and importance of interactions with mineral surfaces in the microporous rock fabric. Nevertheless there is a growing body of in situ and experimental data and theoretical understanding. Some of the areas where significant progress has been made in recent years, and which have been reviewed in this paper, are: Pore water sampling - growing use of squeezing tests, confidence in data, and awareness of limitations and comparability of data from squeezing and solute leaching. Physical chemistry of pore water - coherence of experimental data with the electrostatic model for water-solute-surface interactions, identification of concept of geochemical porosity, estimation of the proportions of bound and free water in pores, and the potentially small amounts of free water in highly compacted mudrocks. Transport of water and solutes in mudrocks confirmation in diverse site conditions that movements of water, solutes and dissolved gases have been diffusion controlled for very long times, in many cases for the entire geological history of a formation; the quality and distribution of concentration data are adequate for modelling of diffusion in some cases. Geochemical modelling increasingly sophisticated and realistic approaches to equilibrium thermodynamics and reactive mass transport calculations, tested with field and laboratory data as far as possible, but with some limitations remaining. There are, however, some remaining questions and uncertainties:
ACKNOWLEDGEMENTS This paper has been prepared under the auspices of the Mont Terri Project. The consortium of partners comprises: SNHGS (Swiss National Hydrological and Geological Survey: patron), Nagra (Switzerland), Andra (France), IPSN (France), Enresa (Spain), JNC and Obayashi (Japan), BGR (Germany), SCK.CEN (Belgium). We particularly acknowledge Lise Griffault (chairman) and other members of the Geochemistry Group of the Mont Terri Project.
10
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Alexander, W.R. & I.G. McKinley 1999. The chemical basis of near-field containment in the Swiss high-level radioactive waste disposal concept. In R. Metcalfe & C.A. Rochelle (eds.), Chemical Containment of Waste in the Geosphere: 4769. Special Publications, 157. London: Geological Society. Appelo, C.A.J. & D. Postma 1993. Geochemise, Groundwater and Pollution. Rotterdam: Balkema. Arcos, D., Bruno, J. , Peiia, J., Turrero, M”.J. & A.M. Fernhndez 2001. 1D reactive transport model for the Opalinus Clay at Mont Terri underground laboratory. In R. Cidu (ed.) Water-Rock Interaction - Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this issue). Baeyens, B., Maes, A. & P.N. Henrion, P.N. 1985. In situ physico-chemical characterization of Boom Clay. Radioactive Waste Management and the Nuclear Fuel Cycle. 6 : 391-408. Bath, A.H. 1994. Evolution of porewaters in mudrocks. In H. von Maravic & J. Smellie (eds), Proc. International Workshop, 5th CEC Natural Analogue Working Group Meeting, Toledo, October 1992. EUR Report 15176EN. Luxembourg: European Commission. Bath, A.H. 1995. Hydrochemical Characterisation of Argillaceous Rocks. In Proc. OECD-NEA Workshop on Hydraulic and Hydrochemical Determination of Characteristics of Argillaceous Rocks, Keyworth, June 1994: 77-92. Paris: Nuclear Energy Agency of the OECD. Bath, A. 200 1. Pore Waters in Mesozoic Mudrocks in Southern England. In R. Cidu (ed.) Water-Rock Interaction - Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this issue). Bath, A.H., Entwisle, D., Ross, C.A.M., Cave, M.R., Falck, M. Fry, W.E., Reeder, S., Green, K.A., McEwen, T.J. & W.G. Darling 1988. Geochemistry of pore-waters in mudrock sequences: Evidence for groundwater and solute movements. In Proc. Intl. Assoc. Hydrogeol. @mp. on Hydrogeology and Safety of Radioactive and Industrial Hazardous Waste Disposal, Orle‘ans, June 1988. 1:87-97. Documents du BRGM no 160. Beaucaire, C., Pitsch, H., Toulhoat, P., Motellier, S. & D. Louvat 2000. Regional fluid characterisation and modelling of water-rock equilibria in the Boom clay Formation and in the Rupelian aquifer at Mol, Belgium. Applied Geochemistry 151667-686. Beaufays, R., Bloomaert, W., Bronders, P., De Canniere, P., Del Marmot, P., Henrion, P., Monsecour, M., Patyn, J. & M. Put 1994. Characterisation of the Boom clay and its multilayered hydrogeological environment. EUR Report 14961 EN, Luxembourg: European Commission. Blanc Ph., Gaucher, E., Sanjuan, B., Crouzet, C., Seron, A. & L. Griffault 200 1. Testing a clay/porewater interaction model through a laboratory experiment. In R. Cidu (ed.) Water-Rock Interaction - Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this issue). Bradbury, M.H. & B. Baeyens 1998. A physicochemical characterisation and geochemical modelling approach for
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Interactions. 17:412-422. Mineralogical Society Series: 4. London: Chapman & Hall. Grainger, P. 1984. The classification of mudrocks for engineering purposes. @at. J eng. Geol. London. 17:381387. Griffault, L., T. Merceron, J.R. Mossman, B. Neerdael, P. De Canni&e, C. Beaucaire, S. Daumas, A. Bianchi & R. Christen 1996. Acquisition et regulation de la chimie des eaux en milieu argileux pour la projet de stockage de ddchets radioactifs en formation geologique. Projet Archimbde argile. Rapport final. EUR Report 17454. Luxembourg: European Commission. Hendry, M.J. & L.I. Wassenaar 2000. Controls on the distribution of major ions in pore waters of a thick surficial aquitard. Water Resour. Res. 36503-513. Hendry, M.J., L.I. Wassenaar & T. Kotzer 2000 Chloride and chlorine isotopes (36Cl and 637Cl) as tracers of solute migration in a thick clay-rich aquitard system. Water Resour. Res. 36:285-296. Horseman, S.T., J.J.W. Higgo, J. Alexander & J.F. Harrington 1996 Water, Gas and Solute Movement Through Argillaceous Media. Report CC-96/1 for NEA Working Group ‘Clay Club’. Paris: Nuclear Energy Agency of the OECD. Michelot, J.L., P. Ricard, A. Barbreau & J.-Y. Boisson 1994. Environmental isotope study of pore water and of fracture calcite in the Tournemire Toarcian Claystones: sampling, analyses, and preliminary interpretation. In Proc. OECDNEA Workshop on Determination of Hydraulic and Hydrochemical Characteristics of Argillaceous Rocks, Keyworth, June 1994: 77-92. Paris: Nuclear Energy Agency of the OECD. Nagra 1997. Geosynthese Wellenberg 1996: Ergebnisse der Untersuchungsphasen I und II. Nagra Technischer Bericht NTB 96-01. Wettingen, Switzerland: National Cooperative for the Disposal of Radioactive Waste Neuzil, C.E. 1994. How permeable are clays and shales? Water Resour. Res. 30:145-150. Pearson, F.J. 1999. What is the porosity of a mudrock? In A.C. Aplin, A.J. Fleet & J.H.S. Macquaker (eds). Muds and Muaktones: Physical and Fluid Flow Properties: 9-2 1 . Special Publications, 158. London: Geological Society. Pearson, F.J. Jr. & A. Scholtis 1995. Controls on the chemistry of pore water in a mar1 of very low permeability. In Y.K. Kharaka & O.V. Chudaev, O.V. (eds). Water-Rock Interaction: Proc. - WRI-8, Vladivostok: 35-38. Rotterdam: Balkema. Pearson, F. J. & H.N. Waber 2001. Origin and Evolution of Pore-Water SoIutes in the Very-Low Permeability Opalinus Clay, Switzerland. In R. Cidu (ed.) Water-Rock Interaction Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this volume). Riibel, A. & C. Sonntag 2001. Profiles of noble gases and stable isotopes across the Opalinus Clay Formation at Mont Terri, Swiss. In Water-Rock Interaction - Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this volume).
Thury, M. & P. Bossart 1999. Mont Tem Rock Laboratory. Results of the Hydrogeological, Geochemical and Geotechnical Experiments Performed in 1996 and 1997: Geological Report 23. Bern: Swiss National Hydrological and Geological Survey. Waber, H.N., S.K. Frape & A. Gautschi 2001. C1-Isotopes as Indicator for a Complex Paleohydrogeology in Jurassic Argillaceous Rocks, Switzerland. In R. Cidu (ed.) WaterRock Interaction - Proc. WRI-10, Villasimius, 10-15 June 2001. Rotterdam: Balkema (this volume). Walraevens, K., J. Cardenal, D. De Smet, D. & W. De Breuck 1998. Palaeo and present-day fluid flow through Eocene clay layers in Flanders: Hydrogeological and hydrogeochemical evidence for the present-day existence of preferential pathways in the Bartonian Clay. In Proc. NEUEC Workshop on Fluid Flow Through Faults and Fractures in Argillaceous Formations, Berne, June 1996: 369-380. Paris: Nuclear Energy Agency of the OECD.
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Water-Rock lnteraction 2001, Cidu (ed.), 02001 Swefs & Zeitlinger, Lisse, lSBN 90 2651 824 2
Interpreting kinetics of groundwater-mineral interaction using major element, trace element, and isotopic tracers S.L.Brantley, M.Bau, S.Yau & B .Alexander Dept. of Geosciences, Pennsylvania State University, Univ. Pk PA 16802 J.Chesley Dept. of Geosciences, University ofArizona, Tucson AZ 85721
ABSTRACT: To determine the residence time of groundwater, hydrologists use environmental tracers or numerical modelling. If mineral reactions within an aquifer could be used to determine the age of packets of water, such calculations would be extremely useful. Such calculations assume plug flow in the aquifer, and assume that laboratory mineral-water reaction kinetics are applicable in the field system. Application of such a model to an extraordinarily "simple" aquifer -- the Cape Cod aquifer in Massachusetts, USA -- reveals several puzzles related to interpretation of mineral reactivity. Despite a mineralogy consisting of more than 90% quartz and alkali feldspar, dissolution of quartz in the aquifer is not expected due to concentrations of Si higher than saturation with respect to quartz. However, flow rates inferred from Si gradients in the aquifer and alkali feldspar dissolution rates measured in the laboratory are too fast by a factor of about 50 as compared to observations and hydrologic models developed for the site. Furthermore, analysis of solute concentrations and Sr isotopes are consistent with only minor dissolution of alkali feldspar. In contrast, observed concentrations are consistent with dissolution of plagioclase (present at very low concentrations) and ion exchange between the aquifer fluid and accessory phases such as glauconite. This conclusion documents that chemistry of accessory minerals, when more reactive than the host minerals, can dominate groundwater solute compositions. While this observation is not necessarily unexpected, it emphasizes that further investigations of in situ reactions as tools for dating groundwater must rely upon thorough characterization of aquifer mineralogy combined with the use of major and trace element as well as isotopic analysis. Without such characterization, misinterpretation of reaction stoichiometries and kinetics in natural systems will be unavoidable. CFC concentration in the atmosphere during the last several decades. Another type of environmental tracer, event tracers, can be used similarly. These tracers are distinctive in that they were introduced into the atmosphere during specific events (e.g. nuclear weapons testing during 1952-196314). Environmental tracers and their approximate range of dating application are summarized by Plummer et al. (1993a,b). Unfortunately, not all of these tracers are applicable in all aquifers, and most of these analyses are expensive or limited to a few laboratories. Here, we use mineral dissolution kinetics to assess the residence time of groundwater in the aquifer that comprises the U.S. Geological Survey (U.S.G.S.) Toxic Substances Research Site at Cape Cod, Massachusetts USA (LeBlanc 1984; Garabedian et al. 1991; LeBlanc et al. 1991). This site represents a "simple" aquifer -- a best case aquifer wherein analysis of mineral dissolution was
1 INTRODUCTION Analysis of the direction and flow of groundwater is necessary to understand rates of water recharge, to assess aquifer vulnerability to contamination, to identify recharge and discharge areas, and to test groundwater flow models. In order to determine the residence time of groundwater (to "date" the groundwater), researchers typically use an environmental tracer. The age of a groundwater sample is the time since the sample of water became isolated from the atmosphere. For short residence time situations, dyes can be injected into the flow system, and the breakthrough curve can be assessed. Also of particular interest are isotopes or compounds that can be analyzed readily in groundwaters to infer ages. For example, chlorofluorocarbons (CFCs) introduced into the atmosphere largely from refrigerants have been related to the age of the groundwater through knowledge of the change in
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Figure 1. Cross section of Cape Cod aquifer showing shaded contaminant plume (modified after LeBlanc 1984).
anticipated to be unequivocal and where ample hydrologic studies have been completed. For this site, it was felt that several natural tracers would unambiguously reveal groundwater ages.
chemistry of the plume and the hydrology has been completed for this site by the U.S.G.S. (e.g. Wood et al. 1990; Garabedian et al. 1991; LeBlanc et al. 1991; Coston et al. 1995; Masterson et al. 1996). Sixty-five filtered (0.45 pm), uncontaminated groundwater samples (samples outside the plume) were collected from 10 MLS on June 5 - 6, 1997. Cation and strontium samples were acidified in the lab with concentrated ultrapure nitric acid. Aquifer sediment samples were collected from the uncontaminated part of the aquifer by the U.S.G.S in 1988. Sediment samples were sieved with a 1 mm polyethylene screen, which passed 79% (wt.) of the sediment (Coston et al. 1995). The age of groundwater increases with increasing depth in the aquifer (Shapiro et al. 1999). Si concentrations (Fig. 2) and pH also increase with depth, as would be expected if feldspar were dissolving and these parameters increased with groundwater age. As minerals dissolve, the net dissolution rate can decrease as the driving force for dissolution (expressed as the 6G of reaction) of the
1.1 Hydrologic setting The aquifer consists of unconsolidated sand and gravel underlain by fine sand and silt (LeBlanc 1984) in western Cape Cod (Fig. 1). A sewage treatment plant at a military base on site released contaminants into the aquifer beginning in 1936, and the contaminant plume has been documented by the U.S.G.S., which has sampled wells throughout the area. The effective porosity of the aquifer sand is 0.39 and the groundwater flow rate estimated from tracer tests is 0.42 m/d (Garabedian et al. 1991; Le Blanc et al. 1991). The water table slopes 0.15% in the S-SW direction, and flow lines are therefore near horizontal. The Cape Cod sand and gravel consists of 90 95% quartz and 5 - 10% alkali feldspar, glauconite, iron oxides, mica and other ferromagnesian minerals (Barber et al., 1992; Coston et al., 1995; Yau, 1998). Median grain size is 0.5 mm, and silt/clay and organic carbon comprise 4 % and 0.1% of sediments respectively (Barber et al. 1992). The grains are coated with Fe-Al-Mn-Si coatings (Wood et al. 1990; LeBlanc et al. 1991; Coston et al. 1995, Yau 1998), and the grains have significant internal porosity which contains clay minerals. 2 METHODS AND RESULTS 2.1 Si as a tracer The U.S.G.S. Toxic Substances program has installed and maintained multilevel samplers (MLS) in many sites within the aquifer. The MLS contain 15 polyethylene tubes which run from the surface to various depths. Other wells onsite are cased with PVC and are screened. Extensive analysis of the
Figure 2. Concentrations of Si (ppm) plotted as a function of depth in the Cape Cod aquifer.
14
to equilibrium, and/or iv) other inadequacies in the model.
water approaches 0, the value for AG at equilibrium. For reactions proceeding spontaneously, the AG must be negative. To determine whether the increase in Si in the Cape Cod aquifer is due to feldspar or quartz dissolution, we calculated the AG for groundwater samples using SOLMINEQ (Kharaka et al. 1988). Everywhere in the aquifer the water is oversaturated with respect to quartz, and undersaturated with respect to feldspar, except for a very few locations which are supersaturated with respect to potassium feldspar. Even using the higher quartz solubility quoted by Rimstidt (1997), the groundwater is supersaturated with respect to quartz except at shallow depths. We can use feldspar dissolution rates determined from laboratory experiments under far from equilibrium conditions to calculate groundwater residence time using the following equation (Gislason & Eugster 1987):
2.2 Sr isotopes as tracers To check whether surface coatings affect dissolution, we have measured dissolution kinetics of aquifer grains with and without coatings (Yau 1998), and these results will be presented elsewhere. However, to check whether our kinetic model is even applicable to this system, we also measured Sr concentrations and isotopic compositions vs. depth. Sr isotopic ratios in Cape Cod groundwater range from 0.71123 to 0.71265 and show a systematic decrease with increasing sampling depth (as exemplified for one well in Figure 3). This is considerably less radiogenic than the s7Sr/s6Srratio of the bulk aquifer sediment (0.72017). Sr isotopic data for mineral separates reveal that the isotopic composition of Sr in this bulk sediment is intermediate between that of a very radiogenic Kfeldspar component (s7Sr/86Sras high as 0.74533) and that of a plagioclase + Fe-silicate separate (87Sr/86Sras low as 0.71 149). The latter, however, shows a s7Sr/86Srratio that is close to that of the groundwater. Together with the depth-dependent increase of Si concentrations in the water, this suggests that dissolution of plagioclase controls the Sr budget of Cape Cod groundwater, whereas Kfeldspar (or any other K-Rb-rich detrital mineral rich in radiogenic 87Sr)does not contribute significant Sr. In spite of the rather small but systematic variation of Sr isotopic ratios with depth, the Sr concentration is variable. Sr concentration generally increases with increasing depth; however, concentration spikes occur at some depths, such as the spike at 10 m depth shown in Figure 3. This spike is accompanied by a spike in the K concentration but is not reflected in the Si or 87Sr/86Srvs. depth plots (Fig. 3). Hence, this additional Sr input does not result from dissolution of silicate minerals such as plagioclase. Sr release from detrital K-Rb-rich minerals (that would be rich in radiogenic s7Sr)can also be excluded. Considering that carbonates are observed to be absent from the aquifer sediment, this suggests that ion exchange with the authigenic Fe-silicate (glauconite) component causes the Sr-K concentration spike. Although rich in K and Rb, this glauconite is an authigenic phase which is too young to have accumulated significant radiogenic 87Sr;the Sr released from this glauconite is inferred to be close to the 87Sr/86Srratio of seawater (Clauer et al. 1992). Such a low ratio resembles the low isotopic ratio of Sr released from dissolving plagioclase.
where [Si] is the concentration of dissolved Si, t is the groundwater residence time along a given flow path, vj is the stoichiometric coefficient (number of moles Si / mol feldspar), kfeld is the dissolution rate constant of felds ar at the pH of the aquifer (mol feldspar cm-2 s- ), and S is the surface area of feldspar (cm2 feldspar/cm3 water). In this model, we are implicitly assuming that water flows through the porous media as plug flow. We also assume little to no change in pH, or that the rate constant is not dependent upon pH (Schweda 1990), which should be approximately true at the near neutral conditions of the aquifer. To calculate residence time (or flow rate) from this equation, we estimate the change in [Si] over some flow distance, the laboratory dissolution rate for feldspar, and the surface area of feldspar per volume of aquifer pore fluid. Using published values for the dissolution rate constant for far from equilibrium conditions (Schweda 1990), our own measured value of the surface area normalized by pore volume (1540 cm2 feldspar cm-3ground water), and a preliminary estimate of A[Si] over a known flow distance observed from our own sampling, we calculate a residence time for this flow path of less than a year, or a flow rate of about 20 m/d. In contrast, published values for flow rates based upon MODFLOW modelling suggest rates on the order of 0.4 m/d (a discrepancy of about a factor of 50). This 50-fold discrepancy could result from i) surface coatings on aquifer minerals (Coston et al. 1995), ii) heterogeneities in flow path that are not accounted for in the modelling, iii) dissolution occurring near
P
15
Apparently, the concentration of Sr (and other solutes) in the groundwater is not exclusively controlled by plagioclase dissolution, but also by accessory minerals in the sediment. This additional Sr input is hidden in the isotopic data because of the close isotopic similarity of the Sr released from these different pools.
provenance of dissolved components in soil and groundwaters (e.g. Izbicki et al. 1994; Johnson & DePaolo 1994; Katz & Bullen 1996; Bullen et al. 1996, 1997; Glow et al. 1997). Many ofthese studies are complicated by the diverse Source of Sr in the study aquifers, which may include significant carbonates and silicates; however, many of the studies are not accompanied by extensive analysis of aquifer mineralogy. Perhaps the most quantitative treatment of the use of Sr isotopes to determine rates of mineral dissolution is summarized in the papers by Johnson and coauthors (Johnson & DePaolo 1994; Johnson & DePaolo 1996; Johnson & DePaolo 1997a,b). These authors have argued that Sr isotopes can help to constrain sources of solutes and movement of ground- or formation waters. Indeed, based on the results reported here, interpretation of Si release rates can be clarified when accompanied by thorough analysis and interpretation of trace element and isotope chemistry as a function of flow in order to constrain mineral reaction in even the "simplest" of aquifers. In ongoing work we are coupling these analyses with analyses of rare earth elements in some of the wells to contrain mineral reactivity even more thoroughly. Without such extensive analysis, accurate interpretation of single tracer systems such as Si or Sr will be difficult at best, given that minor phases with high rates of dissolution can control dissolved concentrations (Brantley et al. 1998). For the Cape Cod aquifer, combined Sr isotopes and rare earth element chemistry will be used to elucidate the stoichiometries and kinetics of reactions and determine what factors cause the apparent slowness of dissolution rates in the field compared to the laboratory. Further interpretation of geochemical kinetics in other field systems should similarly combine several major and trace element along with isotope tracers to better constrain contributions from reacting minerals. ACKNOWLEDGEMENTS Funding from the Dept of Energy Office of Basic Energy Sciences grant DE-FG02-95ER14547 to SLB is acknowledged.
Figure 3. A) Graph of Sr concentration in groundwater vs. depth for well 3. B) Graph of p7Sr/g6Srratio in groundwater vs. depth for well 3.
3 DISCUSSION AND CONCLUSIONS
REFERENCES
A large number of investigations into the use of Sr isotopes in ground- and formation waters have been reviewed by Johnson & DePaolo (1997a). For example, Sr concentrations and Sr isotopic signatures have been used to determine the
Barber, L.B., Thurman, E.M. & D.D. Runnells 1992. Geochemical heterogeneity in a sand and gravel aquifer: Effect of sediment mineralogy and particle size on the sorption of chlorobenzenes, J Contam. Hydrol. 9: 35-54. 16
The effects of longitudinal dispersion and the use of Sr isotope ratios to correct for water-rock interaction, Water Resources Research, 32, 7: 2203-2212 Johnson, T.M., & D.J. DePaolo 1997. Rapid exchange effects on isotope ratios in groundwater systems: 1. Development of a transportdissolution-exchange model, Water Resources Research, 33, 1: 187-195 Johnson, T.M., & D.J. DePaolo 1997. Rapid exchange effects on isotope ratios in groundwater systems: 2, Flow investigation using Sr isotope ratios, Water Resources Research, 33, 1: 197-209 Katz, B.G. & T.D. Bullen 1996. The combined use of 87Sr/86Sr and carbon and water isotopes to study the hyrochemical interaction between groundwater and lakewater in mantled karst, Pergamon, 5075-5087p Kharaka, Y.K., Gunter, W.D., Aggarwall, P.K., Perkins, E.H. & J.D. DeBraal 1988. SOLMJNEQ88:A computer program code for geochemical modeling of water-rock interactions, US. Geol. Surv. Water Investigations Report 88. LeBlanc, D.R., Garabedian, S.P., Hess, K.M., Gelhar, L.W., Quadri, R.D., Stollenwerk, K.G. & W.W. Wood 1991. Large-scale natural gradient tracer test in sand and gravel, Cape Cod, Massachusetts, 1. Experimental design and observed tracer movement, Water Resour. Res. 27: 895-910. LeBlanc, D.R. 1984. Sewage plume in a sand and gravel aquifer, Cape Cod, Massachusetts, US. Geol. Surv. Water Supply Pap. 221 8,28 p. Masterson, J.P., Walter, D.A. & J. Savoie 1996. Use of particle tracking to improve numerical model calibration and to analyze ground water flow and contaiminant migration, Massachusetts Military Reservation, Western Cape Cod, Massachusetts, U. S. Geol. Surv. Open-File Report 96-214, 50p. Plummer, L.N., Dunkle, S.A. & E. Busenberg 1993. Data on chlorofluorocarbons (CC13F and CC12F2) as dating tools and hydrologic tracers in shallow ground water of the Delmarva Peninsula, OpenFile Report- U. S. Geological Survey, OF 930484,56 p. Plummer, L.N., Michel, R.L., Thurman, E.M. & P.D. Glynn 1993. Environmental tracers for age dating young groundwater, Alley, William M. (editor), Regional ground-water quality, p. 255294. ISBN: 0-442-00937-2. Rimstidt, J.D. 1997. Quartz solubility at low temperatures, Geochimica et Cosmochimica Acta 61: 2553-2558.
Brantley, S.L., Chesley, J.T. & L.L. Stillings 1998. Isotopic ratios and release rates of Sr measured Geochimica from weathering feldspars. Cosmochimica Acta 62: 1492-1500. Bullen T.D., Krabbenhoft D.P. & C. Kendall 1996. Kinetic and mineralogic controls on the evolution of groundwater chemistry and 87Sr/86Sr in a sandy silicate aquifer, northern Wisconsin. Geochim.Cosmochirn.Acta 60: 1807 - 1821. Bullen, T.D., White, A.F., Blum, A.E., Harden, J.W. & M.S. Schulz 1997. “Chemical weathering of a soil chronosequence on granitoid alluvium: 11. Mineralogic and isotopic constraints on the behavior of strontium.’’ Geochim. Cosmochim. Acta 61: 291-306. Clauer, N., Keppens, E. & P. Sr. Stille 1992. Isotope constraints on th eprocess of glauconitization, Geology 20: 133-136. Clow, D.W., Mast, MA., Bullen, T.D. & 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 Research, 33: 13351351 Coston, J.A., Fuller, C.C. & J.A. Davis 1995. Pb2+ and Zn2-t adsorption by a natural aluminum- and iron-bearing surface coating on an aquifer sand, Geochimica et Cosmochimica Acta 59: 35353547. Garabedian, S.P., LeBlanc, D.R., Gelhar, L.W. & M.A. Celia 1991. Large-scale natural gradient tracer test in sand and gravel, Cape Cod, Massachusetts, 2. Analysis of spatial moments for a nonreactive tracer, Water Resour. Res. 27: 91 1-924. Gislason, S.R. & H.P. Eugster 1987. Meteoric waterbasalt interactions, 11. A field study in N.E. Iceland, Geochimica et Cosmochimica Acta 58: 284 1-2855. Izbicki, J.A., Bullen, T.D. & R.L. Michel 1994. Use of 87Sr/86Srin ground water to identify the source of deposits underlying the Oxnard Plain and Pleasant Valley, California, Eos, Transactions, American Geophysical Union, 75, Suppl., 280. Johnson, T.M. & D. J. Depaolo 1994. Sr isotope ratios as indicators of water-solid interaction and fluid flow at the Lawrence Berkeley Laboratory site. Eos Transactions. AGU 75, 280. Johnson, T.M., & D.J. DePaolo 1994. Interpretation of isotopic data in groundwater-rock systems: Model development and application to Sr isotopic data from Yucca Mountain, Water Resources Research, 30, 5: 1571-1587 Johnson, T.M., & D.J. DePaolo 1996. Reactiontransport models for radiocarbon in groundwater: 17
Schweda, P. 1990. Kinetics and mechanisms of alkali feldspar dissolution at low temperatures, Ph. D. Dissertation, Stockholm University. Shapiro, S.D., LeBlanc, D., Schlosser, P. & A. Ludin 1999. Characterizing a sewage plume using the 3H-3 He dating technique, Ground Water, 37, 861-878. White, A.F. & S.L. Brantley (eds) 1995. Chemical Weathering Rates of Silicate Minerals, MSA Short Course v. 31, Mineralogical Society of America. Wood, W.W., Kraemer, T.F. & P.P. Jr. H e m 1990. Intergranular diffusion: An important mechanism influencing solute transport in clastic aquifers? Science 247: 1569-1572. Yau, S. 1998. Feldspar Dissolution Kinetics in the Cape Cod Aquifer, Massachusetts, Masters Thesis, Dept. of Geosciences. Pennsylvania State University.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Reducing ambiguity in isotopic studies using a multi-tracer approach Thomas D.Bullen & Arthur F.White Water Resources Division, US. Geological Survey, Menlo Park, California, USA
Cyril W.Childs School of Chemistry and Physical Sciences, Victoria University, Wellington,New Zealand
Juske Horita Oak Ridge National Laboratory, Oak Ridge, Tennessee, USA
ABSTRACT: Naturally-occurring radiogenic and stable isotope tracers are commonly used in studies of water-rock interaction to provide both process and water/solute source information. Several isotope tracers have been proposed as diagnostic process indicators, and have been used for that purpose without a full understanding of important factors controlling their behavior such as isotope distributions and fiactionation mechanisms. Experience shows that these initially promising isotope tracers considered alone are likely to provide ambiguous interpretations, and that additional geochemical andor isotopic constraints are required. Isotope data must be considered in a multi-tracer context if potential ambiguity is to be minimized.
1 INTRODUCTION
This paper explores three examples in which an isotope tracer, widely accepted to be useful alone as a diagnostic indicator of some natural process, has proven to provide ambiguous information about that process. The examples include: 1) the use of strontium (Sr) isotopes as an indicator of relative weathering rates of granitoid minerals; 2) the use of iron (Fe) isotopes as a "biosignature" in the search for evidence of ancient or extra-terrestrial life; and 3) the use of carbon (C) isotopes to distinguish biogenic and abiogenic methane. In each case, the isotope tracer has been viewed as a "silver bullet" that can be used to diagnose the natural process with great confidence. Moreover, in the case of Sr and Fe a "bandwagon"-like acceptance of the robustness of the tracer occurred largely due to the convincing arguments of the early proponents. In each case, however, later work demonstrated that the physical and chemical complexities of natural systems severely limit the isotope's use as an independent diagnostic process indicator. Fortunately, the application of a multi-tracer approach should allow each of these isotope tracers to ultimately be extremely useful for their original intended purpose.
A variety of radiogenic and stable isotope tracers are commonly used in studies of water-rock interaction, primarily because they provide unique geochemical information and their general behavior in hydrogeologic and biogeochemical systems is believed to be well-understood. Isotopes have long held an almost mystical status in the geochemical hierarchy, certainly earned but in part due to the perceived difficulty associated with analyzing isotope compositions as well as to the relatively small number of laboratories that perform those analyses. However, as more studies that use isotopes are published and as we learn more about the factors that control isotope distribution, it is becoming increasingly clear that interpreting variations of a given isotope tracer on their own is more likely to provide an ambiguous vision of the processes being studied. For example, isotopic variability can develop due to processes on the microscopic (e.g., pore space between mineral grains) to megascopic (e.g., aquifer) scales with similar ranges of values. Thus the isotopic compositions determined for a single sample or for a reconnaissance suite of samples cannot be used in the absence of other information to determine the nature or scale of the process that led to those particular compositions. The obvious way to decrease ambiguity is to consider the isotopic compositions in light of additional isotope or geochemical tracers that provide information on the scale issue, and thus employ a "multi-tracer approach".
2 FUNDAMENTALS OF THE "MULTI-TRACER, APPROACH The goal of the multi-tracer approach is to identify additional isotope andor chemical tracers that in some way constrain the geochemical behavior of the isotope system. Especially useful multi-tracer
19
approaches are those that consider two or more components of a molecule or mineral. For example, variations of hydrogen and oxygen isotope compositions of H20 considered together provide a tracer of water evaporation, assuming that waterrock isotope exchange can be ruled out (Hoefs 1987). The evaporation "signal" can be used to distinguish flowpaths taken by evaporated surface waters that recharge an aquifer (e.g., lake waters) from those taken by meteoric waters that recharge the aquifer through the unsaturated zone (e.g., Krabbenhoft et al. 1994). Similarly, variations of nitrogen and oxygen isotope compositions of N03considered together can provide a discriminant of a variety of natural and anthropogenic nitrate materials and the biogeochemical processes that affect them (e.g., Kendall 1998). An alternative approach is to use unrelated isotope or chemical tracers that provide independent information about a particular process. For example, Bullen & Kendall (1998) used C and Sr isotopes together to determine sources of solutes in stream waters and associated weathering processes at two adjacent watersheds developed on calcite-bearing greenstone. The C isotope compositions of the stream waters could be explained entirely by calcite weathering. On the other hand, the Sr isotope compositions were different from that of the bedrock calcite, and similar to those of regional groundwater samples that had much lighter C isotope compositions than those resulting from calcite weathering. However, on a plot of C isotopes versus Sr isotopes the streamwaters formed a mixing array between two "end-member" groundwater types having chemistry controlled by either carbonic acid weathering of calcite or strong-acid weathering of calcite-bearing silicate rocks. Neither of the postulated end-member groundwaters had both C and Sr isotope compositions similar to those of the regional groundwaters, suggesting that chemical variations probably developed along flowpaths channeled along the bedrock-soil interface in the near-stream environment. The examples presented below for the Sr, Fe and C isotope systems are particularly interesting in that they provide insight into the realities and pitfalls associated with practical development of isotope tracer capabilities. In addition, Sr, Fe and C span a broad range of geochemical and isotopic behavior, and their isotopes can be used for a variety of waterrock interaction applications. By studying these isotope systems, we have learned a considerable amount of new information about the way that isotopes behave in general, and thus may be able to better predict how other as yet unexplored stable and radiogenic isotope tracers may behave in nature.
3 Sr ISOTOPES AS AN INDICATOR OF MINERAL WEATHERING RATES Sr isotopes are one of the more commonly used isotope tracers in studies of water-rock interaction. Variations in the Sr isotope composition of rocks develop from the variable, long-term radioactive decay of 87Rbto s7Sr (half life = 4.9 x 10'' y) in their constituent minerals. In granitoids and their weathering derivatives, minerals having elevated Rb/Sr concentration ratios ( e p micas and Kfeldspar) develop greater s7Sr/8Sr ratios over time than minerals having low Rb/Sr ratios (e.g., plagioclase feldspar, hornblende). Similarly, as a general rule older granitoids and derivative rocks tend to have greater 87Sr/86Srratios than younger rocks, although old, Rb-poor rocks will have experienced little radiogenic ingrowth of 87Sr. Therefore Sr isotopes are most useful for distinguishing Sr contributions from Rb-rich vs. Rbpoor minerals, and from old vs. young rocks. In addition, silicate minerals exposed to weatherin on g8 the continents have on average greater s7Sr/ ratios than exposed marine carbonates (Faure 1986). Sr isotopes are an especially powerful indicator of the sources of dissolved cations in waters because the waters inherit the 87Sr/s6Srratio of the rocks with which they interact. Once inherited by the water, the s7Sr/86Srratio is not changed by loss of Sr due to mineral precipitation or cation exchange and thus reflects the source rock "signature". Of course, waters can interact with a variety of minerals and rocks along a flowpath, and can inherit Sr of a different isotopic composition from exchange sites. Thus, the s7Sr/s6Srratio measured in water at any point along a flowpath reflects the unique mixture of Sr derived from upgradient sources. Sr isotopes should therefore be useful as an indicator of both the relative weathering rates of silicate minerals in a specific granitoid terrain, and the relative weathering rates of silicate and carbonate rocks over a lithologically diverse region. Understanding relative mineral weathering rates and contributions to solute fluxes is important for both the development of chemical budgets and the identification of important weathering processes. Sr isotopes have been used for this purpose in a variety of recent studies, including the effort to understand the effects of tectonics and glaciation on the riverine budget of Sr. For example, Edmond (1992) postulated that from a global perspective only Himalayan-style collisional tectonic events are capable of causing major excursions in 87Sr/86Srof seawater as recorded in marine carbonates, such as occurred during the mid-Cenozoic. He attributed the large increase in S7Sr/s6Srof seawater that occurred from 40 to 20 Ma to increased silicate weathering relative to the buffering effect of carbonate
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Even after protracted weathering when plagioclase and biotite are likely to be the only important weathering minerals, other factors such as temperature have significant effects on both mineral weathering rates and the isotopic composition of Sr released to weathering solutions. For example, White et al. (1999a) demonstrated that plagioclase dissolution rates decreased by a factor of 12 while biotite dissolution rates decreased by a factor of only 2 as the temperature of the experimental granitoid weathering columns mentioned above was decreased from 34OC to 5'C. Over the same temperature change, 87Sr/s6Sr ratios of the column effluents increased, but became progressively less than those predicted from the plagioc1ase:biotite dissolution proportions required for WNa mass balance. These observations suggest that 87Sr/s6Srratios in waters draining cold regions may on average be greater than those in waters draining warm regions, relative to the local bedrock composition, due to preferential suppression of the plagioclase dissolution rate relative to that of biotite at cold temperatures. Clearly, the use of s7Sr/86Srratios to assess relative weathering rates of granitoid minerals in cold regions will result in an under-estimate of the actual bi0tite:plagioclase dissolution mass ratio. Although incongruent Sr release, the importance of K-feldspar dissolution and temperature effects all need to be considered when using Sr isotopes to assess relative weathering rates of granitoid minerals, the influence of trace calcite in the granitoids truly creates considerable ambiguity in the method and can cloud interpretations. Calcite is commonly associated with metamorphosed gneisses and schists and hydrothermally-altered granitoid rocks, and is generally recognized as an important contributor to Ca fluxes in associated watersheds (Drever & Hurcomb 1986, Blum et al. 1998). A useful example of the potential impact of calcite dissolution on both the Ca and Sr budgets of an alpine watershed is provided by Clow et al. (1997). Using the Sr isotope approach but including both bedrock- and aeolian-deposited calcite in the calculations, they determined that on an annual basis Ca in streamwater draining the Loch Vale watershed (Colorado, USA) is derived in equal amounts from dissolution of bedrock calcite, bedrock plagioclase and aeolian-deposited calcite. Thus fully two-thirds of the calcium and approximately one-half of the Sr generated by weathering is derived from calcite dissolution. Further, the 87Sr/86Srratio of bedrock calcite is greater than that of streamflow at any time of the year. Clearly the calcite component must be evaluated if calculations of relative weathering rates of the silicate granitoid minerals are to make sense. White et al. (1999b) pointed out that the chemistry of effluents from the experimental
weathering, and release of especially radiogenic silicate-derived Sr from the uplifted Himalayan orogen. He contended that Sr contributions from silicate weathering elsewhere, as recorded by current riverine inputs, would be insufficiently radiogenic to significantly impact the seawater Sr pool. From a more local perspective, Blum et al. (1994) and Blum & Ere1 (1997) proposed that for the Sierra Nevada and Wind River Ranges, Sr derived from weathering of recently-glaciated granitoid is systematically more radiogenic than Sr derived from weathering of granitoid that has not been recent!y glaciated. They attributed this difference to enhanced weathering of biotite relative to that of plagioclase at the more recently glaciated areas. They concluded that continental glaciation may have the effect of accelerating biotite weathering, elevating riverine s7Sr/86Srratios in regions draining silicate bedrock and possibly causing the more subtle fluctuations in the seawater Sr isotope record as have been observed over the past 450 Ka. The use of Sr isotopes for determination of relative weathering rates of granitoid minerals has been based on three fundamental assumptions: 1) that Sr is released congruently from each mineral during weathering; 2) that Sr contributions from all weathering silicate minerals are included in the calculations; and 3) that trace non-silicate minerals in the granitoids do not contribute significant amounts of Sr during weathering. An expanding body of research over the past several years has demonstrated that each of these assumptions is violated in at least some cases. For example, there is evidence from two studies that Sr may not be released congruently from either plagioclase (Brantley et al. 1998) or biotite (Taylor et al. 2000) during dissolution experiments under controlled laboratory conditions. In both studies, Sr was released preferentially with respect to stoichiometric cations and radiogenic Sr was released preferentially with respect to unradiogenic Sr, particularly when mineral surfaces were fresh. Similarly, there is evidence from flow-through column dissolution experiments using four different fresh and weathered granitoids (Bullen et al. 1998) that K-feldspar dissolution rates may be of similar magnitude to those of plagioclase and biotite during incipient granitoid weathering, and thus release of relatively radiogenic Sr from K-feldspar must be considered. This is an important finding because Kfeldspar is generally perceived to be unreactive, and thus relative-rate calculations tend to consider only plagioclase and biotite dissolution as significant components. The fact that these effects are most pronounced when mineral surfaces are fresh suggests that they need to be considered in studies of relative weathering rates in glaciated terrains, where fresh mineral surfaces are likely to predominate.
21
process. Moreover, hydrothermal calcites in the Sierra Nevada and Wind River granitoids probably have s7Sr/86Srratios that are intermediate between those of plagioclase and biotite, and thus calcite dissolution can account for the greater 87Sr/86Sr ratios of streamwaters draining recently glaciated areas in those regions. Ambiguity can be greatly reduced or eliminated by considering the Sr isotopes together with other geochemical information. For example, silicate and carbonate 'lend-membertt compositions may be deduced on plots of s7Sr/86Srratios versus NdCa or Si/Ca ratios (cf., Clow et al. 1997). An example is shown in Figure 1, in which these ratios in experimental flow-through column effluents using a granodiorite from Yosemite Valley, California are compared. Chemical variations indicate that at reaction times less than -100 hours, dissolution of the freshly-ground feldspar surfaces dominated the solute flux. During the next several hundred hours, calcite dissolution became the dominant control. Finally dissolution of biotite and plagioclase became dominant as the calcite was slowly consumed (White et al. 1999b). From the lowest to the highest NdCa and Si/Ca ratios the data form a simple mixing array between two 8'Sr/86Sr ratios: -0.71 1, the inferred composition of the calcite, and -0.720, the inferred composition derived from dissolution of biotite and fiagioclase in the absence of calcite. Note that the Sr/'%r ratio of the calcite is considerably greater than that of plagioclase in this rock (0.7065). This method should be applicable to field data as a means of extracting information about Sr contributions from different minerals. Alternatively, simple watershed chemical mass balance calculations in the style of Garrels & Mackenzie (1967) are especially useful for indicating the need to consider unanticipated trace minerals that might
columns containing fresh granitoid mentioned above was in each case initially dominated by calcite dissolution, which then progressively decreased in importance during reactions lasting up to 1.7 yr. The chemistry of effluents from the columns containing weathered granitoid showed little to no evidence of calcite dissolution, suggesting that calcite had been effectively stripped from the granitoids in their field settin!. Mass balance calculations reveal that the s7Sr/8Sr ratio of calcite is in each case intermediate to that of plagioclase and biotite, and thus the usefulness of Sr isotopes alone as an indicator of relative weathering rates of the silicate minerals is essentially lost. White et al. (1999b) noted that stream waters draining recently glaciated, granitoiddominated catchments tend to have high Ca/Na ratios relative to plagioclase stoichiometries, a direct reflection of the amount of calcite exposed in the bedrock. Thus it appears that the use of Sr isotopes on their own to determine relative weathering rates of silicate minerals in recently-glaciated terrains is at best problematic, and probably a futile effort. Of course the ultimate conclusions of Edmond (1992) (i.e., that only Himalayan-style collisional tectonic events are capable of causing major excursions in 87Sr/s6Srof seawater as recorded in marine carbonates) and Blum et al. (1994) and Blum & Ere1 (1997) (i.e., that relatively radiogenic Sr is preferentially released from granitoids as a response to glacial processes) remain valid in spite of the complications outlined above. However, the ability of silicate mineral weathering processes to account for those conclusions is clearly in question. In fact, Blum et al. (1998) stress the importance of calcite weathering in controlling the Ca and Sr flux in rivers draining the High Himalayan Crystalline Series rocks, in contrast to the contention of Edmond (1992) that silicate weathering is the important
Figure 1. Sr isotope and chemical compositions of Yosemite (CA) experimental granitoid column effluents during first 15,000 hours of reaction. Arrows show direction of increasing reaction time. Mineral labels show predominant mineral contributors at various reaction times. In this rock, 87Sr/s6Srratios of component minerals are: plagioclase, 0.7065; hornblende, 0.7069; K-feldspar, 0.7075; biotite, 0.8053; calcite, 0.7 1 1.
-
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impact the Sr isotope systematics. Of course this approach is itself assumption bound and can be problematic. Obvious12 any additional parameters used to constrain 87Sr/ Sr ratios must be applicable to the dynamics of the particular field situation. 4 Fe ISOTOPES AS A “BIOSIGNATURE” Iron (Fe) has four stable isotopes (54Fe, 56Fe, 57Fe and ”Fe), and thus there is the potential for variability of stable isotope compositions to develop in nature through either abiotic or biotic processes. For example, atomic/molecular vibration theory predicts that Fe isotopes may be fractionated during both equilibrium and uni-directional abiotic mineralhid reactions due to the relative stability of the heavier isotopes in more tightly-bound minerals and aqueous species (Urey, 1947, Polyakov, 1997). On the other hand, microbes are capable of using Fe during both dissimilatory and assimilatory redox processes as either an electron donor or acceptor (e.g., Lovley et al. 1991, Mandernack et al. 1999). Microbially-mediated redox processes may result in Fe isotope fractionation due to a variety of kinetic considerations, such as reactant- and transitioncomplex bond strengths and diffusional gradients across cell membranes. Depending on the particular multi-step reaction pathway, the amount of Fe isotope fractionation observed during microbiallymediated processes may be greater than or less than that observed during abiotic processes. Recently it has been proposed that microbiallymediated Fe isotope fractionation, on the order of 1 to 2%0 in terms of the 56Fe/54Feratio, is of greater magnitude than the fractionation resulting from abiotic processes in nature (Beard and Johnson, 1999, Beard et al., 1999). This proposal was based on a comparison of the -1.3%0 fractionation of the 56Fe/54Fe ratio observed during dissimilatory reduction of Fe in the Fe-oxyhydroxide mineral ferrihydrite by Shewenella algae under laboratory conditions with the near uniform isotopic composition of Fe (range of 0.4%0)in terrestrial and lunar igneous rocks. If true, then Fe isotopes might be useful as an effective “biosignature” for a variety of applications such as remote sensing of microbially-mediated processes (e.g., monitoring the progress of subsurface degradation of organic contaminants in aquifers; Bullen & McMahon 1998) and in the search for evidence of ancient or extraterrestrial life (e.g., constraining the origin of Banded Iron Formation, discerning evidence of microbial activity in Martian meteorite ALH 8400 1; Beard et al. 1999). This latter application has
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generated considerable excitement and controversy in the scientific community. However, Anbar et al. (2000) convincingly demonstrated that Fe dissolved in acid media could be abiotically fractionated by more than 10%0 in terms of the 56Fe/54Feratio in a simple anion resin exchange column experiment. They explained the fractionation mechanism in terms of equilibrium partitioning of Fe isotopes among at least three coexisting dissolved Fe-chloro-aquo complexes, with only one of the complexes having affinity for resin exchange sites. Other workers have questioned the interpretation of equilibrium isotope partitioning, postulating instead “kinetic isotope effects” to explain Fe isotope fractionation (Beard et al. 2000). Although the relevance of exchange resin column experiments to natural systems is debatable, the results clearly indicated a need to assess both Fe isotope fractionation mechanisms and the importance of abiotic Fe isotope fractionation in nature particularly in low-temperature systems. There are now several completed or ongoing studies of Fe isotope fractionation in natural systems. These include studies of microbiallymediated Fe redox processes in a fuel-spill impacted aquifer (Bullen & McMahon 1998), precipitation of the mineral ferrihydrite (Bullen et al. 1999), Fe cycling in a seasonally-reduced soil (Wiederhold 2000) and Fe transport in an acid mine drainageimpacted stream (Mandernack et al. 2000). Each of these studies has approached the issue from a reaction flowpath perspective. Comparison of the results of these studies suggests that although the relative importance of biotic and abiotic reactions involving Fe differs in each study setting, the range of Fe isotope fractionation observed in each setting is similar and on the order of a few per mil. Based on these studies, we now have strong evidence that abiotic Fe isotope fractionation occurs in nature, and that the amount of fractionation is essentially the same as that due to microbial processing of Fe. Furthermore, the “signal” of abiotic Fe isotope fractionation, whether caused by equilibrium or unidirectional processes, is likely to be preserved in the geologic record due to processes such as mineral zonation and sedimentary layering and armoring. Obviously these results suggest that the use of Fe isotopes on their own as a “biosignature” is problematic at best and probably not possible. The best evidence for significant abiotic Fe isotope fractionation in nature comes from the ongoing paired field and laboratory investigation of precipitation of the common mineral ferrihydrite from Fe(I1)-rich water (Bullen et al. 1999, Bullen et
al., submitted; unpublished data). In the field settings on the flanks of two andesitic volcanoes on the North Island of New Zealand (Waiuku Spring on Mt. Ruapehu and Kokowai Springs on Mt. Taranaki), ferrihydrite precipitates rapidly from Fe(I1)- and CO2-rich spring waters as the water reaches the surface, oxygenates and flows downstream. The mineralogy, water chemistry and water isotope compositions at each site have previously been well characterized (Childs et al. 1982, 1986). At each site, the calculated oxidation rate of ferrous ions based on thermodynamic theory is sufficient to account for the rate of ferrihydrite precipitation observed, and thus microbial activity is not required to mediate the precipitation process. The two sites provide an interesting comparison, in that Waiuku spring is S04-rich while Kokowai spring is S04-poor. An important conclusion of the field study and the accompanying experimental study is that SO4 apparently provides a control on the amount of Fe isotope fractionation observed. Fe isotope compositions of aqueous Fe in stream waters and spatially-associated ferrihydrite samples from each site are plotted in Figure 2. The Fe isotope compositions are given as 856Fe, the per mil deviation of the measured 56Fe/54Feratio of the sample fi-om that of the U.S.G.S. rock standard BIR1, an Icelandic basalt. These analyses were made using thermal ionization mass s ectrometry (TIMS), employing a mixed 57Fe-5Fe “double-spike” amendment to allow correction of analyticallyinduced stable isotope fractionation (Johnson et al. 1999, Mandernack et al. 1999). At both field sites the greatest fractionation between aqueous Fe and ferrihydrite is observed in the samples collected closest to the spring. The fractionation factor aferrihydrite-aqueous Fe is 1.0019 for Kokowai Springs, and 1.0009 for Waiuku Spring. Thus, the fractionation is greater in the S04-poor system. On the other hand, the observed rate of ferrihydrite precipitation is greater at Kokowai Springs, probably due to the greater turbulence of stream flow immediately downgradient from the spring. Further downstream at each site the Fe isotope compositions of aqueous Fe and ferrihydrite converge. This may be due to a decreased effective fractionation factor downstream, but is more likely due to contamination with colloidal Fe that passed through the 0.10 pm filter at the time of sampling. A series of steady-state laboratory experiments was conducted to simulate the natural ferrihydrite precipitation process. In each experiment, ferrous chloride solution was oxygenated in a pH-stat reactor utilizing a flow-through design that minimized potential isotopic back-reaction between the solid and aqueous phases. Experiments were carried out over a range of pH from 5.4 to 6.2. In
Figure 2. Fe isotope compositions of aqueous Fe and coexisting ferrihydrite samples from New Zealand field sites. The % of Fe remaining in solution is with respect to the Fe concentration of the spring water. one set of experiments at pH=6.0, SO4concentrations were varied to test the influence of SO4 on Fe isotope fractionation. The important results of these experiments are shown in Figure 3. First, as shown in Figure 3a there is an overall trend of increasing Fe isotope fractionation between ferrihydrite and aqueous Fe with increasing relative proportion of Fe(II)HC03+(aq). In addition, at given pH the Fe(I1) oxidation rate in the experiments using a 5% Col95% air gas mixture was approximately twice that in the experiments using a 1% co2-99% air gas mixture. In contrast, as shown in Figure 3b there is an overall trend of decreasing Fe isotope fractionation between ferrihydrite and aqueous Fe with increasing S04-concentration (and relative proportion of Fe(II)S04(aq)). Both of these results suggest that there is a strong control of Fe isotope fractionation by Fe(I1) aqueous speciation. The most intriguing and surprising feature of the field and experimental data is that in all cases the Fe in ferrihydrite is isotopically heavier than the coexisting aqueous Fe. Fe(I1)-oxidation at pH >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.
28
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.
36
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
0.02
0.178
0.01
0-9.5
<0.05
0.05
0.00 18
<0.000005
Iron (Fe) Mercury (Hg)
5.4
<0.05 #
0.82
- Not measured # Not measured at site, but after 3 days storage in a gas bulb
studied (Sigvaldason 1963, Lawler et al. 1996) and also samples collected just after a suspected volcanic event (Kristmannsdottir 1999). A few samples from subaerial springs from the edges of the Grimsvotn caldera (Table 1) aq well as samples from dri,llholes in the lake itself (Agustsdottir et al. 1992, Agustsdottir & Brantley 1994) have been obtained. A sample was collected from open water at the eruption site in Vatnajokull in 1996 (Gislason et al. 1997) two months after the eruption was over. Samples from volcanically affected water from the river Jokulsa a Fjollum draining to the north of Vatnajokull have also been studied (Kristmannsdottir et al. 1999). In Table 1 a typical composition of geothermal water (total flow) from an Icelandic high-temperature area, a typical composition of water from the Skeidara river at the peak of surnrner and the range
of composition for waters sampled from eleven jokulhlaups in the river through the later half of last century are shown. In the table concentrations of most major components in the waters are shown but only those of a few trace elements as little data are available on trace elements until 3-4 years ago (Kristmannsdottir 2000a). Total dissolved solid concentration of the water from jokulhlaups is high and the main components are carbonate, sulfate, silica, alkali- and alkaline earth metals. The pH may drop 1-2 units below normal and the water may contain some hydrogen sulfide, but for the most part it is not detected. Iron has been recorded in concentrations exceeding 9 mg/l and commonly 2-3 mg/l (Kristmannsdottir et al. 2000a). Sample treatment has been variable and this may explain some but not all variation. In the few samples in which mercury has been detected it is in very low concentrations. The con42
centrations of some other heavy metals will increase relative to their normal water concentrations, but data is only available from the last 2-3 jokulhlaups. The jokulhlaups known to be connected to volcanic events do not distinguish themselves much in the chemistry of water, but some correlation may be seen in increasing concentrations of sulfur species and iron. As the lake is large and takes several years to fill up the timing of a volcanic event in relation to the water level will have a great effect on the reactions in the lake and subsequently the composition of the jokulhlaup water. The giant jokulhlaup following the eruption in Vatnajokull in 1996 was the best studied in every way so far, also geochemically (Gislason et al. 1997). Besides monitoring all the rivers draining the eruption site below Vatnajokull samples were also collected from the eruption site itself two and three months after activity ceased (Table 1). When the eruption started the lake was almost empty and it took about a month to fill up by the meltwater from the eruption site. The composition of the water was not very different from water from previous jokulhlaups. The most pronounced difference is shown in much higher concentrations of sulfur species this time, mainly in the form of sulfate and thiosulfate. Fluoride concentration was also fairly high, whereas chloride, alkali metals and magnesium were found in somewhat lower concentrations than in former jokulhlaups. It was deduced from the calculated carbonate equilibrium that the water has mostly reacted at rather high temperatures, up to 80 "C at the eruption site (Gislason et al. 1997) and then the water had flowed towards the Grimsvotn lake, melted glacier ice and cooled down undenvay. When the lake was filled up and the water left the lake the water temperature was probably about 5-10 "C. The water from the eruption site (Table 1) distinguishes itself mainly by higher concentrations of sulfur species, fluoride and silica than recorded from any previous jokulhlaup waters. With reference to the concentration of S, C1 and F in the water the pH in the meltwater may have been as low as 2.7 prior to reaction with the volcanic tephra. Samples of lake water in Grimsyotn obtained from drillholes through the ice cover (Agustsdottir et al. 1992, Agustsdottir & Brantley 1994) show varying concentrations by depth and for the two different drillholes but the concentration of most elements fall within the range of concentrations in jokulhlaups as shown in Table 1. In Table 1 an analysis, not previously published, of a hot spring which appeared on the edge of the Grimsvotn lake after a volcanic eruption in 1983 and was sampled in 1984 is shown. The water phase of the spring is shown in the table, but at the same time a gas phase which bubbled up with the spring water and through open water at the edge of the lake was also sampled. This gas contained 99.1 % (vol.) C02, about 0.02 % H2S and 0.6
% C&. Calcium, SO4, C1 and total carbonate con-
centrations for the spring sample show values within the limits for the jokulhlaup water. The concentration of H2S is very low, but the sample was not analysed for H2S upon sampling and it may have been oxidized during storage. The concentration of total carbonate is similar to that of jokulhlaup water and the gas phase accompanying the water was almost pure Col. Silica concentration is about five times higher than the highest one recorded in any of the jokulhlaup waters. Sodium concentration is about double of the jokulhlaup water and both potassium and magnesium concentrations considerably higher than in any jokulhlaup water encountered so far. Fluoride concentration is triple to that of previous jokulhlaups and about the same as water collected from the 1996 eruption site. The pH of the spring water on the other hand is just about 7, which is relatively high compared to most of the jokulhlaup waters. The concentration of the spring water is very similar to that of Icelandic carbonated springs of similar temperatures (Arnorsson 1982, Kristmannsdottir 1992). The chemical composition of the flood water (Kristinsson et al. 1986, Kristmannsdottir et al. 1999) from the small cauldrons not far from the Grimsvotn subglacial lake is very similar to that of jokulhlaup water from Grimsvotn, but less mineralised. The waters generally also have a lower pH, and higher concentrations of sulfur species. In previous times giant floods due to volcanic events in Vatnajokull have occurred from the rivers draining to the north of Vatnajokull (Sigbjarnarson 1990, Tomasson 1990, Thorarinsson 1950), especially Jokulsa a Fjollum. During the 1996 eruption in Vatnajokull a small part of the drainage from the eruption site was detected in Jokulsa a Fjollum (Kristmannsdottir 1999), but no major flood occurred. The chemical changes observed were an increase in total dissolved solids, conductivity and total carbonate. Traces of sulfide and mercury, which have never been observed at normal conditions were detected. The stable isotope ratios (6D and 6**0) were found to be abnormal and lower than measured in precipitation anywhere in Iceland. The excess 6D is extraordinary low, the 13C values are also abnormal and 14C age appears to be very high. The pH of the water is near to normal, which may reflect the fact that the flowpath is very long giving the water opportunity to react with sediments carried by the water as well as the atmosphere. As pointed out above many of the rivers draining the glacier Myrdalsjokull show a more or less constant geothermal influence. In 1999 there was a significant flood in one of the rivers probably due to due to a suspected volcanic event (Kristmannsdbttir 1999). Samples were obtained late in the flood, and showed a highly increased mineralization, mainly as
43
bicarbonate, alkali and alkaline earth metals and to a lesser degree sulfur species.
atmosphere. As pointed out by Gislason et al. (1996) the release of nutrient salts and introduction of Fe, Mg and other metal fluxes to the ocean may have implication for the short term carbon cycle as planktonic growth rates are determined both by nutrients and the level of metabolically important metals.
6 DISCUSSION Geothermal or volcanic influx into glacial meltwater will enhance reactions with basaltic tephra. During volcanic activity the amount of tephra suspended in the meltwater will be very high and increase reaction rate even more. The interaction of volcanic or geothermally acidified glacial meltwater with basaltic tephra results in a chemical composition similar to carbonated hot springs elsewhere in Iceland as demonstrated by the spring water collected on the edges of the Grimsvotn lake just after the 1983 eruption. At normal conditions such water and an accompanying gas phase would enter the lake, be diluted by meltwater and react with the suspended tephra "sediments" and available atmospheric air. The subsequent reactions with atmospheric air will raise the pH and vent out some of the volatiles like C02, H2S and mercury. As long as the system is confined within a subglacial lake the access to atmospheric air will be limited and the water will be undersaturated with oxygen. The temperature reached by the system will have great influence on the reaction pattern within the system. Precipitation of carbonate may follow a sudden heating of the meltwater (Gislason et al. 1997) whereas reaction at lower temperatures may induce its dissolution. It is assumed that during subglacial eruptions water temperatures are much higher than in the circulation systems in glacial lakes like Grimsvotn. During the 1996 eruption in Vatnajokull (Gislason et al. 1997) temperatures as high as 80 "C may have been reached at the eruption site and the composition of the water may have been modified by the precipitation of carbonates. The higher temperatures are also indicated by e.g. a higher concentration of silica and a lower magnesium concentration than observed in jokulhlaup waters. Very high concentrations of sulfur species as well as fluoride in the sample from the eruption site are signs of the volcanic influx. When the water is freely exposed to atmospheric air in the flowpath towards the sea subsequent reactions with atmospheric air will raise the pH further and vent out most of the volatiles with subsequent reactions with the suspended sediments involving exchange of protons and metals as well as oxidation of sulfur species and formation of bicarbonate ions. Such reactions will be the more efficient the longer the subaerial flowpath, the larger the sediment load and are also influenced by temperature and type of flow. Typically the pH in rivers with a long flowpath will be generally higher than in those with a short one (Kristmannsdottir et al. 1999,2000a). Later when the floodwater is released to the ocean it may cause an increase in the fixation of CO2 in the
7 CONCLUSION Contrary to subaerial volcanic eruptions where great amounts of gases, other volatile elements, as well as volcanic ash are released to the atmosphere almost all those emissions are dissolved and suspended in the meltwater during subglacial eruptions. The interaction of volcanic and geothermal activity with big glaciers is a localized phenomenon almost confined to Iceland. Locally the correct interpretation of the chemical composition of a glacial river sample may be of the utmost importance in order to tell if there is an imminent danger of a volcanic eruption or whether there is a jokulhlaup from a geothermal lake or a cauldron. The complex reactions resulting from such interactive processes may have a rather general effect and induce significant global consequences by the influence of the meltwater discharged to the ocean on the fixation of CO2 in the atmosphere. ACKNOWLEDGEMENTS The Public Roads Authority did support work on the project concerning the development of an automatic warning system for volcanic eruptions and extreme floods as well as most other Icelandic research projects on volcanic and glacial interaction. The Icelandic Technical Science Fund and the Natural Disater Relief Insurance Fund also supported the project.
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Sugisaki, R. ang Augustithis, S.S. Eds.). Theophrastus, Athens: 187-222. Amorsson, S., Sigurasson, S. & H. Svavarsson 1982. The chemistry of geothermal waters in Iceland. I. Calculation of aqueous speciations from 0 to 370 “C. Geochim. Cosmochim. Acta 46: 1513-1532. Bames, I. 1984. Volatiles of Mount St. Helens and their origins. J. Volcanol. Geotherm. Res. 22: 133-146. Bjornsson, H. 1974, Explanation of jokulhlaups from Grimsvotn, Vatnajokull, Iceland. Jokull24: 1-26. Bjomsson, H. 1992. Jokulhlaups in Iceland, prediction, characteristics and simulation. Annals of Glaciology 16: 95-106. Bjomsson, H. & M.T. Gudmundsson 1993.Variations in the thermal output of the subglacial Grimsvotn caldera, Iceland. Geophysical Research Letters 20: 2 127-2 130. Bjornsson, H. & H. Kristmannsdottir 1984. The Grimsvotn geothermal area, Vatnajokull, Iceland. Jokull34: 25-50. Bjomsson, H. & E. Palsson 1990. Volcanoes beneath Vamajokull, Iceland. Evidence from radio echo-sounding, earthquakes and jokulhlaups. JokuIl40: 147-168. Bjomsson, H., Bjomsson, S. & Th. Sigurgeirsson 1982. Penetration of water into hot rock boundaries of magma at Grimsvotn. Nature 295: 580-581. Bjomsson, H., Palsson, F. & M.T. Gudmundsson 1997. “Jokulhlaup on Skeidarhsandur from the view of Grimsvotn and Skeidarh glacier” (In Icelandic). Proceedings of Symposium of the volcanic eruption in Vatnajokull 1996, Icelandic Geological Sociely: 17-18. Brown, G.H., Sharp, M.J., Tranter, M., Gurnell, A.M. & P.W. Nienow 1994. Impact of post-mixing chemical reactions on the major ion chemistry of bulk meltwaters draining the Haut Glacier d’Arolla, Valais, Switzerland. Hydrological Processes 8: 465-480. Casadewall, T.J., de la Cruz-Reyna, S., Rose, W.I., Bagley, S., Finnegan, D.L. & W.H. Zoller 1984. Crater lake and posteruption hydrothermal activity, el Chichon volcano, Mexico. J. Volcanol. Geotherm. Res. 23: 169-191. Cox, M.E. 1983. Summit outgassing as indicated by Radon, Mercury and pH mapping, Kilauea volcano, Hawaii. J. Volcanol. Geotherm. Res. 16: 13 1- 151. Edner, A.H., Faris, A.G.W., Sunesson, A.A., Svanberg, A.S., Bjamason, J.O., Kristmannsdottir, H. & K.H. Sigurasson 1991. Lidar Search for Atmospheric Atomic Mercury in Icelandic Geothermal Fields. J. Geophys. Res. 96: 29772986. Einarsson, P. & B. Brandsdottir 1997. “A geological frame and preceedings of the sublglacial volcaninc event in Vatnajokull 1996”. (Haraldsson, H. Ed.), Vatnajukull-Eruption and Jokulhlaup in 1996 (in Icelandic). The Public Poads Administration of Iceland, Reykjavik: 9- 17, Elefsen, S., Snorrason, A., Haraldsson, H., Gislason, S.R. & H. Kristmannsdottir 2000. Real time monitoring of glacial rivers in Iceland. Proceedings Extremes of the extremes, IAHS publ., 271(Snorrason et al. Eds.): In press. Gerlach, T.M. & T.J. Casadewall 1986. Fumarole emissions at Mount St. Helens volcano, June 1980 to October 1981: Degassing of a magma-hydrothermal system. J. Volcanol. Geotherm. Res. 28: 141-160. Giggenbach, W.F. 1992. The composition of gases in geothermal and volcanic systems as a function of tectonic setting. Proceedings Water-Rock Interaction, (Kharaka Y. K. & Maest A. S. Eds.), 873-878, Rotteredam: Balkema. Gislason, S.R., AndrCsdottir, A. & A. Sveinbjomsdottir 1992. Local effects of volcanoes on the hydrosphere: Example from Hekla, southern Iceland. Proceedings Water-Rock Interaction, (Kharaka Y. K. & Maest A. S. Eds.), 477-481, Rotteredam: Balkema. Gislason, S.R., Amorsson, S. & H. Armannsson 1996. Chemical weathering of basalt in southwest Iceland: Effects of
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sen 2000a. Geochemical warning for subglacial eruptions. Background and history. Proceedings Extremes of the extremes, , IAHSpubl., 271 (Snorrason et al. Eds.): In press. Kusakabe, M., Ohsumi, T. & S. Aramaki 1989. The lake Nyos gas disaster: chemical and isotopic evidence in waters and dissolved gases from three Cameroonian crater lakes, Nyos, Monoun and Wum. J. Volcanol. Geotherm. Res. 39: 167185. Lawler, D.M., Bjomsson, H. & M. Dolan 1996. Impact of subglacial geothermal activity on meltwater quality in the Jokulsa a Solheimasandi system, southern Iceland. Hydrol. Proc. 10: 557-578. Olafsson, M. & K. Kristmannsdottir 1989. The influence of volcanic activity on groundwater chemistry within the Namafjall geothermal system, North Iceland. Water-Rock Interaction, ( Miles, D.L. Ed.), 537-540, Rotteredam: Balkema. Muffler, L.J.P., Hedenquist, J.W., Kesler, S.E. & E. Izava 1992. Japan-US seminar on magmatic contributions to hydrothermal systems. Report no. 279, Geological Survey of Japan, pp. 1-6. Palsson, S. & G.H. Vigfiisson 1996. The database of sediment 1963-1995. Report OS-96032NOD-05 B Reykjavik: Orkustofnun (In Icelandic) Palsson, S., Zophoniasson, S., Kristmannsdottir, H. & P. Jonsson 1999. The first jokulhlaup from Grimsvotn in 1996. Report OS-991 15 Reykjavik: Orkustofnun (In icelandic),. Saemundsson, K. 1979. Outline of the geology of Iceland. Jokull29: 7-28. Sigbjamarson, G. 1990. Jokulhlaups and their flow paths (In Icelandic). Vatnid and Landid: Orkustofnun, 129- 143. Sigvaldason, G. 1963. Influence of geothermal activity on the chemistry of three glacial rivers in southern Iceland. Jokull 13: 10-17. Sigvaldason, G. 1964. Some geochemical and hydrothennal aspects of the Askja eruption. Beitr. Mineral. Petrograph. 10: 263-274. Sigvaldason, G. 1965. The Grimsvotn thermal area. Chemical analysis ofjokulhlaup water. Jokull 15: 125-128. Steinthorsson, S. & N. Oskarsson 1983. Chemical monitoring of jokulhlaup water in Skeidara and the geothermal system in Grimsvotn, Iceland. Jokull33: 73-86. Sveinbjornsdottir, A.E., Amorsson, S., Heinemeier, J. & E. Boaretto 1998. Geochemistry of natural waters in Skagafjordur, N-Iceland, 11, Isoptopes. Proceedings WaterRock Interaction 9, (Arehart, G. B. and Hulston, J. R. Eds.), 1431-1434, Rotteredam: Balkema. Thorarinsson, S. 1950. “Jokulhlaup and volcanic eruptions in the river basin area of Jokulsa a Fjollum“ (In Icelandic). Nattlirufredingurinn 20: 113-133. Thorarinsson, S. 1959. “The possibility to predict next eruption of Katla” (In Icelandic with English summary). Jokull9: 618. Thorarinsson, S. 1967. Hekla and Katla. The share of acid and intermediate lava and tephra in the volcanic products through the geological history of Iceland: Iceland and MidOcean Ridges. Soc. Sci. Zslandica 32: 190-199. Tilling, I. & B.F. Jones 1996. Waters associated with an active basaltic volcano, Kilauea, Hawai: Variation in solute sources, 1973-1991. Geol. Soc. Amer. Bulletin 108: 562577. Tomasson, H. 1990. Glaciofluvial sediment transport and erosion. Arctic Hydrologv, Present and Future Tasks. Report no. 23, (Gjessing, Y. ed.) Norwegian National Committee for Hydrology: Oslo, 28-36. Trabant, D.C., Waitt, R.B. & J.J. Major 1994. Disruption of drift glacier and origin of floods during the 1989-1990 eruptions of Redoupt volcano, Alaska. J. Volcanol. Geotherm. Res. 62: 369-385.
Varekamp, J.C. & P.R. Buseck 1983. Hg anomalies in soils. A geochemical exploration method for geothermal areas. Geothermics 12: 29-47. White, D.E. & G.A. Waring 1963. Volcanic emanations. Data of Geochemistv, 6Ih ed. U. S. Geol. Sum. Pro$ Paper, 440K: 1-29. Zophoniasson, S. & S. Palsson 1996. “Flow in Jokulhlaups from the Skajici cauldrons, sediment concentration and TDS’. Report OS-96066NOD-07 Reykjavik: Orkustofnun (In icelandic).
46
Water-Rock Interaction 2001, Cidu (ed.), 02007 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The inverse modeling of water-rock interaction G .Ottonello Permanent Earth Sciences Technical Observatory at DIPTERIS, University of Genova,Corso Europa 26, 16132 Genova,Italy
ABSTRACT: Interpreting water chemistry in the framework of reaction kinetics is a complex procedure that involves three distinct (and superimposed) inverse model approaches: the first one devoted to the appraisal of the continuum, the second one to the (local) resolution of reaction surfaces and the third to the evolution of these last parameters in a time mode scale. The first two steps imply modifications of the informational content of the system, which may be properly analyzed in terms of marginal probability density in the space of the model. This analysis for the third step is more difficult, due to the usual non linearity between model parameters and observables.
1 INTRODUCTION
for an appropriate treatment of inverse problem theory: Parameterization of the system: "Discovery of a minimal set of model parameters whose values completely characterize the system GfFom a given point of view)". Forward modeling: "Discovery of the physical laws allowing, for given values of the model parameters, of making predictions as to the results of measurements on some observable parameters". Inverse modeling: "Use of the actual results of some measurements of the observable parameters to infer the actual value of the modelparameters". In our conceptual framework, moreover, the term "system" acquires the precise thermodynamic significance assigned to it long ago by Lewis and Randall (i.e. 'I Whatever part of the real world which is the subject of a thermodynamic discussion"; Lewis & Randall 1970) and the characterization is implicitly complete. Some features of the inverse model approach to water chemistry are here outlined with the aid of a case study. The arising considerations are of general validity.
Probably "the inverse approach" is the most ''direct" way for a Geochemist to deal with the physicalchemical complexity of a water-rock interaction process. This statement is something more than a joke. Due to the different time scales of human lives and geological events, we are like visitors at a PhotoShop exhibition of past events. Instead of being fixed on images, events are translated into altered mineralogical parageneses (or altered water chemistries) each one pertaining to a given time (or time scale, when moving from macroscopic to microscopic). Also, due to the different size scales of human beings and geological formations, we are like probes that move here and there and sample appropriate points of the physical system. How "appropriate" are these points is the central question of the inverse approach. To appreciate this fact we must turn onto more rigorous arguments and give mathematical definitions of the conceptual framework on whose basis we must operate. The following definitions (and much more) are taken from Tarantola (1987) to whom readers are referred
47
2 MODEL SPACE AND MODEL CHART
the covariance between chemical data (albeit translated into activity terms). The theoretical probability density O(d,m) is a joint density function that describes the expected physical correlation between model and observable parameters. Practically, O(d,m) represents the state of information obtained solving the forward problem by an "inexact" physical theory. For an exact theory, denoting g(m) the generic function of model parameter that leads to a calculated value (deal): dcal =g(m) (31 the conditional probability (i.e. referred to the state of null information) may be described by a delta Dirac function, i.e.: @(dim)= 6(d - g(m)) (4) with
Denoting with n the number of model parameters, to each point in the model space N it may be associated a point in space M, isomorphic to a part of R" (with R = real line); the model space N is a n-dimensional nonlinear differentiable manifold (i.e., a topological space whose local coordinate system is obtained from differentiable analytic functions of those in euclidean space) and M is a "chart" of N (Tarantola 1987). Along this line of reasoning we may equally well define a data space D which collects all observables (i.e. results of measurements) and a joint space X = DxM (parameter space) whose elements are constituted by the couples x = (d,m), with d = observables (data) and m = model parameters. In this acceptation, solving the inverse problem means simply to translate the information arising from the data space D (actually from the data chart D) into the model space M1 . Practically this is achieved by a combination of theoretical, aprioristic and experimental informations. The arising "a posteriori" information is represented by the probability density CY (d,m) (Tarantola and Valette, 1982): o(d, m) = P ( 4 m>o(d, m) m)
44
@(dlm)=O(d.m) %l(m)
(5)
which gives null probability to d # g(m) and probability 1 to d = g(m). For an inexact theory, the conditional probability function depends on the errors associated to each model parameter. Conceptually thus, estimating O(d1m)means to assign "error bars" to an otherwise exact theoretical relation d = g(m). It must be noted incidentally that equation 5 is quite general since the error bar associated to the ith predicted datum depends on the current value of m. Stemming from the above definitions, the solution of the general inverse problem is represented by the "a posteriori" information in the model space M, represented by the integral over all the d observables in the data space D of their respective individual probability densities ci (d,m) (eqn. 1.65 in Tarantola 1987):
(1)
where p(d,m) is the prior probability density, O(d,m) is the theoretical probability density and o(d,m) represents the state of null information. The prior probability density p (d,m) represents the joint prior information obtained in the joint space X when combining the density functions arising from the information on the data space D (observables, d) and on model parameters (M,m):
dd,m)= P D (d>pM (2) By definition, the a priori information on model parameters is independent of observations, but often part of the a priori information is indeed obtained from a rough analysis of the data set and the general probability density p(d,m) does not satisfy strictly equation 2 since does exist correlation between d and m. As it will be shown later on, this is our case, since the "approach to continuum" implies the choice of continuous differentiable functions in the reaction space whose conformation depends upon
3 THE DATA SPACE AND THE MODEL SPACE OF THE WATER - ROCK INTERACTION PROCESS 3.1 Choice of the data space axes Although the net result of irreversible exchanges taking place whenever water flows through mineral grains (reactive transport) is in term of exchanged mass, the most convenient data space we may adopt when describing the effects of reactions taking place within an heterogeneous system is that one based on chemical potentials (p). There are two main (both logical and historical) reasons to state this:
note that, by linguistic abuse, the term model space is synonymous of model chart in the current literature
48
i) For a system initially in a perturbed state (disequilibrium) and evolving under constant temperature (T) and Pressure (P) conditions, reactions among components proceed spontaneously in the sense of Gibbs free energy minimization: system = C p i n i +min (7)
T are two essential parameters. However, with the adopted data parameterization, the effects of P and T are embodied in the numerical value of the activity terms and P, T do not belong (at least explicitly) to the data space. As already stated, reactions proceed spontaneously in the sense of an internal entropy production. However the rates of attainment of equilibrium through elemental exchanges (i.e. reaction rates, see later) are not identical for all components in all phases. Hence also time (t) is an essential parameter. However, as we will see later on, time is better represented in a fractional context, identifying with 0 the fractional time at the beginning of the water-rock interaction process and 1 the fractional time at completion of exchanges. The overall degree of advancement of the process (6) is thus transformed into a fractional overall degree of advancement of the process Marini et al. 2000) which is the main axis (at least conceptually) of the data chart. The space of data chart encompasses thus only the field of (adimensional) real positive numbers, which renders the computation of model parameters less complicate.
with ni = number of moles of the ith component.
ii) Irreversible exchanges taking place locally in a Representative Elementary Volume (REV) of the system (see later on) are conceived as due to a positive amount of internally produced entropy (diS, Prigogine 1961):
(c;
with vi = stoichiometric reaction coefficient of the ith component and dk = infinitesimal overall advancement of the process. It must be noted that, since chemical potentials can be expressed in several ways, equation (9) is more general than equation (8). For instance, operating under constant internal energy (U) and volume (V) we may eventually write:
3.2 Choice of the model parameters Basically, any water-rock interaction process may be reconducted to a set of independent reactions taking place between the minerals present in the rock matrix (both pre-existing or newly formed phases) and the permeating solution. The model parameters are thus almost invariantly represented by kinetic reactivity terms, with instantaneous equilibrium constituting the particular case of detailed balancing between forward and backward velocities (cf. Lasaga 1981; Ottonello 1997). A fundamental assumption to be made from the very beginning to address the complexity of natural systems concerns the idealization of the chaotic mineral-water interfaced structure as a porous medium composed of an aggregate of mineral grains and pore spaces permeated by a pore fluid in local equilibrium. This is reconducted to a connected set of REVs (Bear 1972; Lake 1989) which characterize locally the properties of the system. In this continuum representation the physical-chemical variables are locally averaged within a REV and continuous at the macroscale of the REVs connection. As shown in Figure 1 we may conceive in this way the locally reversible - overall irreversible exchanges between water and rock matrix as a non-zero summation of relative reaction rates.
and so on. Although the model space t.E is defined intrinsically (i.e., without reference to any particular system of coordinates), for numerical computation it is necessary to represent each point of the data space D by its coordinates in a given coordinate system. Chemical potentials are thus referred to a particular state of reference (standard state) and the adopted parameterization of the system composition (observables) is thus in terms of "thermodynamic activity" of components (aiYa,p,~): /
n
\
with pea= standard state potential of the ith component in the a phase; R= gas constant and T= absolute temperature. Since the intensive variables P,T affect the local value attained by the chemical potentials, then P and
49
where Ei is the relative reaction rate of the ith reaction, vj ,i indicates the stoichiometric coefficient of the jth solute in the ith reaction, and 5 represents the bulk (i.e. non-fractional) overall progress variable. Moreover, the relative reaction rates are related to the absolute rates through:
where Rovera11 is the overall reaction rate of the bulk process (Delany et al. 1986):
Figure 1. Schematic representation of an irreversible exchange trend regarded as a sequence of internally buffered local equilibria. Each discrete point corresponds to a condition of detailed balancing in a reaction progress mode with identical forward and backward reaction rates. In this condition the overall affinity to equilibrium is zero. However, a vector connecting two neighboring points corresponds to an irreversible exchange associated with a nonzero summation of relative rates (from Sciuto & Ottonello 1995b; note that symbol CJ in Figure 1 corresponds to symbol E in the present context).
overall = C l R i l = CSiI'il I
I
i=Cil
"5
(18)
I
The important fact ensuing from the above relations is that:
The driving force that moves the composition along the chemical path of the connected REVS is the overall affinity Aoverall of the chemical reactions (De Donder 1920; Prigogine 1961; see also equation 9): I
(16)
1
Equation (19) emphasizes that the essential model parameters are the specific reaction rates ( E i ; defined per unit surface) and the reaction surfaces Sj of the various minerals present in the solid aggregate. It is worth to stress now that the specific reaction rates Ri are themselves function of thermodynamic activities of solutes. In fact, the Transition State Theory (TST) representation of laboratory measured rate constants of minerals, written in terms of forward (dissolution) parameters takes the form (Aagaard & Helgeson 1982):
U5
In this conceptual framework, the compositional trend of the fluid can be reconstructed through a sequence of discrete points representing locally reversible and overall irreversible mass exchanges (Sciuto & Ottonello 1995a,b; Marini et al. 2000). These irreversible mass exchanges can be quantified by the following formulation, for a given component
where k+,i is the dissolution rate constant of the ith reaction, q+,i is its kinetic activity pr duct, and A+,i is the thermodynamic affinity of the it' reaction:
where n i j stands for the molality of the jth aqueous solute in the ith reaction, represents the fractional overall progress variable and
<
1
A+,l=-RTln -+2 .
(21)
Q+,i in equation (21) is the ion activity product (or react'on quotient, cf. Aagaard & Helgeson 1982) of the itL reaction and q is usually defined a Temkin's average stoichiometric number of the it' reaction (Aagaard & Helgeson 1982; Oelkers 1996; but see also Denbigh (1971) for a correct appraisal of the significance of in a macroscopic notation).
50
Glynn & Brown 1996) the forward equation may be eventually linearized around a reference model (mo):
Alternatively, reaction rates are often described as the activity term rate laws. For example the reaction rate of calcite has been represented as: R, = kl,,[H’] +k2,i[H2C03]+k,,i[H20]+...... (22) where terms in square brackets are thermodynamic activities and the various rate constants k l , k2, k3, ... depend exponentially on temperature (Plummer et al. 1978; Busenberg & Plummer 1982; Chou et al. 1989). We see thus that, in both cases (Q+,i in eqn. 21; square brackets in eqn. 22) the activity terms do condition the local value attained by the specific reaction rate. Obviously then, when expanding equation (19) by combining it with a certain theoretical appraisal of the intrinsic significance of R i (i.e. equations 20-21 or, alternatively, equation 22) it turns out that the specific reaction rates are influenced by the prior information (activity terms) and only the reaction surfaces Si are true unconstrained model parameters. Unfortunately the concept of “reaction surface” is still under debate. Its current appraisal is largely phenomenological and vitiated by some inconsistencies which cannot be exhaustively treated in the present context. In treating a case study as application of the above outlined principles we will keep thus consistent with the mathematical significance of our operations, disregarding some of their physical-chemical implications.
d,, = gb-4 g(m, ) + G 0 (m - mo (23) with the linear operator Go representing the derivative of g at m = mo. Concerning the second question, rarely the data population is sufficient to allow a meaningful statistical analysis. Moreover, often the data are themselves a (non-linear) thermodynamic manipulation of true observables (chemical analyses). I will deal here with a case study where the data population is log-normal; i.e. has a normal (Gaussian) distribution in a logarithmic space. The third question is the most difficult to answer for a non mathematician. I will afford it with an “a posteriori” analysis of model results for our case study along two traditional guidelines: 1) a parameter is well resolved from the data set if the a-posteriori error bar is much reduced with respect to the a-priori one (more generally, if the a posteriori marginal probability density is significantly different from the a priori marginal probability density). 2) accepted that: a) the parameter ma represents the discretization of a continuous variable Y(t) (location in space and time); b) the covariance operator m = Y(t) adequately represents the dispersion from the mean; c) errors are correlated; then, the greater is the error correlation, the worst is the resolution (in time and space) of the dataset.
4 SOME PRELIMINARY CONSIDERATIONS IN THE RESOLUTION OF A CASE STUDY Mathematically, the first concerns one has to address when solving the inverse problem are the following: 1) Is the forward problem linear ? In other terms, does exist a linear operator G such that the data (d) may be linearly translated into model parameters ( d = G m)? 2) Is the distribution of data in the space of data D normal? 3) Are the probability densities used to describe observational uncertainties, forward modeling uncertainties and prior uncertainties on model parameters Gaussian ?
Generally, being the problem non linear, I may anticipate that there is no explicit expression of the solution and the analytic approach I will follow in the second (crucial) step is the chi-squares criterion applied to a combined (Monte Carlo - Simplex Steepest descent) exploration of the model space.
5 INVERSE MODELING IN A CASE STUDY: THE BISAGNO VALLEY SPRING WATERS
5.1 Step one: the approach to continuum
Concerning the first point, since we have usually to deal with complex chemical paths in the space of reaction (see later on) the response is generally “no”. Only in the very special case of simple mass balance calculations (such as those discussed for instance by
The chemical evolution of the Bisagno valley spring waters, ranging from initial Ca-HC03 composition towards final Na-HC03 facies was investigated by
51
Marini et al. (2000). The aim. of the paper was to assess the possibility of reconstructing the surface area of the mineral phases dissolved (or precipitated) during the evolution of natural waters, using water
chemistry. The Bisagno Valley catchment was chosen because of its relatively homogeneous geological-hydrogeological setting, making it a simple natural case to test inverse problem techniques. The chemical path of the investigated spring waters is relatively well constrained. The simple plot of total Ca concentration vs pH (Fig. 2, upper part) is indicative of an evolutive trend from rain waters to “immature” Ca-HC03 waters, most of which appear to be slightly oversaturated with respect to calcite (this is typical of waters coming from carbonate-rich aquifers, Appelo and Postma, 1996), to “mature” Na-HC03 waters. Calcium-HC03 and Na-HC03 waters depict a continuous trend in the plot of Ca vs f& (computed from carbonate alkalinity and pH; Fig. 2, lower part), excluding the very “immature” samples BIS-72 and BIS-75.Inspection of the activity plot of log [ka2+/(a~+)2]vs log [a~~z+/(a~+)2] for carbonate minerals at 15°C and 1 bar (Fig. 3) shows that all the spring of the Bisagno valley, including the Na-HC03 waters BIS-13, SC-102 and SC-103, plot close to the phase boundary between calcite and ordered dolomite, whose location is little affected by temperature changes. In the activity plot of log (aNa+/aH+)vs log (aK+/aH+) for silicate minerals at 15°C and 1 bar (Fig. 4) most waters fall in a well defined alignment into the kaolinite field, whereas Na-HC03 waters are situated close to the invariant point of coexistence of K-feldspar, albite and muscovite. Continuous trends may be also envisaged in the activity plots of log [ ~ ~ 2 + / ( a ~vs + ) log ~] (aNa+/aH+)(Fig. 5) and of log [ a ~ ~ z + / ( a ~vs+ )log ~] (aNa+/aH+)(Fig. 6) for silicate minerals at 15°C and 1 bar. In these latter two activity diagrams Na-HC03 waters plot near the invariant points of stable equilibrium between either laumontite or 7 Aclinochlore, albite, kaolinite, and the aqueous solution for the pressure and temperature of interest. This does not mean that all these phases are actually precipitated in the natural system, but surely that NaHCO3 waters are representative of the condition of nearly completion of irreversible mass exchanges in the system. Here we are not concerned with all the subtilities of water chemistry (the bulk process can be thought as the result of continuous rock titration [dissolution of Al-silicates and carbonates] mainly driven by the conveision of CO2 to HCO3- see Marini et al. 2000, for an extended discussion) but rather with the informational content represented by the 135 water samples collected and analyzed by the Authors.
Figure 2. Correlation plots of (upper part) Ca concentration vs pH showing the springs of the Bisagno valley (circles = CaHC03 waters; triangles = Na-HCO, waters) and rain waters (stars). The two lines of positive slope refer to open-system calcite dissolution, under fco2 values of 0.03 and 0.003 bar. The two lines of negative slope refer to calcite saturation at 0” and 25°C. (Lower part) Ca concentration vs fco2 for the springs of the Bisagno valley (from Marini et al. 2000).
52
reaction path modeling calculations; cf. appendix 1 in Marini et al. 2000), the activity trends of the remaining solutes and the molality trends (obtained from activities by accounting for individual ionic activity coefficients) were readily derived by fitting the analytical data with polynomial expansions in I; based on the observed correlations in the activity diagrams of Figures 3 , 4 , 5 and 6 (Table 1).
Figure 3. Activity plots of log [kaz+/(a~+)2] vs log [ a ~ ~ z + / ( a ~ + ) ~ ] for carbonate minerals, at 15°C and 1 bar. Circles = Ca-HC03 waters; triangles = Na-HC03 waters. Paths of forward (x) and inverse (+) geochemical modeling are superimposed.
] (aNa+/aH+) for Figure 5. Activity plots of log [ a ~ , z + / ( a ~ +vs) ~log silicate minerals at 15°C and 1 bar. Simbols as in Figure 3.
Figure 4. Activity plots of log (aNa+/aH+) vs log (aK+/aH+)for silicate minerals at 15°C and 1 bar. Simbols as in Figure 3.
To achieve the "continuum limit" (the discrete set of open subsystems replaced with a continuous set of mathematical points labeled by the coordinate vector <; cf. Lichtner 1985) a simple linear conversion between fractional degree of advancement of the process and acidity of the system was first envisaged (equation. 3 in Table 1) and, stemming from this simple relationship (justified by preliminary forward
Figure 6 : Activity plots of log [aM,2+/(aH+)21"s log ( a ~ a + / a ~for +) silicate minerals, all at 15°C and 1 bar. Simbols as in Figure 3.
53
Table 1. Polynomial expressions relating thermodynamic activity of solutes, %, and molality, ni, to the fractional overall progress variable of the process, (Marini et al. 2000).
log asio,,
with x; ,U,o2respectively mean, standard deviation and variance of the normal population. The gain in information that Marini et al. (2000) obtained modeling the data around the interpolant functions (Table 1) may be conveniently represented in terms of Shannon (1948) analysis of information contents:
= -4.1826
log nsio,,aq = -4.1 826 loga,, = -7.2 - 2.06
log nH+= -7.1 6 19 - 2.03006 + 0.0307C2 log aNa+= -3.609 1 - 0.6893 6 + 1.89 166
+ 1.9305r2
lognNa+=-3.5656-0.72736
with fl(s) = probability density of the continuous distribution, and f 2(s)= probability density of the original distribution. Being the two populations log-normal, the information content may be readily derived by substitution of eqn. 27 in 28 and the opportune assignment of x i , (7, 0, terms. Results of this exercise are resumed in Table 2. The results listed in this Table may be regarded as an “overall” gain in information obtained assuming that activity paths are continuous differentiable in the space of reaction (4).We may appreciate that the informational gain is translated from the activity space to the space of composition with virtually no change (Table 2) and that the minimum gain of information is observed for Na’ . This last fact is particularly informative because it emphasizes the importance of activity diagrams in assisting the inverse modeling procedure. In fact, Marini et al. (2000) first computed log (aNa+/aH+)as a function of 4 and then utilized it as a pivot axis to obtain additional information on the activity distributions of K+, Ca2’ and Mg” through the activity-activity plots (Figs 4, 5, 6 respectively). The information added by activity-activity plots is thus missing for Na’ (which explains the limited increase of informational content for this element) but present
loga,, =-4.8208- 1.52986+2.6248c2 -0.2406c3 -0.6770c4 logn,, = -4.7759- 1.56974 i-2.66586’ -02406c3 -0.6770c4
log aca2+= -2.80 11 - 0.4277 6 log nca2+= -2.6352 - 0.5673 6 + 0.143 16
logaMgz+ = -3.8923
- 0.27846
lognMgz+ = -3.7349- 0.40576 + 0.1306C2 logaAi,+ = -13.2633- 6.0< log nM3+= -12.9175
- 6.27486
+ 0.28 18Y2
logaHcoj = -2.2582- 1.04176 + 1.0685c2
lognHco, = -2.2147- 1.07976 + 1.10746,
It is quite obvious that the manipulation of the initial activity data (obtained from speciation calculations based on the initial water chemistry) to make them consistent with the approach to continuum (linearization around a reference model) it is itself an inverse modeling procedure that affects the initial amount of information of the original dataset. To analyze to some extent the consequences of this preliminary operation in terms of information content I anticipate that either the distribution of the initial activity data and of the linearized data is lognormal. The corresponding probability density function is thus f(x)=---exp[
1
1
--$[log$)2]
Table 2. Shannon’s relative information content attained with the approach to continuum.
(26)
(2n)ll’s x
with xo=mean s = standard deviation of the log-normal population s2= variance of the log-normal population.
(9
Denoting x* =plog
-
with p = log,lOand y
=
x;, -3.543 -3.507 -4.993 -4.955 -3.988 -3.853 -2.948 -2.807
1
(activity data are adimensional) transforms f(x) in a normal distribution f(x*) with density f(x*)=----..p[-7(x* 1 1 -x;,)’] (27) (27$12
(7
20
54
(3
x;,
ci
0.263 0.265 0.104 0.108 0.057 0.061 0.088 0.092
-3.537 -3.501 -4.970 -4.932 -3.993 -3.861 -2.960 -2.822
0.308 0.311 0.350 0.352 0.262 0.271 0.242 0.252
0.033 0.033 0.635 0.6 19 1.OS 1 1.059 0.624 0.623
for K+, Ca2' and Mg2+. Let us consider again the activity-activity plots of Figures 3, 4, 5 , 6. It is quite obvious that there are four variables (i.e. log ma']/[H+], log [K+]/[H?], log [Ca2']/[H'I2, log [Mg2']/[H'I2) correlated by four interpolating functions. The information is thus redundant (at least in terms of activity ratios). In fact the trend depicted in Figure 3 (crosses) does not represent the best interpolant (maximum likelihood) of the data population but it is rather the information translated from the other three plots. We may regard thus this trend as effectively "theoretical" with respect to the diagram population and may analyze properly its m o u n t of marginal density information. It is rather evident that the (exact) theoretical relationship between the two variables is linear and that a linear operation relates thus data (d) to model parameters (m) (d = G m). We do not know much about the errors affecting each datum but we know that these affect both axes (we could assume that the two repeated analyses of sample BIS 13 give us an idea of a typical a standard deviation, induced by both analytical uncertainties and temporal effects; cf. Figures 2 to 6). Assuming that the null information probability density on model parameters and on data parameters is constant and that we have no a priori information on model parameters, i.e. :
Figure 7. Posterior probability density of the approach to continuum. Contourings are in logarithmic scale. Minima of the misfit function 50 are also localized for all samples (1) and CaHC03 waters (2), for comparative purposes.
versus log [Mg2']/[H']2 activity plot, obtained assuming for simplicity negligible errors on the x axis and 0 = 0.035xdOt,, for uncorrelated Gaussian errors. We see that the maximum probability density of the theoretical trend (identified by logarithmic contourings in the model chart of Figure 7) does not coincide with the maximum likelihood of the whole samples population (point 1 in Figure 7) and that things ameliorate if we discard mature Na-HC03)= waters in the data population (point 2 in the same Figure). This fact deserves a further analysis in terms of Bayes factors. A Bayes factor may be conceived as the ratio of the probability densities relating two alternative hypotheses. In our notation this may be presented in terms of conditional probability densities as:
w M(m) = const.
a,,(d) = const. (29) p M(m)= oM(m) = const. then the prior probability density representing the information we have on the theoretical (exact) relationship between d and m is (cf. equation 4) :
f MID (rn ' 'Id, I "'
and the marginal probability density representing the posterior information on model parameters is (cf. equation 6) oM(m)= do(d,mbd=exp
if the errors alternatively
-did)2
2
i
o2
1
B=F-JqQ i.e., for a given do,
(33) f MID ' , f MID " are two distinct
conditional probability densities in m. In our particular case the two alternative hypotheses are that the given datum do pertain or not to the space of the model that defines the theoretical trend. To solve the problem we must first compute elliptical contours of equal probability density (Meyer 1975):
(31)
are uncorrelated Gaussian, or
if the errors are uncorrelated exponential (Tarantola 1987). Figure 7 shows the contouring of the posterior probability density computed for the log [Ca2']/[H']2
k in equation 34 is a parameter that characterizes the contour, oxand oyare the variances of the sample on
55
x and y axes, xo and yo are (do) sample’s coordinates and r is the correlation coefficient of the data population. Stemming from the fact that the theoretical trend depicted in Figure 3 is linear: y = a + bx = 2.0755 + 0.91 19x
Table 3. Coordinates and uncertainties of selected data. Values in italics are relative to the final extended forward computation (6 1 1 ; see text for explanations). sample
(35)
BlS13 SC102
equating the partial derivative in dx of equation (35) to the partial derivative of the elliptical equation (34), the points of tangency between elliptical contour and theoretical trend are readily obtained by solving the following system:
SC103 BlS72 BIS75
b20: -2rbo,oy + o i y, = a + b x ,
(37)
(Jx
oy
r
14.062 14.062 14.267 14.267 14.162 14.162 10.438 9.734
14.684 14.684 14.638 14.638 14.697 14.697 10.822 10.011
0.30 0.30 0.28 0.28 0.29 0.29 0.22 0.20
0.42 0.42 0.43 0.43 0.44 0.44 0.31 0.29
0.9416 0.9416 0.9416 0.9416 0.9416 0.9416 0.9416 0.9416
a
b
k
B
BIS13
2.0755 -171.02 2.0755 -171.02 2.0755 -171.02 2.0755 2.0755
0.9119 13.082 0.9119 13.082 0.9119 13.082 0.91 19 0.9119
3.300 0.198 8.577 0.072 3.690 0.014 75.297 108.81
1.033 I . 770 0.198 1.911 0.762 1.982 3.7B-7 3.OE-11
SC 103
b o ~ ( a - y o ) + r o , o y ( a - b x-oy o ) + x 0 o i
Yo
sample
sc102 x, =
~0
BIS72 BIS75
(38)
The Bayes factor B of the datum do is then obtained through opportune simplifications as (Nicholls 1998):
The elliptical contour of tangency is kt =3.3 and the coordinates are xt=14.287; yt=15.104. Applying the above conditions to equation 38 results in a Bayes factor B=1.033. According to Jeffrey’s (1961) interpretation for ranges of B, for B > 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,
<
(41-5)
SiO, a SiO,,,,, Al,Si,O,(OH),
Values of the vector of solute concentrations, Equation (40), and the time dependence law (46), were substituted at each discrete value of 4 and the system was solved in terms of surface areas Si. Practical calculations were conducted by the authomated routine MINUIT (James & Roos 1977) minimizing the chi-squared summation:
(4 1-3) (4 1-4)
(I>
+ 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<
0 0 0 0 1 0
dCa 2'/d< dAI3'/d<
0 0 0 0 0 1
dSiOqq, /dC
0 0 0 1 0 0
and generalized as (Marini et al. 2000; cf. also Chilakapati 1995): (44) where I stands for the identity matrix (dimension Ntot xNtot , being Ntot the species over which the system is solved), n refers to the vector of solute concentration (length Ntot ), v is the stoichiometric reaction matrix (dimension Ntot XNr, with Nr = number of utilized reactions) and Rr, is the transposed reaction rate vector (cf.eq. 19), i.e. -
dt
RC,l =SjRi dr
with CD=: Cd + CT. Since the equation solving the forward problem of the second step is linear, in the least squares approach the maximum posterior probability density of the second step could be estimated computing the maximum likelihood point m M L through the misfit
(45)
57
function (equations 4.7a,b, 4.8, in Tarantola 1987; with modification): 1
dohs c-d s(m) =-[(g(m)2 = minimum
(dm)dohs )]
(49)
for m = mML.
Assuming errors to be uncorrelated the misfit function becomes effectively a "least squares" function (see Tarantola 1987; chapter 4 for an appropriate treatment):
Adopting the same notation, the misfit function (47) may be represented as:
'[
sf(m)=-
2
(51)
i€ID
and the probability density associated to function (51) may be expressed as follows (Spiegel 1961): f(x)= ...,t.xN-2exp[--$)
interpolant, however an investigation of the computed surfaces shows that they decrease exponentially with time (Fig.s 9, 10) and that the interpolation errors are uncorrelated (R2 = 0.024 for albite, 0.0189 for K-feldspar, 0.0508 for dolomite and 0.0677 for calcite). 5.3 Third step: the continuous representation of effective reaction surfaces Strictly speaking, this step could have been embodied in the second one (and indeed we already did it in the error analysis of step 2), with some risk ! The plot of effective reaction surfaces as a function of time indicates that all trends are non linear. Solving the third step means again to apply a least square method and to identify the maximum likelihood through a misfit function of type 50. Table 4 resumes this exercise and Figures 9,lO visualize the results. Errors are minimized through power interpolations or (in the case of K-feldspar) logarithmic interpolation.
(52)
with N = number of data. Equation (51) is significant only for N 2 5 (and becomes eventually Gaussian for a large data population). Being the data population in our case much reduced (N = 4), the probability density analysis through equation (52) in not significant and we must limit ourselves to analyze the error correlation. The system is thus solved in a stepwise fashion (i.e. at various discrete values of c), results are interpolated on 6 and interpolation errors are then plotted versus 4 to assess the existence of correlation. This procedure is rather rough since the error assessment depends upon the choice of the best
Figure 9. Evolution of the effective reaction surface of carbonate minerals with time in the Bisagno Valley case study.
Figure 10. Evolution of the effective reaction surface of silicate minerals with time in the Bisagno Valley case study.
(approach to continuum) involves the analysis of thermodynamic activity trends in logarithmic spaces (i.e. activity diagrams). The formulation of continuous differentiable functions in the space of reaction based on activity data implies a modification (generally an increase) of the informational content of the system which may be expressed in terms of Shannon's analysis. The assessment of the modification of informational content involved in the second step (local resolution of reaction surfaces) is more difficult and may be carried out only assuming that the model parameters (reaction surfaces) represent the discretization of a continuous variable Y(t) (location in space and time), that the covariance operator m = Y(t) adequately represents the dispersion fiom the mean and that errors are correlated. This analysis is prodromic to the third step: accepted that (as in the case study) the evolution of model parameter is non-linear, we must know abinitio (or alternatively build up, by a stepwise approach) the type of function that relates model parameters to the main axis of the model chart (time). Limiting the number of model parameters of the third step (e.g. constant reaction surfaces, for instance) means thus to limit the mathematical elasticity of the model with deep consequences on the accuracy of the forward computation which have been extensively discussed by Marini et al. (2000).
Table 4. Surface areas or reacting phases (cm2) vs time (years). Interpolating equations and r2 factor are listed. Phase
Equation
r2
Calcite Dolomite Albite K-feldspar
1nS = -2.01 55 - 1.7844 x In t 1nS = -3.8332-2.1064~ In t 1nS = 9.0012-0.3639xlnt S = 86.183 - 1 1.498 x lnt
0.9998 0.9961 0.9815 0.3980
Here, again, the model parameters of the previous step become the data population of the third one but the number of model parameters depends upon the "a priori" choice of the degree of the various polynomials. Limiting the number of parameters means thus to limit the mathematical elasticity of the model and a stepwise approach is thus almost compulsory. Performing now the complete forward calculation with the third step model parameters through the computer package EQ6 (Wolery & Daveler 1992) and extending the computational limits to >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.
<
ACKNOWLEDGEMENTS I am indebted with Alberto Tarantola for critical reading of a preliminary version of this manuscript and for his precious advice. I benefited of continuous enlightening discussions with Luigi Marini and Marino Vetuschi Zuccolini which are also gratefully acknowledged.
<
REFERENCES Aagaard, P. & H.C. Helgeson 1982. Thermodynamic and kinetic constraints on reaction rates among minerals and aqueous solutions. I. Theoretical considerations. Am. J. Sci., 282: 237-285. Appelo, C.A.J. & D.Postma 1996. Geochemistry, groundwater andpollution. Rotterdam: A.A. Balkema, Bear, J. 1972. Dynamics of Fluids in Porous Media. New York: Elsevier. Busenberg, E. & L.N. Plummer 1982. The kinetics of dissolution of dolomite in C02-H20 systems at 1.5 to 65 "C and 0 to 1 atm Pc02. Amer. Jour. Sci., 282: 45-78. Chilakapati, A. 1995. RAFT: A simulator for reactiveflow and transport of groundwater contaminants. PNL- 10636, Pacific Northwest Laboratory Report. Chou, L., Garrels, R.M. & R. Wollast 1989. Comparative study of the kinetics and mechanisms of dissolution of carbonate minerals. Chem. Geol., 78: 269-282.
6 CONCLUSIONS The apparently simple geochemical approach to the physical-chemical complexity of a water-rock interaction process is indeed the result of three superimposed inverse modeling steps. Each step involves modifications of the informational content of the system which may be quantified by analyzing the marginal probability densities in the data space D or in the space of the model M. The first step 59
De Donder, Th. 1920. LeCons de Thermodynamique et de Chimie-Physique, Paris: Gauthier-Villars. Delany, J.M., Puigdomenech, I. & T.J. Wolery 1986. Precipitation Kinetics Option for the EQ6 Geochemical Reaction Path Code. Livermore, CA: Lawrence Livermore National Laboratory, , UCRL-53642. Denbigh, K.G. 1971. The Principles of Chemical Equilibrium Third edition, Cambridge: Cambridge University Press. Glynn, P. & J. Brown 1996. Reactive transport modeling of acidic metal-contaminated ground water at a site with sparse spatial information. In Reviews in Mineralogy, Vol. 34, 13 1191, P.C. Lichtner, C.I. Steefel & E.H. Oelkers (eds.), Washington D.C.: Mineralogical Society of America. JefEeys, H. 1961. Theory of Probability. Oxford: Oxford University Press. Lake, L.W. 1989. Enhanced Oil Recovery, Englewood Cliffs, NJ: Prentice Hall. Lewis, G.N. & M. Randall 1970. Termodinamica, Roma: Leonardo Edizioni Scientifiche. Lichtner, P.C. 1985. Continuum model for simultaneous chemical reactions and mass transport in hydrothermal systems. Geochim. Cosmochim. Acta, 49: 779-800. Marini, L., Ottonello, G., Canepa, M. & F. Cipolli 2000. Waterrock interaction in the Bisagno Valley (Genoa, Italy): Application of an inverse approach to model spring water chemistry. Geochim. Cosmochim.Acta., 64: 2617-2635. Meyer, S.L. 1975. Data analysis for scientists and engineers. New York: John Wiley and Sons. Nicholls, J. 1998. Estimation of probabilities of three kinds of petrologic hypotheses with Bayes theorem. Math. Geol., 30: 817-835. Oelkers, E.H. 1996. Physical and chemical properties of rocks and fluids for chemical mass transport calculations. In Reviews in Mineralogy, Vol. 34, 131-191, P.C. Lichtner, C.I. Steefel & E.H. Oelkers (Eds) Washington D.C: Mineralogical Society of America. Ottonello, G. 1997. Principles of Geochemistry, New York: Columbia University Press, . Plummer, L.N., Wigley, T.M.L. & D.L. Parkhurst 1978. The kinetics of calcite dissolution in C02-water system at 5O to 60°C and 0.0 to 1.O atm C02. Amer. Jour. Sci., 278: 179216. Prigogine, I. 1961. Introduction to Thermodynamics of Irreversible Processes. 2nd revised edition, New York: Interscience Publishers. Sciuto, P.F. & G. Ottonello 1995a. Water-rock interaction on Zabargad Island (Red Sea), a case study: I) application of the concept of local equilibrium. Geochim. Cosmochim. Acta, 59: 21 87-2206. Sciuto, P.F. & G. Ottonello 1995b. Water-rock interaction on Zabargad Island (Red Sea), a case study: 11) from local equilibrium to irreversible exchanges. Geochim, Cosmochim. Acta, 59: 2207-2214. Shannon, C.E. 1948. A mathematical theory of communication. Bell System Tech. J., 27: 379-423 Spiegel, M.R. 1961. Statistics, New York: Mc Graw - Hill Tarantola, A. 1987. Inverse model theoiy: Methods of data fitting and model parameter estimation, Amsterdam, Oxford, New York, Tokyo: Elsevier Tarantola, A. & B. Valette 1982. Inverse problems = quest for information. J. Geophys., 50: 159-170. Wolery, T.J. & S.A. Daveler 1992. EQ6, A computer program for reaction path modeling of aqueous geochemical systems: Theoretical manual, user's guide, and related
documentation (version 7.0). Report UCRL-MA-110662 PT IV. Livermore: Lawrence Livermore National Laboratory
60
Water-Rock Interaction 2001, Cidu (ed.),0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrothermal WaterRocMOrganiclMicrobeInteractions E .L .Shock Group Exploring Organic Process In Geochemistry,Department of Earth & Planetary Sciences, Washington University, St. Louis,MO 63130 USA
ABSTRACT: The discovery of microorganisms in hot springs and deep high-temperature subsurface environments has changed models of hydrothermal systems and the ways they can be studied. Adding organic and genetic analyses of water and rock samples to more traditional types of inorganic analysis greatly expands the scope of studies in active and fossil hydrothermal systems. As hydrothermal processes are recognized as supporting ecosystems, geochemical and biochemical methods will be conjoined into new analytical tools. The results will include new insights into how genes and environment affect each other, new descriptions of ecosystem functions, new quantitative models of the extent and dynamics of the subsurface biosphere, the exploration of a new fossil record, and renewed interest in the emergence and early evolution of life. Great progress has been made in the study of water/rock interactions without much consideration of organic compounds in hydrothermal fluids or of the microorganisms that thrive in hot springs, hydrothermal systems, and high-temperature subsurface environments. Recent revolutions in analytical and theoretical organic geochemistry, microbiology, and molecular biology permit researchers to integrate organic compounds, microorganisms, and their metabolic pathways into studies of waterhock interactions. My purpose here is to illustrate how this expansion of scope will enable many new discoveries, which will in turn profoundly change the pursuits of hydrothermal geochemistry and biogeochemistry, as well as the underlying reasons for why waterhock interactions are studied. Hydrothermal ecosystems host an enormous diversity of high-temperature microorganisms that: catalyze otherwise sluggish oxidation-reduction reactions, * drive mineral dissolution and precipitation, * possess enzymes of enormous practical use, populate the deepest branches of the universal tree of life, and have left a fossil record distinct from that dominated by photosynthesis. Various lines of evidence lead to hypotheses that high-temperature microorganisms populate a large volume of the Earth’s crust, * affect the production of natural gas, the generation of petroleum, the formation of ore deposits, and the rheology of faults,
anchor the biogeochemical cycles that support life at the surface, are the best available analogues for the most ancient life on Earth, and * provide realistic models for imagining life on other planets.
I SOME BACKGROUND Despite early reports of high temperature life (Brewer, 1866; Setchell, 1903; Copeland, 1936; ZoBell, 1958), including the familiar observation of heat generation by decomposing haystacks, compost heaps, coal piles, and other anthropogenic accumulations of organic matter, microorganisms have been thought of as inhabitants of soils, streams, lake sediments, marine muds, pig intestines and other relatively low-temperature environments. This view began to change dramatically in the 1960’s and 70’s following pioneering work at Yellowstone, New Zealand, Iceland and other hot spring systems by Thomas Brock and his colleagues (see review by Brock, 1978). It became evident that microorganisms are thriving at temperatures up to boiling in fluids that are the products of waterkock interactions. With the subsequent discovery of organisms that thrive at temperatures > 100°C and pressures > 1 bar, our view of hydrothermal systems has changed permanently (Deming and Baross 1993; McCollom and Shock, 1997; Delaney et al., 1998; Summit and Baross, 2001). Now they are seen as habitats, windows on a plausible deep subsurface
0
0
0
61
biosphere, and as analogues for early Earth ecosystems and those that may exist on other planets (Shock, 1997; Jakosky and Shock, 1998). Rather than freaks of evolution, high-temperature microorganisms (thermophiles) are now viewed as integral parts of hydrothermal systems. They are seen as arbiters of many water-rock reactions, capable of elemental separations, organic compound transformations and isotopic fractionations, and we now know that they are likely to have pursued these lifestyles for the vast majority of Earth history. Efforts to reveal the relations among all forms of life have lead to the construction of universal phylogenetic trees based on sequences of RNA and DNA (Woese et al., 1990; Pace, 1997; 2001; Woese, 2000). The result is the recognition that thermophiles populate the deepest branches of such trees. Although there is ongoing contentious debate about how such trees are made and what meaning should be imposed on or read from them (Pennisi, 1998, Galtier et al., 1999), the fact remains that evidence from highly conserved genetic material common to all life indicates that the course of evolution has been from high temperature to low. Consistent with this observation is the fact that the use of geochemical sources of energy (chemosynthesis) preceded the use of light energy (photosynthesis) in the production of biomolecules. The apparent ability of many high temperature organisms to pursue their chemosynthetic lifestyles in the dark makes them prime candidates for populating the deeper and hotter regions of the Earth’s crust where ample sources of geochemical energy are generated through geological processes. It follows that many reactions that support life in hot near-surface environments may also support life throughout large volumes of the crust. Because many of these energy sources are oxidation-reduction (redox) reactions, it is plausible that microorganisms catalyze the redox geochemistry that is central to weathering, diagenesis, ore formation, petroleum generation and hydrothermal alteration. The genetic diversity of microorganisms far exceeds that represented by all of the visible forms of life, and it should be no surprise that the metabolic diversity of microorganisms is enormous compared to the familiar models of oxygenic photosynthesis and oxidative respiration that dominate our everyday world. In a survey of the thermodynamics of overall metabolic reactions (Amend and Shock, 2001), we identified 197 redox reactions that have been shown to be mediated by microorganisms, and for which standard Gibbs free energies can be calculated at elevated temperatures. Examples of hightemperature microbially-mediated redox reactions involving inorganic forms of 0, N, S, C, Fe, Cu, U and/or organic compounds are listed in Table I. In addition, there is evidence for lower temperature microbial involvement in redox reactions involving
manganese (Lovley 1993; ZOOO), chromium (Chen and Hao, 1998; Wang, 2000), arsenic and selenium (Stolz and Oremland, 1999; Oremland et al., 2000), chlorine (Coates et al., 1999), mercury (Chen et al., 1997; Hobman et al., 2000), and gold (Southam and Beveridge, 1994), as well as vanadium, molybdenum, silver, and technetium (Lovley, 1993), and there is no obvious reason why these reactions could not be used by thermophiles (Amend and Shock, 200 1). There are other microbially-mediated reactions for which thermodynamic calculations are not yet possible, such as degradation of pesticides (MacRae, 1989; Fournier et al., 1997), chlorinated hydrocarbons (Henson et al., 1989; Hanson et al., 1990; Haas and Shock, 1999), and organophosphorus compounds (Kertesz et al., 1994), and there are undoubtedly many more microbially-mediated reactions that remain undiscovered. Given the extent of the known diversity, it seems likely that any geochemical source of redox disequilibrium can be tapped by some microorganism somewhere. The enzymes that high-temperature microorganisms use to catalyze this wide range of reactions are encoded in relatively small genomes. As an example, the genome of the thermophilic chemolithotrophic methanogen Methanococczrsjannaschii is composed of roughly 1.66 x 106 base pairs, and codes for just less than 1800 proteins (Bult et al., 1996). Genomes of thermophilic organisms are being sequenced in labs around the world. As of January 2001, among the 38 published complete genomes 9 are from thermophiles, and 3 more thermophilic genomes are among the 18 complete genomes awaiting publication (TIGR, 200 1). As these newly aquired data are deciphered and interpreted, new phylogenetic relations and metabolic capabilites will be determined that will once again change the way we think about the diversity of microorganisms and their evolutionary history. After all, a thermophilic genome is the product of evolution in a geochemical environment. Genomic information can be used to design new analytical methods that show great promise for probing the dynamic behavior of biogeochemical processes (see below). In fact, it may soon be as routine to obtain genetic data on water and rock samples as it is to obtain trace element and isotopic analyses today. 2 CONDITIONS FOR A HYDROTHERMAL ECOSYSTEM There are several conditions that must be met for a hydrothermal system to support life. Chief among these are geochemical energy sources, which are only present if the system fails to reach stable thermodynamic equilibrium. Although there is great diversity in the number and variety of redox reaction that can support microorganisms, many involve
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1 able 1. txamples ot the diversity ot high temperature microbially-mediated redox reactions (Amend and Shock, LUU 1 )
Reaction
Species known to catalyze
H2+ 1/202= H20 NO3- + 4H2 + H+ = NH3 + 3 H 2 0 SOL2+ 4H2 + 2H' = H2S + 4H20 S + 3/202 + H 2 0 = SOL2+ 2H' S + H? = H2S NO3- + H2S = NO2- + H 2 0 + S CO2 + 4H2 = CH4 + 2H20 5CH3COOH + 8N03- + 8H' = 4N2 + 1 OC02 + 14H20 2CH3CHOHCOOH + 3S0L2 + 6H' = 6C02 +3H2S + 6H20 3CH3(CH2)*CH20H+ 4So3-*+ 8H' = 4H2S + 6CH3COOH + 3 H 2 0 2FeS2 + 1 5/202 + H 2 0 = 2Fe'3 + 4s04-2+ 2H' ZnS + 2 0 ~= ~ n+ sod-' + ~ 2CuFeS2 + 17/202 + 2H' = 2Cu" + + 4S0L2 + H 2 0 U02 + 1/202 + 2H' = U02+2+ H20 12Fe'3 + CH3CHOHCOOH + 6H20 = 3FeC03 + 9Fe'* + 18H'
Aqu f e x pyrophilus Pyrolobusfumarii Archeoglobus lithotroyhicus Sulfolobus acidocaldarius Pyrodictium occultum Ferroglobus placidus Methanopyrus kandleri Pyrobaculum aerophilum Desulfotomaculum thermobenzoicum Desulfotomaculum thermosapovorans Sulfirococcus yellowstonii Metallosphaera sedula Sulfolobus metallicus Metallosphaera prunae Bacillus infernus
elements that are relatively scarce except in rare circumstances. Rates of redox reactions involving minerals have not yet received the attention paid to pH-driven dissolution and precipitation reactions, but various lines of evidence can be adduced to illustrate the temperature ranges where disequilibrium occurs and the magnitude of the energy supply that can be available. In general, major energy sources in hydrothermal ecosystems correspond to disequilibrium states in the C-S-Fe-N-HzO system, and such states become increasingly common as temperatures fall below about 250°C. Organic compounds are necessarily part of this system, and their contribution to the availability of energy can be readily addressed with thermodynamic data (Shock and Helgeson, 1990; Shock, 1995; Amend and Helgeson, 1997a; 1997b; Dale et al., 1997; Helgeson et al., 1998; Richard and Helgeson, 1998; Plyasunov and Shock, 2000; Amend and Plyasunov, 2001) as illustrated below. Recent advances in analytical methods mean that the demand for compositional data engendered by theoretical studies can now be met. Members of GEOPIG and our colleagues are researching hydrothermal ecosystems at Yellowstone. Our analytical data collected in the field and lab are revealing the diversity of forms that iron, sulfur and organic carbon can take in these systems. With these data we can begin to inventory the availability of energy that can support thermophilic life. As an example, in Obsidian Pool, a hot spring famous for its microbial diversity (Barns et al., 1994; 1996; Hugenholtz et al., 1998a), preliminary results show that magnetite (Fe304) hematite (Fe2O3) and pyrite (FeS2) are present in the sediment, and that pyrite appears to be a product of the reactions occurring in situ. This allows us to propose several reactions leading to pyrite that may be microbially mediated including:
Fe+2+ 7/4 H2S + 114 S04-2 pyrite + H20 + 3/2 H+
(1)
hematite + 1/4 S04-2+ 15/4 H2S + 112 H' 4 2 pyrite + 4 H20
(2)
and magnetite + 11/2 H2S + 1/2 S O i 2 + H+ --+ 3 pyrite + 6 H20
(3) These reactions are all written as fully coupled redox reactions in the Fe-S-HZO system. As such they are likely to reflect overall reactions pursued by consortia of microorganisms rather than single reactions conducted by individual organisms. Nevertheless, it can be seen that changes in pH, sulfate concentration, and flux of H2S in the gas phase will affect the energy availability from these reactions. In our measurements from 1999 and 2000, the temperature of Obsidian Pool has varied from 77.2 to 83.2"C, and pH from 6.48 to 6.78. Concentrations of Fe+2 determined by field spectrophotometry averaged 0.03ppm, which is about 50% of the total Fe concentration we measure by inductively coupled plasma-mass spectrometry (ICP-MS). A field measurement of H2S in the gas phase gave 1600ppm, and lab analyses of sulfate by ion chromatography (IC) vary from 37.3 to 67.4ppm. Using these data, the corresponding logarithmic activity product (log Qr) for reaction (1) varies between -0.65 and -0.72, that for reaction (2) between 9.0 and 9.2, and that for reaction (3) between 15.4 and 15.8. These values of Qr can be combined with standard Gibbs free energies for these reactions, calculated with SUPCKT92 (Johnson et al., 1992) using data and parameters from Shock et al. (1997)
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and Helgeson et al. (1978), to evaluate overall Gibbs free energies and the states of disequilibrium for each of these reactions. As an example, the standard state Gibbs free energy (AG,") for reaction (1) at 77.2"C equals -13060. cal mol-' and can be used with the expression AGr = AGF + 2.303 RT log Qr, together with the appropriate value of log Qr (-0.65), to evaluate the overall Gibbs free energy, which equals -14090. cal mol-'. It follows that at these conditions the maximum amount of energy available per mole of pyrite produced from reaction (1) was -14090. cal. This value is somewhat variable in this location; it equaled -14570. cal at conditions that prevailed at 8 1.4"C. Curiously, the energy yields per mole of pyrite produced from reactions (2) and (3) are -15225. to -14745. cal, and -15430. to -14890. cal, respectively, suggesting that the disequilibrium with respect to pyrite is nearly equal among these three reactions. If the comparison is made per electron transferred (equivalent in these reactions to the energy per mole of H2S consumed), the energy from reaction (1) varies between -8050. and -8325. cal (e-mol)", that from reaction (2) varies from -8420. to -8120. cal (e-mol)-', while that from reaction (3) varies between -7860. and -8 100. cal. (e-mol)-', supporting this conclusion. Additional analytical data permit evaluation of the energy available from primary iron-bearing silicates that may be present in the rhyolite and other rocks undergoing alteration in hot spring systems like Obsidian Pool. As an example, additional data for potassium (IC), and aluminum (ICP-MS) from Obsidian Pool water samples allow evaluation of the amount of energy that would be released by dissolution of annite (KFe3(AlSi3)010(OH)2)and precipitation of pyrite and quartz (Si02) represented by the redox-coupled reaction
methods that must be followed, and which are made more difficult by the near wilderness setting. Nevertheless, we have found that concentrations of dissolved organic carbon (DOC) in several locations can be as high as lOppm and occasionally >50ppm, which exceed values reported by other investigators (Ball et al., 1998). The full range of compounds that are present is a matter of ongoing research, but preliminary results show a strong correlation between concentrations of compounds derived from vascular plants (tannins and lignins) and DOC in the Yellowstone hot springs we have studied (Shock et al., 2001). Of course, vascular plants do not grow in the hot springs, but they are abundant in nearby forests and meadows. This suggests that these compounds survive transport through soils, groundwater and hydrothermal systems, and provide an imprint of the forested ecosystem on the composition of the seemingly disconnected hydrothermal ecosystems. Our recent IC results for the carboxylate anion formate (HCOO-) from many of these same springs, together with field measurements of dissolved oxygen by spectrophotometry, and bicarbonate concentrations obtained from alkalinity titrations indicate that as much as 52000. cal mol-'is available from formate oxidation via HCOO- + 112 02(aq)
-+ HC03-
(5)
even though formate concentrations rarely exceed lppm. This is equivalent to about 26000. cal (emol)-'. Comparison with results for iron mineral alteration summarized above indicates that heterotrophic oxidation reactions may be major sources of energy for hydrothermal ecosystems, even in cases where the total concentration of organic compounds may seem relatively low (see: Amend et al., 1998). It follows that renewed efforts to characterize hydrothermal organic compounds and quantify their concentrations will spawn great improvements i n models of energy flow throughout hydrothermal ecosystems.
annite + 312 H+ + 2114 H2S + 314 sodw2 -+ 3 pyrite + 3 quartz + K+ + AI(OH)4- + 5 H20 (4) Using the same criteria for comparison, there is between -14670. and -14140. cal per mole of pyrite produced from reaction (4), or between -8380. and -8080. cal (e-mol)-', but proportionately less depending on the annite content of the biotite. Once again, we see that the conditions in Obsidian Pool are consistent with the formation of pyrite at the expense of other iron-bearing minerals. Surprisingly, it does not seem to matter whether ferric or ferrous iron is present in the reactant phases. All other iron minerals are unstable relative to pyrite, and all are about equally unstable. Organic analyses from Yellowstone hot springs are relatively rare (Clifton et al., 1990; Ball et al., 1998) despite the long record of geochemical research there. One reason may be the relatively cumbersome sampling protocols and preservation
3 FROM HOT SPRINGS TO THE SUBSURFACE BIOSPHERE Although we have collected samples for only a few years, analytical data summarized here suggest that the composition of Obsidian Pool does not vary greatly. If so, then this hydrothermal ecosystem can be thought of as an open system of interlinked chemostats that are only gently perturbed by natural fluctuations. On the other hand, in subsurface environments it is imaginable that reactants are not renewed nearly as quickly as in a hot spring. Under these circumstances, behavior converging toward a closed system may prevail, and one or another coinpound may become a limiting reactant as the system
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is driven toward equilibrium and the energy dissipated. Mineral alteration reactions (2) - (4) can be used to illustrate the differences in energy availability represented by various sources of iron at conditions where pyrite is the stable product. Using data from Obsidian Pool as an example, and assuming closed system behavior, sulfate would be the limiting reactant in all three cases. Variations in stoichiometry and overall Gibbs free energies allow us to calculate differences in energy per kg of fluid represented by these three alteration reactions. Hematite alteration yields about 46 cal (kg fluid)-’, magnetite alteration about 35 cal (kg fluid)-’ and annite alteration about 20 cal (kg fluid)-’ for the case of sulfate limitation. If, on the other hand, the minerals are to become limiting, then waterhock ratios (W:R) and modal abundances enter the calculations. In order for these minerals to replace sulfate as the limiting factor for energy availability from these reactions, they would have to be present at mol percents considerably < 1 at W:R = 1. It is no doubt true that portions of the subsurface are oligotrophic even though prevailing temperatures are suitable for thermophiles (Whitman et al., 1998). Indeed, regardless of initial supplies of energy, closed systems (or chemically damped systems) will become sterile as microorganisms produce an overabundance of metabolic products on the way toward equilibrium. Long term support of abundant life in a high-temperature subsurface system requires that the system is open to mass transfer. This is the familiar condition required to support ore deposition on any scale that may be economic. It follows that many of the criteria that must be met to account for abundant alteration and mass transfer are the same as those that can support life. Subsurface life is driven by fluid flow, and it is likely that the greatest concentrations of subsurface biomass occur in deep aquifers and fault systems where fluid flow can be on the order of meters per year. In fault zones that act as incubators for thermophiles, seismic events may drive pulses of microorganisms together with fluids. Seismic events accompanied by compositional changes (injection of new fluids, exposure of fresh rock) can trigger subsurface biomass blooms. Incubating regions in faults may be lined with biofilms, which, if abundant, could change the rheology of fault zone materials. Microbial activity, especially methanogenesis, may be capable of driving up gas pressures, which raises the intriguing possibility of a biological mechanism for stress evolution. Such notions of “seismobiology” or “bioseismology” may seem fanciful. But, it should be kept in mind that the drive to make methane during petroleum generation can not be satisfied because equilibrium methane partial pressures greatly exceed the pressures attainable in crustal rocks (Shock, 1988; 1989; 1994; Helgeson et al., 1993). The same results apply if thermophilic
methanogens are responsible for catalyzing the production of natural gas, and even in situations where autotrophic methanogenesis prevails - especially it’ the source of the CO2 consumed is carbonate minerals. 4 GENES, GENOMES AND GEOCHEMISTRY The advent of the polymerase chain reaction (PCR) has enabled researchers to inventory the microorganisms present in an environment by their genetic composition without having to culture the organisms. In effect, the polymerase enzyme makes copies of nucleic acids, and the rate at which copies can be made has been increased dramatically by using polymerases from thermophilic organisms. Researchers often select small subunits of RNA from the ribosome (the cell machinery that produces enzymes) for PCR amplification, and can work with environmental samples that sometimes harbor a hundred or more distinct ribosomal RNA (rRNA) sequences (see reviews by: Ward et al., 1992; Giovannoni and Cary, 1993; Amann et al., 1995; Pace, 1997). This has lead to explosive growth in the number of microorganisms for which at least some genetic information is available (for example there are -15,000 rRNA sequences from Bacteria; Dojka et al., 2000) and has lead to new assessments of microbial diversity in many environments (see: Hugenholtz et al. 1998b; 2001; Dojka et al., 1998; Tanner et al., 1998; Crump and Baross, 2000; Teske et al., 2000; Kaye and Baross, 2000; Massana et al., 2000; among many others). At the same time, we know little about the physiology of the vast majority of the organisms for which we have genetic data. So, it is currently easier to inventory the microorganisms in an environment, and determine how they are related to one another, than it is to discover how they function in biogeochemical processes. Several recent breakthroughs are helping to forge links between phylogenetic and functional information from genetic samples. As an example, it is increasingly possible to quantify the populations of specific types of microorganisms by combining rRNA-specific probes with fluorescent in situ hybridization (DeLong et al., 1999; Christensen et al., 1999), denaturing gradient gel electrophoresis (Muyzer et al., 1993; Murray et al., 1996), terminal restriction fragment length polymorphism (Liu et al., 1997; Clement et al., 1998), and other recent molecular methods (Suzuki et al., 2000; Takai and Horikoshi, 2000; Rondon et al., 2000). It is also increasingly plausible to combine genetic data with microbial activity analyses using microautoradiography (Lee et al., 1999; Ouverney and Fuhrman, 1999), or microelectrodes (Ramsing et al., 1993; 2000; Schramm et al., 1996).
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‘lhe availability of complete genomes for microorganisms is a first step toward a comprehensive analysis of the function of microbial genes, identification of the encoded enzymes, and a greater understanding of how microorganisms interact with their environments. It will soon become common to construct genetic probes that help to identify which types of enzymatic processes actually occur in an environment (Edgcomb et al., 1999; Amann and Ludwig, 2000). These methods will simultaneously reveal the dynamics of biogeochemical processes, and the agents responsible for change. It may in fact be possible to obtain quantitative compositional data on microenvironments by identifying which genes are active, and in which particular combinations. We are embarking on a new era of discovery regarding the actual mechanisms of biogeochemical processes, and the potential for understanding the complex dynamics of water/rock/organic/inicrobe interactions is better than at any previous time.
over the surrounding water (Zn, Mg, Sr) and some are even depleted slightly (Na, Rb). Although much needs to be done to improve the quantitative rigor of this type of work, presently available data suggest that trace element ratios and associations in hydrothermally altered rocks may be indicators of ancient microbial communities. Several metals that have dissimilatory or assimilatory roles in biochemical processes may also exhibit fractionation of stable isotopes in these processes. Recent attention has focussed on iron isotope ratios as possible signatures of biological activity (Beard et al., 1999; Anbar et al., 2000), but other isotopes of U, W, Cr, and MO, may also be affected by microbial metabolism or production of nietalloenzymes. New advances in multicollector ICPMS are being promoted as the means to obtain I > < ) 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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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 <0.05 <0.10 <0.10 0.94<0.10 1.00 max 7.22 9.16 4.90 12.8 1.88 40.5 4.50 24.0 mean 5.49 1.36 0.42 2.09 0.33 4.01 1.42 4.47 vw.mean 5.42 1.19 0.17 1.33 0.25 3.03 0.86 2.89 median 5.33 0.99 1.18 0.21 2.40 0.97 2.94 EF 1.00 3.95 29.33 1.70 1.41 9.62 Some samples displayed values below detection limit for K (1% Ca (I), Mg (4) and NO3 (5).
3.1 Precipitation acidity The collected rainwater samples display pH values from 3.80 to 7.22 (Tab. 1). Such a wide range of values is quite common for rainwaters that are dilute unbuffered solutions. Pure water in equilibriumwith atmospheric CO2 has a pH of 5.65. Small additions of acids or bases, of natural or anthropogenic origin, induce strong deviations from this theoretical value. Ross et al. ( 1989) demonstrated that concentrations of few ppm of SO2 in air lower the pH to values below 4. The main rainwater acidifLing agents are NOx, SO2 (mainly anthropogenic), HC1, HF and organicacids (mainly of natural origin). On the contrary, CaC03, often present in soil dust, is the main proton acceptor. Thus, acidity values measured in rainwater are the result of complex reactions due to the contemporaneous presence of many of these compounds (Dongarra& Francofonte 1995).
Figure 1. Location and schematic lithologic map of the area. l=volcanic rocks; 2=clays and sands; 3=carbonate rocks.
2.3 Climate The climate of the Nicolosi area can be classifed as sub-tropical humid (Pinna 1977). Thirty-year average values are 15 "C for temperature and 1036 mm for rainfall (Regione Sicilia 1998). Pluviometric maxima generally occur in the autumn-winter seasons (meanmonthly values: Oct. 125 mm; Dec. 110 mm). During the period of this study, maximum weekly values were measured at the end of January (139 mm) and at the beginning of December (104 mm). while minimum values (no rainfall) were re84
rangingfrom < 0.1 to 12.8 mgll. The Na+ contents are only slightly lower (12.2%), with values ranging from 0.1 to 9.16 mg/l, while Mg2+ (2.6%) and K+ (1.8%) have contents from < 0.1 to 1.88 mgll and from < 0.05 to 4.9 mgllrespectively. C1- is the dominant anion (3 1.1%) and displays values rangingfrom 0.94 to 40.5 mgll, while SO:- (29.7%) and NO3(8.9%)have contents from 1.O to 24.0 mg/l and from < 0.1 to 4.5 mgllrespectively. The NdCl ratios of the studied rainwater samples are very close to that of seawater (Fig. 3a) indicding a prevailing origin from sea spray. A few samples show a chloride enrichment of probable volcanic origin which will be further discussed in the following section. If we assume that Na’ is entirely of marine origin (which is indeed a rough assumption, if one consider that a not neglgible crustal contribution is likely to exist), it is possible to calculated the enrichment factors of the each X major ion with respect to seawater composition, following the equation:
Measured pH values (Tab. 1) display a median value (5.33) and a volume-weightedmean (5.42) that are both lower then 5.65 indicatinga prevalent input of acid compounds. Subdividingthe samples in two classes on the basis of their pH value (>5.65>), we can observe that the more acidic samples generally display low cation contents, while being enriched in nitrate (Tab. 2). Chloride and sulphate contents display no significantcorrelation with pH values. This means that acid compounds deriving from NOX are probably the more effective in lowering rainwater pH in the Nicolosi area. Strong acid compounds, such as HC1, HF and SO2, released by volcanic a0 tivity from the central craters, are probably partially buffered through reaction with the contemporaneously releasedvolcanic ash. Table 2. Median values of rainwater samples subdivided in two classes on the basis of the measured pH values. Na Ca Mg C1 NCX SO4 pH<5.65 0.54 0.85 0.21 2.41 1.68 3.05 pH>5.65 1.29 1.89 0.32 2.40 0.90 3.02
where X and Na are the concentrations (in mdl) in the rainwater samples (r) and in seawater (sw), respectively. As evidenced by Table 1, where the computed average enrichment factors are reported, all ions display EF values higher then 1, Ca2’ and S042-being the most enriched. This implies that some additional source of chemicalsin rainwater has to be taken into account. The scatter diagramsof Figure 2 are great help in revealingthe origins of major ions in rainwater. As a whole, the diagramssuggesta significant contribution from the intake of solid dust and other fine-grained crust materials. For example, we note that the MgNa ratios in raiwaters (Fig. 2a) are comprised between the marker ratios of seawater (0.12) and of the Etnean rocks (1.13). K/Na and of Ca/Na ratios (Figs. 2b, c) indicate that rainwaters are enriched in K and Ca with res ct to both seawater and Etnean rocks. Excessof Kpecould be related to the use of synthetic fertilisers while that of Ca2’ to soil dust coming from the carbonate rocks of the Hyblean plateaux. The highest Ca2’ content was measured in a small volume sample collected in August after 5 weeks without precipitations. Its Ca/S04 ratio, very close to that of gypsum, indicates a likely contamination from Saharan dust (Caboi et al. 1992), which is consistent with the south-western winds (“scirocco~~) blowing by the time of sampling. The collected samples generally display S0,JCl ratios higher then the typical values of both seawater (7.2) and plume particles and gases (1.2-30), indicatingthat most of the sample have an additional SO4 source. Most of the samples display a S04/N03 ratio of about 2 (Fig. 3b) not far from that reported in EEA (1995) for average emission rates of gaseous pollutants in Italy (-l),
Figure 2. Scatter diagrams of some major cations in the sampled rainwaters.ER and SW are the characteristiccation ratios of the Etnean rocks (average)and of seawater, respectively.
3.2 Major ion content Collected rainwater samples display large variations in ion content. Ca2’ is the dominant cation (13.7% on volume weighted mean basis) with concentrations 85
produced the 1991-93 eruption, while the higher sulphate contents of the period November-Decembel are related to the shallow degassing of the same magma.
pointing to a main anthropogenic source for both compounds. The rare samples characterisedby very high S04/N03ratios (16-240) give their sulphate excess to gypsum in soil dust (1 sample) or to SO2 degassingat the summit craters.
4 CONCLUDING REMARKS The present study evidences that major ions in rainwater fiom the Etnean area may originate fiom different sources, such as sea spray (Na and Cl), volcanic plume and ash (Cl, SO4 and, secondarily, cations), soil dust both of local (NayK and Mg) and more distal origin (Ca) and atmospheric pollution from human activities (NO3 and SO4). The anthropogenic source appears to be the main cause of the low pH values of rainwaters. Finally, this study highlightesthe significantcontribution of volcanic activity to the chemistry of rainwaters, especially during paroxysmal periods. REFERENCES
Figure 3. Variation with time of the Cl/Na and SOdNO, ratios in the sampled rainwaters.
3.3 Variationsrelated to volcanic activity
Fig. 3 shows the evolution of the CVNa (a) and Sono3 (b) ratios during the sampling period. Clearly, both ratios are characterised by a relevant temporal variability, which fiuther evidences the existence of several sources for these ions. Excepting for the sample collected in August, which is contaminated by Saharan dust gypsum, all rainwater samples displaying high Sono3values were collected in periods of high SO2 fluxes andor intense strombolian activity. Also, high CVNa values in rainwaters are recorded during phases of enhanced strombolian activity at the summit craters. Notably, peaks in the CVNa ratio are shifted, and generally precede periods of high Sono3values. This finding can be explamed by the different behaviour of C1 and S during magmadegassing.Pennisi & Le Cloarec (1998), in fact, correlate high Cl/S ratios in the volcanicplume to deep magma degassing (P> 100 W a ) , while low values are related to shallow “eruptive” degassing (P<20 MPa). Accordingly, we may sug gest that the temporal variability of rainwater chemistry can be ascribed to the mutable composition of the Etnean plume during the evolution of the volcanic system. Thus, the h g h chloride contents in rainwater of the period September-October are probably related to deep degassingof the magmathat 86
Aiuppa, A., Allard, P., D’Alessandro, W., Michel, A., Parello, F., Treuil, M. & M. Valenza, 2000. Mobility and fluxes of major, minor and trace metals during basalt weathering and groundwater transport at Mt. Etna volcano (Sicily). Geochim Cosmochim. Acta 6411 1: 1827-1841. Allard, P., Carbonelle, J., Dajlevic, D., Le Bronec, J., Morel, P., Robe, M.C., Maurenas, J.M., Faivre-Pierret, R., Martin, D., Sabroux, J.C. & P. Zettwoog, 1991. Eruptive and diffuse emissions of CO;?fiom Mount Etna. Nature 351: 387-39 1. Bruno, N., Caltabiano, T. & R. Romano, 1999. SOz emissions at Mt. Etna with particular reference to the period 1993-1995. Bull. Volcanol. 60: 405-41 1. Calvari, S., Coltelli, M., Neri, M., Pompilio, M. & V. Scriban0 1994. The 1991-1993 Etna eruption: chronology and lava flow-field evolution. Acta Vulcanol. 4: 1-14. Caboi, R., Cidu, R., Cristini, A., Fanfani, L. & P. Zuddas 1992. Influence of saharan dust on the chemical composition of rain in Sardinia, Italy. In: Proc. of the f h Internat; Symp. On Water-Rock interaction - BXI 7- Park City, Utah, USA, 13-18 JuIy 1992. Y.K. Kharaka and A S . Maest Eds., Balkema Rotterdam, 469-472. Dongarra, G. & S. Francofonte, 1995. Quality of rainwater: a geochemical process of water-air-rock-life interaction. Envir. Geol. 25: 149-155. EEA 1995. Enviroment in the UE, The Statistical Compendium. Pennisi, M. & M.F. Le Cloarec, 1998. Variations of C1, F and S in Mount Etna’s plume, Italy, between 1992 and 1995. J. Geophys Res. 103(B3): 5061-5066. Pinna, M. 1977. Climatologia. UTET, Torino. Regione Siciliana 1998. Climatologia della Sicilia. Palerrno, Tipogrda Priulla . Ross, M., McGee, E.S. & D.R. Ross, 1989. Chemical and mineralogical effects of acid deposition on Shelburn Marble and Salem Limestone test samples placed at four NAF’AP weather monitoring sites. Am. Mineral. 74: 367-383.
Water-Rock lnteraction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrothermal systems as indicators of paleoclimate: an example from the Great Basin, Western North America G .B .Arehart & S .R .Podson University of Nevada, Reno; Reno, Nevada, USA
ABSTRACT: Many hydrothermal systems are dominated by water of meteoric origin, the stable isotopic composition of which is, in part, a function of climate extant at the time the hydrothermal system was active. Therefore, measurement of stable isotopic parameters in fossil hydrothermal systems provide a signal averaged over relatively large time scales that potentially can yield insights into longer-term climate changes. Two excellent indicators of the H isotope composition of meteoric waters from hydrothermal systems include fluid inclusions and hydrous minerals. In the Great Basin of western North America, hydrothermal systems have been active intermittently since at least Jurassic time. Data from the last half of the Tertiary in the Great Basin correlate well with existing data from weathering-related minerals of known age (e.g. alunite, kaolinite) and parallel global paleoclimate curves. A few data exist for older hydrothermal systems and can be utilized to constrain meteoric water back to the Jurassic, but with relatively large uncertainties. Additional data are required to better define this older portion of the secular curve for the Great Basin.
1 ORIGIN AND ISOTOPIC COMPOSITION OF METEORIC FLUIDS
is clearly not strictly a function of latitude, as originally postulated by Arehart and O'Neil (1993). Rather, there is a broad band of lowest values that runs in a NE direction from Reno (Fig. 1a). However, once this regional trend is taken into consideration, the stable isotopic composition of meteoric fluids is expected to reflect primarily paleoclimatological variations over the past -50 Ma in the Great Basin, rather than reflecting changes in local altitudinal or latitudinal conditions.
The stable isotopic composition of meteoric fluids has been shown to reflect a number of parameters, including differences in paleolatitude, differences in paleoaltitude, or changes in paleoclimate (e.g. Taylor 1979, Arehart & O'Neil 1993). Changes in latitude required to affect the 6D values are much larger than those known for western North America over the time span (-50 Ma) represented by most of the samples described below (Denham & Scotese 1988). Stable isotope compositions of precipitation can vary significantly with altitude, but in most cases groundwaters are long-term averages such that local variations due to altitude, as well as seasonal variations, are minimized or eliminated (e.g. Thordsen et al. 1992). In the Great Basin, Arehart and O'Neil (1993) showed that if local altitude effects were dominant, the data for their sample set would be expected to be more random. This is supported by the fact that deuterium analyses of present-day springs across Nevada are broadly similar for any given latitude, regardless of the altitude of the spring (OWeil & Silberman 1974, Mariner et al. 1983, G. Arehart, J. O'Neil, S. Poulson, unpub. data). However, it is shown below that the variation in spring composition
2 HYDROTHERMAL SYSTEMS AND METEORIC WATER Many hydrothermal systems that develop in the upper crust of the earth are dominated by meteoric fluids. Meteoric fluid circulation is driven by thermal anomalies in the upper crust, often associated with elevated geothermal gradients from shallow intrusive activity. The oxygen isotope composition of hydrothermal fluids may reflect the oxygen isotope composition of the parent fluid (meteoric, metamorphic, magmatic) with modifications from water-rock interaction at depth and elevated temperatures. Although some geothermal fluids are essentially unexchanged (e.g. 87
Figure l a (left). Present-day contour map of the northern Great Basin for 6D values of warm and hot springs. Data modified and updated from Mariner et al. 1983. White dots are locations for which alunite paleoclimate data are available. 1b (right). Paleo-isotopic contour map of the northern Great Basin for 6D values at ca. I6 Ma. Data from O'Neil & Silberman 1974, Arehart & O'Neil 1993, and this study. Black dots are locations for which fluid inclusion paleoclimate data are available. R = Reno Wairakei, NZ;Truesdell 1984) and have 0 isotope compositions identical to local meteoric water, other geothermal fluids (e.g. Iceland; Truesdell 1984) show isotopic evidence of extensive exchange of oxygen between local meteoric waters and the rocks through which they have circulated. Therefore, measurement of oxygen isotopes in hydrothermal fluids is not useful as an indicator of the isotopic composition of the local meteoric fluid at the time of hydrothermal activity unless it is clear how much exchange has taken place with the rocks during hydrothermal circulation. In contrast to exchange of oxygen between water and rock and consequent modification of the isotopic Compositionof a geothermal fluid, minimal effects are expected on the hydrogen isotopic composition of most geothermal fluids. Because rock contains very small amounts of hydrogen, fluids passing through most rocks will not be modified significantly in their hydrogen isotope composition unless they have passed through large volumes of rock and interacted extensively. This effect is well-known in the literature (e.g. Taylor 1979). Therefore, the hydrogen isotope composition of meteoric-hydrothermal systems should be reflective of the isotopic composition of local meteoric water at the time the geothermal system was active. Magmatic, metamorphic, and seawater fluids also may contribute water to hydrothermal systems. Such fluids are likely to have stable isotopic signatures significantly different from meteoric water-dominated systems. However, in most cases, these fluids can be recognized by other chemical tests and can be eliminated from any data set that relates to meteoric
water. In particular, seawater and magmatic fluids usually have much higher salinities (>5 wt. % NaCl equivalent) than meteoric-hydrothermal (<5 wt. % NaCl equivalent) systems. Metamorphic fluids often can be recognized by the geologic environment (deeper crustal) and the presence of significant (>-3%) CO,. None of the samples included in this study have demonstrable components of magmatic, metamorphic, or seawater.
3 DETERMINATION OF METEORIC FLUID COMPOSITIONS I"FOSSIL SYSTEMS In most cases, hydrothermal activity is evanescent, but may leave behind some mineralogical record of its former presence. These mineralogical proxies commonly comprise vein material containing quartz, phyllosilicates (sericite, clays), pyrite andor other minerals. Presuming that these minerals formed in equilibrium with the hydrothermal fluids, the isotopic composition of the former hydrothermal fluid can be calculated from equilibrium fractionation equations (Kyser 1986, Gilg & Sheppard 1996 and others) provided the temperature of deposition (equilibration) is known or can be estimated. Oxygen-bearing minerals can, therefore, provide an accurate measure of the oxygen isotope composition of the hydrothermal fluid. However, because of the problem of water-rock exchange described above, oxygen generally cannot be utilized as a measure of the meteoric fluid composition. In contrast, hydrogen-bearing minerals are much more likely to provide a reasonable estimate 88
meteoric water.) Generally, there are both altitude and latitude effects that appear to influence the meteoric water isotopic patterns, resulting in a NE-trending zone of lowest 6D values with higher values in N-S and E-W belts adjacent to the zone of lowest bD. Fig. Ib is a more areally restricted contour map drawn for the period 13-19 Ma, based on alunite (Arehart & O'Neil 1993) and fluid inclusion (O'Neil and Silberman 1974, Arehart & O'Neil 1993, this paper) data from several localities. The number of data for this time period are significantly less than for the present day, and there a few data outliers (as there are with present-day data). However, in general, the paleo-isotopic data result in patterns of isotopic contours very similar to those of the present day. Our interpretation of this pattern is that the altitudinal and latitudinal effects on meteoric water isotopic composition were very similar at ca. 16 Ma to those extant today.
of the hydrogen isotope composition of the parent meteoric fluid because of the larger fraction of H in the water relative to H in the rock. Therefore, this paper focuses exclusively on the H isotope signature of fossil geothermal systems. In addition to mineral proxies, direct measurements of the isotopic composition of the (now vanished) hydrothermal fluid can be made on fluid inclusions in hydrothermal minerals. Both 0 and H measurements can be made, but as with the mineral proxies, 0 isotope measurements may not reflect the parent meteoric fluid composition. In addition, there is the possibility of post-entrapment exchange of 0 with any mineral (e.g. quartz) that contains 0 in its structure. This post-entrapment exchange should not affect H isotopes in anhydrous minerals. There also are problems with fluid inclusion measurements that must be addressed before the data are accepted as representative of parent meteoric fluids. Fluid inclusions must be determined to be of the same age as the host mineral (or at least very closely related in time). Secondary inclusions could form in the vein minerals at some later date and not necessarily be representative of meteoric fluids at the time the system was active. Often this problem can be resolved by microscopic inspection and determination of the approximate temperature of deposition.
4 GREAT BASIN PALEOCLIMATE DATA Figure 2. Secular 6D curve for meteoric fluids normalized to 1 16"W longitude, 4 1 ON latitude. White diamonds are data which do not appear to fit this curve (see text). Data from Table 1.
Given the caveats above, we have assembled the extant data from hydrothermal systems in the Great Basin since -75 Ma (Table 1). Insights into the variability in the 6D values of meteoric water in both time and space can be obtained from these data by: 1. Comparison of paleo-isotope data to present-day spring isotope data; and 2. Construction of a secular curve for meteoric water using combined age and paleo-isotope data. Fig. l a is a contour map of 6D values for presentday geothermal springs and wells, based on the data of Mariner et al. (I 983) modified to include several new data. Because of the mixing effect of subsurface flow, this map is generally representative of the 6D values of average meteoric waters. (In the present dry climate, the majority of subsurface waters are more truly representative of winter precipitation because most summer precipitation evaporates before getting into the groundwater system. However, the relative amount of summer precipitation is small relative to the amount of winter precipitation, therefore, the groundwater is very nearly representative of total
From these paleo-isotopic data, then, we can construct a temporal pattern of isotopic variation for the Great Basin (Fig. 2). A similar figure (Arehart & O'Neil 1993) was originally based primarily on alunite data for which a latitudinal correction only was applied (i.e. all data were normalized to 41 O N with a 5%0/degree latitude correction). Fig. 2 is a modification of the secular curve of Arehart and O'Neil for paleoisotopic composition of water in the Great Basin between 0 and 75 Ma, based on data from fluid inclusions and hydrous silicates (kaolinite, sericite) from hydrothermal systems combined with the original alunite data set. Although the curve has been extended to 75 Ma, the constraints on the data beyond ca. 40 Ma are considerably less than for the younger data. These data (Table 1) have been normalized to a location of 41 ON and 116"E, based on the contour maps of Fig. 1. We believe that such a normalization
89
Table 1. Paleo-isotopic data for alunite (A), fluid inclusions (F) and kaolinite (K) samples.
5 SUMMARY AND CONCLUSIONS
Location Roberts Mtns Sleeper Bodie Gilbert Alligator R Post Round Mtn Post Round Mtn Aurora Preble Goldfield Coinst ock Jarbidge Rabbit Creek Trade Dollar Eastern Star Midas Rawhide Tenm i le Manhattan B uc kho ri i Rain Tonopah Rain Rain Wonder Preble Gold Quarry Gold Quarry Gold Quarry Lone Mtn Carlin Trend Getchell Tuscarora Humboldt
Stable isotope measurements from hydrothermal systems that were dominated by meteoric waters (primarily epithermal systems, but other systems as well) in the Great Basin of North America provide important insights into both spatial and temporal variations in the composition of paleo-meteoric water. From these water compositions, inferences about paleoclimates can be drawn. Isotopic composition of meteoric waters across the Great Basin at ca. 16 Ma appear to have been broadly similar to present-day patterns. Temporal variability in the isotopic composition of meteoric waters in the Great Basin yields a secular curve (Fig. 2) that probably reflects changes in paleoclimate over the time period from late Cretaceous to present. Similar curves could be drawn for other areas of the world that have experienced significant hydrothermal activity. Extension of such secular curves to the more distant past is possible, but with decreasing accuracy.
Type
b
A A F F A A A A
5.3 5.4 8.0 8.0 8.3 8.6 9.5 9.5 9.8 10.0 11.3 11.6 13.0 14.0 14.4 15.0 15.3 15.3 16.0 16.0 16.0 16.6 18.8 19.0 20.0 20.7 22.0 23.0 25.9 27.9 28.9 38.0 39.0 39.0 39.0 73 .O
A
F A A F F A F F F F F F A
A F A A F A A A A K K K K F
bD corr bD -1 12 -120 -128 -140 -98 -108 -111 -129 -129 -140 -137 -137 -94 -1 10 -126 -126 -124 -107 -124 -133 -1 13 -104 -1 15 -90 -133 -138 -139 -139 -106 -1 15 -136 -141 -121 -121 -139 -139 -120 -128 -97 -1 12 -1 16 -130 -102 -1 15 -1 15 -130 -1 12 -90 -141 -141 -125 -125 -139 -139 -131 -122 -132 -132 -149 -149 -144 -144 -140 -140 -145 -140 -153 -148 -136 -136 -109 -95
REFERENCES Arehart, G.B. & J.R. O'Neil. 1993. D/H ratios of supergene alunite as an indicator of paleoclimate. Journal of Geophysical Research Mon 8: 277-284. Denham, C.R. & C.R. Scotese 1988. TerraMobilis - a plate tectonics program for the Macintosh: version 2.1, Earth in Motion Technologies, Houston, TX. Gilg, H.A. & S.M.F. Sheppard 1996. Hydrogen isotope fractionation between kaolinite and water revisited. Geochimica et Cosmochimica Acta 60: 529-533. Kyser, T.K. 1986. Equilibrium fractionation factors for stable isotopes. Mineralogical Association of Canada Short Course 13, 1-84. Mariner, R.H., Presser, T.S. & W.C. Evans 1983. Geochemistry of active geothermal systems in northern Basin and Range province. Geothermal Resources Council Special Report #13,95-119. ONeil, 1.R. & M.L. Silberman 1974. Stable isotope relations in epithermal Au-Ag Deposits. Economic Geology, 69: 902-909. Taylor, HP 1979. Oxygen and hydrogen isotope relationships in hydrothermal mineral deposits. In Barnes, H.L. Geochemistry of Hydrothermal Ore Deposits. New York, Wiley, 236-277. Truesdell, A.H. 1984. Stable isotopes in hydrothermal systems. Reviews in Economic Geology. 1 : 129-142. Thordsen, J.J., Kharaka, Y.K, Mariner, R.H, & L. White 1992. Controls on the distribution of stable isotopes of meteoric water and snow in the greater Yellowstone Proceedings National Park region, USA. International Symposium on Water-RockInteraction 7: 591-595.
is more accurate than applying a simple latitudinal correction. It is clear from Fig. 1 that patterns of isotopic distribution in meteoric waters were similar in the Great Basin at least during the late Tertiary. Extension of these patterns to earlier Tertiary times is probably more valid than the simple latitudinal correction. Both weathering-related alunite, which can be dated directly (Arehart & O'Neil 1993) and fluid inclusion data (O'Neil & Silberman 1974, this study) yield similar patterns through time. Generally, 6D values of meteoric fluids are fairly constant from 0-12 Ma, rise to a maximum at ca. 13- 17 Ma, decrease between 1740 Ma, and slowly rise between 40-75 Ma. Four fluid inclusion data at ca. 15 Ma fall off of this general trend and may represent either secondary inclusions (age younger than 15 Ma); an incorrect age for the inclusions; or possibly local climate variability. Additional examination of these anomalous data is warranted. 90
Water-Rock lnferaction 2001, Cidu (ed.), 02001 Swefs & Zeitlinger, Lisse, ISBN 90 2651 824 2
Temporal variations of 3He/4He ratios of dissolved helium in groundwaters of Mt Etna, Southern Italy W.D ' Alessandro Istituto di Geofsica e Vulcanologia - Sezione di Palermo, Italy
F.Parel1 &, B .Parisi Dipartimento di Chimica e Fisica della Terra, Palermo, Italy
P.Allar & P.Jean-Baptiste Laboratoire des Sciences du Climat et de I'Environnement CNRS-CEA, Saclay, France
ABSTRACT: He isotopes were measured in groundwaters of Mt Etna in order to provide fbrther insights on the relationship between volcanic activity and the geochemistry of dissolved gases. The intense seismicity and the strombolian activity that characterized Mt Etna since the end of 1994 was accompanied b a variation in the fluid regime, leading to the input of magmatic He and C02. He and CO2 contents and 3He/ He isotopic ratios showed interesting variations in some of the selected sampling sites, during the last months of 1994 and in the middle of both 1995 and 1996, in concordance with the renewal of summit crater activity and the intensity of both volcanic tremor and earthquakes.
J
1 INTRODUCTION
2 BACKGROUND
In volcanic areas, helium isotopes are widely utilized as powerfbl markers of the origin of fluids, as the upper (8 f 1 R,) and lower mantle (35 R,) and the continental crust (< 0.1 RJ show distinct R/R, ratios (R, =1 .38.10-6). Typically, magmatic gases are relatively 3He-enriched with respect to crustal or atmospheric gases containing more 4He. Parallel to the survey of fkmarolic gases, it has been recently considered important to monitor helium dissolved in groundwaters in order to better understand the spatial distribution of magmatic upwelling, even in areas where gas manifestations (i.e. fbmaroles or mofettes) do not appear (Allard et al. 1997, Federico 1999). The periodic monitoring of He isotope composition appears to be an interesting tool to evaluate the variation of the volcanic activity, which in turn modifies the proportion of the three components (magmatic, crustal and atmospheric) and, ultimately, the R/R, ratio. This study focuses on the temporal variations of 3He/4He ratios in groundwaters collected on Mt Etna. This latter is a very active volcano, which frequently presents an intense explosive activity at summit craters, sometimes accompanied by lava flow; a continuous magmatic degassing occurs at summit craters. On Mt Etna, He isotope composition was determined in some selected sites from the main degassing areas, in the south - western and eastern sectors, respectively. The period of observation spans from April 1994 to September 1996.
Mount Etna is a highly active alkali-basaltic stratovolcano, which has built upon tensional faults located in eastern Sicily, at the collision boundary between the African and European plates (Barberi et al. 1974). Its activity has been controlled by the intersection of two main fault systems, striking respectively NNW-SSE and "E-SSW, and a shallower E-W one, cutting a 18-20 km thick continental crust whose upper part is made of mesozoic-pleistocenic deposits (Lo Giudice et al. 1982). An intense soil degassing of magma-derived CO2 and He occurs along the flanks of the volcano mostly coincident with zones of weakness (Allard et al. 1991; D'Alessandro et al. 1992; Giammanco et al. 1998). In these fractured and seismically active zones (i.e., Paterno - Belpasso to the SW and Zafferana S.Venerina to the E), groundwaters are enriched in dissolved salts, pointing to a greater water-basalt interaction, in response to CO;? inflow into the aquifer ( h a et al. 1989, Allard et al. 1997, Brusca et al. in press). C02-rich groundwaters contain a mantle-derived magmatic component, with a 3He/4He ratio of 6.9 f 0.2 R, (Allard et al. 1997). Apart from the summit zone, this magmatic component is preferentially concentrated in Col-rich groundwaters that issue from two remote sectors of the southsouthwest and eastern volcano flanks. The 3He/4He ratio of the magmatic end-member coincides with that of helium trapped in the He-rich olivine crystals of Etna basalts (average = 6.7 f 0.4 R,) pointing to its negligible dilution by radiogenic He from the 91
crustal basement. These values, lower than the typical MORI3 value of 8 Ra,indicate a partly contaminated upper mantle zone that is upwelling beneath this region (Marty et al. 1999).
4 RESULTS AND DISCUSSION Results of isotope analyses of dissolved He are listed in Table 1, together with temperature and pC02 values. 3He/4Hevalues range from 1.1 < R/% < 6.7 in the SW sector and from 0.9 < R / R a < 3.8 in the E one. Helium concentrations show a wide variability, ranging from 4-10-*to 1.34.10-~ cc/g STP. In Figure 2, temporal variations of both I U R a ratios and 3He concentrations are shown, together with pCOz values. After the 1991-93 eruption, Etna remained in a quiescent state until June 1994, when an increasing explosive activity marked the beginning of the now-active eruptive period (Coltelli et al. 1998). This is characterized by ash emissions and explosions at summit craters and, occasionally, high altitude lava flows, gradually stopping within some hundreds meters lengths.
3 SAMPLING AND METHODS Figure 1 shows the location of sampling points. Temperature and pH of groundwaters were determined upon sampling. Theoretical pC02 has been computed from measurements of alkalinity and pH at sampling temperature. For He isotope analyses, waters were sampled in special copper tubes from which the dissolved helium was subsequently extracted under vacuum, following a routine procedure (Jean-Baptiste et al. 1992). Dissolved He contents are expressed in cc/g STP, with an overall uncertainty of f 1%. He isotope measurements were made at LSCE with a VG 3000 mass spectrometer, connected to a high-vacuum inlet line. The spectrometer is equipped with a double collector for the simultaneous measurements of 3He' and 4He' ions. The 'He' beam is collected by a Faraday cup and integrated by a digital counter, while the He' beam is detected by an electron multiplier. 3He/4He ratios (R) were determined against an air standard and are referred to the atmospheric ratio (1 .38.10-6) as R&. The overall uncertainty on IURa is f 3%.
Figure 1. Location of sampling hydrogeological basins are outlined.
points.
Table 1. Partial pressure of CO2, He contents and isotope composition of Etna groundwaters. He concentration is expressed as cc per gram of water. Data for temperature and pCOz are mean values. R/Ra He (10-'cc/g STP) pC02 T Sample range mean range mean atm "C 1 1.1-1.8 1.6 5-6 6 0.2 4.4 4.4 55-63 2 4.4-4.5 71 0.2 5.4 5.9 10-60 3 4.6-6.7 30 0.9 7.6 4 1.2-6.4 3.6 4-70 17 0.7 9.9 5.5 5-17 5 4.5-6.0 8 1.3 9.2 6 1.3-1.6 1.5 5-20 7 0.1 10.0 2.6 7 1.2-3.2 5-134 15 0.3 16.7 1.4 5-21 8 1.1-1.6 7 0.2 12.9 5-24 0.5 21.1 9 1.9-3.8 3.2 8 3.0 8-14 10 10 2.5-3.2 0.1 17.0 11 0.9-1.0 1.0 4-7 5 0.1 13.8 Licodia: 1; Romito: 2; Difesa: 3; A. Rossa: 4; A. Grassa: 5; Ilice: 6; Guardia: 7; S. Giacomo: 8; Petrulli: 9; S. Paolo: 10; Vena: 11.
This eruptive period also comprised some of the most important seismic sequences registered during the entire time interval spanned, with very strong episodes in November-December 1994 and FebruaryMarch 1995 (Spampinato et al. 1998). All these events released relatively high energy. During July 1995, the Bocca Nuova crater has been interested by episodes of intense strombolian activity with the emission of blocks and bombs (La Volpe et al. 1999). Moreover, a striking increment of volcanic tremor amplitude in the last months of 1995 has been recorded, linked to the renewal of eruptive activity at NE crater (fire-fountain episodes). Furthermore, an intense seismic activity spread throughout the volcanic area. The last important seismic crisis occurred in the period AugustNovember 1996, and it has been recognized to be mainly related to the regional tectonic structures (Spampinato et al. 1998).
Limits of
92
Figure 2. Temporal variations of FURavalues, 3He contents and PCOZin some selected groundwaters. The secondary Y axis on the right side refers to 3He contents, expressed as cc/g STP.PCOZis expressed in atm. The daily occurrence of seismic events is also reported.
In correspondence to the first intense phase of seismic activity and increment of volcanic tremor (1995), some waters (e.g. Petrulli and S. Paolo wells, secondly Romito spring) show a strong increase of He concentration, while the R/R, ratios remain almost constant. An input of magma-derived He could be identified in the aforementioned groundwaters during this stage of activity, accompanied by an increase of p C 0 ~in only S.Paolo well. In Petrulli and Romito wells, on the contrary, a pC02 decrease corresponds to the 3He increase. This finding is probably to be attributed to the different water solubility of He and CO2 and, consequently, to their chemical fractionation during gas ascent and interaction with groundwaters. This produces a delay in the response time of dissolved CO2 with respect to He. As expected, the variations of measured parameters, that are response to the evolution of volcanic activity, are more easily appreciated where the background values are rather low. In fact, especially in Romito spring (Fig. 2a), Difesa well (not
The variation in the stress regime of the system is to be related to changes in both fluid composition and diffusive processes. Stress-induced variations of gas pressure - i.e. pC02 - also influence groundwater composition as dissolved CO2 produces waters strongly aggressive with respect to the host rocks. As supported by chemical and isotopic data, this process results in a very intense rock weathering (D'Alessandro et al. 1997; Aiuppa et al. 2000). On Mt Etna, a general increase in dissolved He contents has been observed during the rise of new magma batches or stress field variations, sometimes paralleled by increases in the R/R, ratios, explained as magmatic "He input. Temporal variations in the RR, ratios can be also related to isotope fractionation between 3He and 4He during magma ascent and degassing as suggested by Nuccio & Valenza (1998). As a matter of fact, in the early separation of gas bubbles during magma uprising, 'He preferentially partitions into the gas phase, thus leading to higher R/R,ratios.
93
Mt. Etna volcano (Sicily). Geochim. Cosmochim. Acta 64: 1827-1841. Allard, P., Carbonelle, J., Dajlevic, D., Le Bronec, J., Morel, P., Robe, M. C., Maurenas, J. M., Faivre-Pierret, R., Martin, D., Sabroux, J.C. & P. Zettwoog 1991. Eruptive and diffuse emissions of COz from Mount Etna. Nature 351: 387-391. 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. Anza, S., Dongarra, G., Giammanco, S., Gottini, V., Hauser, S. & M. Valenza 1989. Geochmica dei fluidi dell'Etna: Le acque sotterranee. Mineral. Pelrogr. Acta 32: 23 1-251. Barberi, F., Civetta, L., Gasparini, P., Innocenti, F., Scandone, R. & L. Villari 1974. Evolution of a section of the AfricaEurope plate boundary: paleomagnetic and volcanological evidence from Sicily. Earth Planet. Sci. Lett. 22: 123-132. Brusca, L., Aiuppa, A., D'Alessandro, W., Parello, F., Allard, P. & A. Michel 2001. Geochemical mapping of magmatic gas-water-rock interactions in the aquifer of Mount Etna volcano. J. Volcanol. Geothernz. Res. (in press). Coltelli, M., Pompilio M., Del Carlo, P., Calvari, S., Pannucci, S. & V. Scribano 1998. Etna - Eruptive activity. Acta Vulcanol. 10: 141-148. D'Alessandro, W., De Domenico, R., Parello, F. & M. Valenza 1992. Soil degassing in tectonically active areas of Mt. Etna. Acta Vulcanol. 2: 175-183. D' Alessandro, W., De Gregorio, S., Dongarra, G., Gurrieri, S., Parello, F. & B. Parisi 1997. Chemical and isotopic characterization of the gases of Mount Etna (Italy). J. Volcanol. Geotherm. Res, 78: 65-76. Federico, C. 1999. Interaction between magmatic gases and the hydrological system at Vesuvius (Southern Italy): evidences from water and gas geochemistry. PhD thesis, University of Palermo. Giammanco, S., Gurrieri, S. & M. Valenza 1998. Anomalous soil CO2 degassing in relation to faults and eruptive fissures on Mount Etna (Sicily, Italy), Bull. Volcanol. 60 (4): 252-259. Graham, D., Giacobbe, A., Spera, F. & G. Tilton 1992. Chemical and isotopic variations in historical lavas from Mount Etna, EOS Tram. Am. Geophys. Union 73: 61 1. Jean-Baptiste. P., Mantisi, F., Dapoigny, A. & M. Stievenard. 1992. Design and performance of a mass-spectrometric for measuring helium isotopes in natural waters and for a lowlevel tritiuin detennination by the 3He ingrouth method. Appl. Radiat. Isotop. 43: 881-891. La Volpe, L., Manetti, P., Trigila, R. & L. Villari 1999. Volcanology and chemistry of the earth's interior. Italian research activity (1995-1998) report to IAVCEI. Boll. GeoJ Teor. Appl., 40: 163-298. Lo Giudice, E., Patane, G.. Rasa, R. & R. Romano 1982. The structural framework of Mount Etna, Menz. Soc. Geol. It. 23: 125-158. Marty, B., Trull, T., Lussiez, 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: 23-29. Nuccio, P.M. & M. Valenza 1998. Magma degassing and geochemical detection of its ascent. In Arehart and Hulston editors, Proc. intern. Symp on WRI-9, Taupo, New Zealand, 3OMarch-3 ilpril1998: 475-478. Rotterdam: Balkema. Spampinato, S., Gambino S., Patane D.. Privitera, E.. D'AmiCO G., Di Prima, S., Pellegrino, A., Scuderi, L., & 0. Torrisi 1998. Etna - Seismic activity. Acta Vulcanol. 10:149154.
shown) and in Acqua Grassa spring (Fig. 2c), which present R& background values relatively high, the relative variations of 3He/4Heratios are very limited and do not match those of He contents and pC02. During the 1996 paroxysmal phase, some of the studied samples show important variations, even stronger than in previous eruptive crisis. Unlucky, due to some gaps in the sampling sequence, only S.Giacomo well show a complete trend to be described during all the period. In this site, a decrease of pC02, 3He content and R/R,ratio can be observed in the first months of 1996, followed by an opposite strong and concordant variation in JulyAugust. In Petrulli well, though the sampling frequency is defective, 3He contents and lUR, ratios in 1996 are significantly higher than 1995 values. In Guardia well (Fig. 20, a sharp negative peak in pC02 and 3He content and R/R,ratio is observed in August 1996, followed by a strong increase in September 1996, immediately after the phase of strombolian activity and lava flows at summit craters (La Volpe et al. 1999). 5 CONCLUSIONS Periodic measurements of both He isotopes and He and CO2 contents in groundwaters reveal variations related to the renewal of volcanic activity. In particular, He and CO2 contents generally show an increment that is interpreted as a new input of magmatic fluids. Mostly during the 1996 phase, magmatic input is thought to be very important, with a significant increment of dissolved 3He and CO2 concentrations. In some cases (Guardia and S.Giacomo in 1996), the resumption of volcanic activity is paralleled by a sharp decrease of dissolved gas contents and €UR, ratios, followed by a similar positive variation. This fact may result from an air dilution produced by stress-induced fracturing, which precedes the new input of volcanic gases. At last, although our data result incomplete and discontinuous in many cases, some significant aspects could be pointed out. By comparing different sampled waters, showing large variability in all the studied parameters, the drainage gallery S.Giacomo seems to be the most interesting. Its abundant flow and peculiar topographic location (beneath the Valle del Bove depression rim) makes it particularly sensitive to the effects induced by variation of volcanic activity, thus appearing the most suitable for monitoring purposes. REFERENCES Aiuppa, A., Allard. P., D'Alessandro, W., Michel, A., Parello, F., Treuil. M. & M. Valenza 2000. Mobility and fluxes of major. minor and trace metals during basalt weathering at
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Magadi and Suguta: the contrasting hydrogeochernistry of two soda lake areas in the Kenya Rift Valley W.G .Darling British Geological Survey, Wullingford, UK
ABSTRACT: The Magadi and Suguta soda lake areas respectively form the southern and northern terminal drainage basins of the main Kenya Rift Valley. Magadi is a perennial lake but with an extensive trona crust. Thermal springs occur at many points around the lake. The northern Suguta Valley has surface water only intermittently, apart from the very small Lake Logipi. Thermal springs tend to discharge onto sediment fans. Chemical and stable isotopic data fi-om the springs of each basin have revealed some fundamental differences. Chemically these can be summarised as a greater loss of carbonate in the Suguta and a major loss of sulphate at Magadi, while isotopically the waters of the Suguta are more evaporated than those of Magadi. The hydrogeological habitat of the Suguta promotes the recirculation of waters, leading both to the removal of carbonate in surface crusts and to a well-oxygenated system in which sulphate can flourish. The latter does not occur at Magadi, leading to large amounts of sulphate reduction.
1 INTRODUCTION
2 RESULTS
Soda ash (Na2C03) is an important raw material for the chemical industry. While it can be made synthetically, production from soda lake deposits is becoming increasingly important (Kyle 199l), and any additional information on their systematics is therefore of some interest. Two such deposits are found in the Suguta and Magadi areas, which form respectively the northern and southern terminal drainage basins of the main Kenya Rift Valley. Under today's climatic conditions, both these basins suffer a great excess of potential evaporation over rainfall (Vincens & Casanova 1987, Casanova et al. 1988). Magadi is at present a perennial lake (Fig. la). However, with a salinity of up to 290 gA evaporite deposits of sodium carbonate (mainly trona) have formed over much of the surface. Thermal springs occur at many points around the lake, where their input of relatively dilute waters prevents a thick trona crust from forming. The northern Suguta Valley (Fig. lb) has surface water only intermittently, apart fiom the very small Lake Logipi at its northern end, which usually retains some perennial water. Springs are thermal and most discharge onto sediment fans with surface crusts of trona and other carbonates. W i l e each area appears to be a simple variation on a soda lake theme, chemical and isotopic data on the thermal spring waters of each basin have revealed some fundamental differences.
2.1 Chemistry Table 1 gives major element analyses for thermal springs in both areas. Alkalinity is expressed as HCO3, though carbonate must dominate. It is immediately apparent that Ca and Mg barely exist in solution because of removal by precipitation of their carbonates. Otherwise, salinities are up to 20 and 45 g 1-' for Suguta and Magadi respectively. When plotted against C1, presumed to be conservative, several species show clear differences between Magadi and Suguta (Fig. 2). There is removal of Na and HC03 at Suguta compared to Magadi, but also a major gain of SO4. Relationships are in general quite linear, showing that simple concentratiodremoval processes are operating (e.g. Jones et al. 1977). Other species such as K, and Si (not shown), can have rather less clear relationships with C1, implying more complex processes governing their concentrations at particular springs (ibid.). 2.2 Stable isotopes Values of 6l80, ?j2H and 6l3C~lcfor both areas are given in Table 1. Figure 3 shows that the 0 and H values range from relatively depleted (-3.7 %O 6l80) to highly enriched (+3.7 %o) along an evaporational trend. By contrast, 6l3C~rc values have a narrower
95
Figure 2. Plots of Na, HC03, SO4 and K versus C1 for the thermal springs of Lake Magadi and the Suguta Valley, in mmol 1-'.
range: ultimately they reflect the dominance of silicate hydrolysis brought about by the locally high flux of mantle CO2 (Darling et al. 1996). This is characterised by 6I3C values of around -3 %o, which accounts for the similarly enriched values found in the DIG. However, these reveal little about differences in processes between the two area and are not considered further here.
Figure 1. Maps showing the location of thermal spring Sampling sites (a) at lake Magadi, southem Kenya, and (b) in the Suguta Valley, northern Kenya.
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Table 1. Temperature, pH, major element chemistry and stable isotopic data for the Magadi and Suguta thermal springs. Alkalinity expressed as HC03 Sitename
Site Temp No. "C
Lahz Magadi LittleMagadi 3 NE Lagoon 11 N Lagoon 73 Graham'sL. 13 BirdRockL. 16 SW Lagoon 18 NW Lagoons 9 Suguta Valley LogipiNW 190 LogipiNE 238 Elboitong N 237 Elboitong S 236 Kageinya 188 sv3 185 Kamuge 187 Namarunu 189
pH
Na
K
Ca
85 67 63 40 41 45 45
8.85 8.95 8.90 9.55 9.60 9.55 8.80
11100 11100 13300 12200 12700 12300 9640
220 157 240 100 109 115 112
0.4 0.6 0.5 0.6 0.7 0.6 0.7
61 70 92 95 68 68 50 66
8.30 8.85 9.00
4170 6200 6690 6770 5420 1160 925 5500
79.4 120 126 129 85.5 12.2 6.9 113
0.5 0.5 1.4 0.8 3.0 1.0 4.2 0.5
9.50 8.25 7.75 8.80
Mg HCO, CI mgl-' <0.1 <0.4 <0.7 <0.4 <0.4 <0.4 <0.4
19600 5350 20900 5850 24900 6600 23100 6400 23800 6450 23400 6450 16700 4900 5160 5720 8240 7940 3350 2290 1740 6540
0.1 <0.1
0.5 <0.1
0.3 <0.1
1.3 0.2
30 20
Zj2H%O 10 0 -10
-20
0
1
2
3
Si
Li
B
F 6 ' * 0 6*H 6I3C 960
151 134 132 160 169 189 204
3420 590 5250 920 4500 1060 4400 1070 2280 3700 250 140 160 38 3200 814
33.0 39.0 35.2 23.8 25.6 32.0 49.2
1.30 7.8 0.78 7.8 1.17 9.2 0.16 6.8 0.17 7.7 0.16 7.4 0.45 7.7
180 -0.7 130 -1.0 -0.1 -2.4 140 -2.4 130 -2.5 80 -3.1
18.5 0.05 5.1 110 30.8 0.15 6.8 82 58.8 0.25 6.8 87 46.2 0.31 6.8 87 81.5 0.23 13.6 190 54.0 0.26 0.8 53 38.4 0.22 0.5 37 26.0 <0.01 6.5 140
+2.2 +2.3 +3.7 +3.6 +1.7 -2.7 -3.7 +3.0
-3 -7 -2 -16 -14 -15 -17
-3.3 -1.8 -1.7 -1.7 -2.0 -3.2
+15 +I4 +20 +21 +9 -13 -19 +14
-1.2 -2.0 -1.5 -1.6 -4.0 -4.2 -4.7 -0.9
The reason for this difference is connected to contrasts in hydrology between the two areas. At Magadi it is apparent from the 0 and H isotopic evidence presented here (and also by Hillaire-Marcel & Casanova, 1987) that spring waters flow more or less directly fi-om the Rift flanks (except for the more northerly springs which may be fed partially by underflow from the evaporatively-enriched Lake Naivasha. Jones et al. 1977, Darling et al. 1990). They acquire much of their solute load by flow through the extensive evaporite deposits of the area. Because the springs discharge more-or-less directly into the lake (Fig la), there is nowhere to lose carbonate other than through trona deposition at the lake surface. In the Suguta valley, the thermal springs do not discharge directly into a lake. They are fed by a combination of recharge on the rift flanks and leakage from the Suguta River, which contains an element of evaporatively-enriched water from Lake Baringo (Darling et al. 1996). However, the spring waters are isotopically heavier even than those of the northern Magadi springs. This additional enrichment is likely to be due to an element of recirculation in waters of the Suguta springs. During such recycling, loss of sodium and (bi)carbonate as trona etc. would occur relative to chloride. In the Suguta sulphate shows a broadly linear, apparently evaporation-associated increase (except for Kageinya), whereas at Magadi there has been a considerable loss (Fig. 2). Bacterial action under reducing conditions has been proposed as the most likely cause of sulphate depletion in the case of Magadi (Jones et al. 1977): if more recirculation is occurring in the northern Suguta Valley, this would presumably result in a greater amount of oxygen in the system and therefore much less scope for bacterial reduction to operate.
40
- 4 - 3 - 2 - 1
SO4
4
6I8O %o Figure 3. Delta-plot for thermal spring waters in the Magadi and Suguta areas. WML is the world meteoric line (Craig 1961).
3 DISCUSSION At Magadi, simple evaporative enrichment in Na was shown by Jones et al. (1977) to exist over several orders of magnitude from dilute rift-boundary streams to lake brines. Over the same range there was a small loss of carbonate species (ibid.). The losses in these two species are significantly greater in the Suguta springs, and are clearly related to the precipitation of Na in the form of carbonates such as trona and nahcolite.
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ACKNOWLEDGMENTS
Perhaps surprisingly for an area subject to volcanism, there is little evidence for major inputs of reduced sulphur in most of the Suguta Valley. However, the concentration of sulphate at Kageinya is so much greater than predicted by simple evaporation (Fig. 2) that it seems possible there has been a local input of HZS into the system which has subsequently become oxidised. In summary, therefore, the hydrogeochemistry of these terminal basin discharges suggests that the Magadi springs can be regarded as a ‘one-pass’ system compared to the ‘partial feedback’ of the Suguta springs. These differences, and those of the resulting soda depositional types, are to a great extent the consequence of geological factors. The highest concentrations of dissolved solids in the Suguta Valley are found in the Elboitong group of springs (opposite Namarunu, Fig. lb). These are at approximately half the concentration of the Magadi springs. There may be as much if not more soda in the Suguta Valley, which has a significantly greater area than Lake Magadi and its environs. However, without a relatively narrow, fault-bounded basinal structure such as exists at Magadi (Fig. la), it is difficult to envisage the Suguta trona deposits ever becoming as concentrated and commercially viable as they are at Magadi. Notwithstanding this, large-scale hydrological changes have occurred in both basins within the Quaternary alone, due to climatic fluctuations (Casanova et al. 1988, Hillaire-Marcel & Casanova 1987). Additionally, developments in rifting activity are difficult to predict, so the Suguta cannot be ruled out as a fkture soda resource.
Thanks are due to BGS colleagues past and present for field and laboratory assistance. The work was supported by the UK ODA (now DFID). This paper is published with the permission of the Executive Director, British Geological Survey (NERC).
REFERENCES Casanova, J., Hillaire-Marcel, C., Page, N., Taieb, M. & A. Vincens 1988. Stratigraphie et palaeohydrologie des episodes lacustres du Quaternaire recent du rift Suguta (Kenya). C. R. Acad. Sci. Paris 307: 1251-1258. Craig, H. 1961. Isotopic variations in meteoric waters. Science 133: 1702-1703. Darling, W.G., Allen D.J. & H. Armannsson 1990. Indirect detection of subsurface outflow fiom a rift valley lake. J Hydrol. 113: 297-305. Darling W.G., Gizaw, B. & M.K. Arusei 1996. Lakegroundwater relationships and fluid-rock interaction in the East African Rift Valley: isotopic evidence. J. African Earth Sci. 22: 423-43 1. Hillaire-Marcel, C. & J. Casanova 1987. Isotopic hydrology and palaeohydrology of the Magadi (Kenya)-Natron (Tanzania) basin during the late Quaternary. Palaeogeog., Paleaoclimatol., Palaeoecol. 58: 155-181. Jones, B.F., Eugster, H.P. & S.L. Rettig 1977. Hydrochemistry of the Lake Magadi Basin. Geochim. Cosmochim. Acta 44: 53-72. Kyle, J.R. 1991. Evaporites, evaporitic processes and mineral resources. In J.L.Melvin (ed), Evaporites, Petroleum and Mineral Resources: 478-533. Amsterdam: Elsevier. Vincens, A. & J. Casanova 1987. Modem background of Natron-Magadi basin (Tanzania-Kenya): physiography, climate, hydrology and vegetation. Sci. Geol. Bull. 40: 9-2 1.
4 CONCLUSIONS Both the Magadi and Suguta areas contain evaporite deposits which are re-mobilised by thermal waters. At Magadi the springs discharge on the edge of the lake, where water is removed by evaporation and trona is deposited on the lake surface. In the Suguta Valley, most springs discharge away from open lake water, resulting in the deposition of trona crusts but also allowing some recirculation of water. This leads to depletion of sodium and carbonate relative to chloride in the Suguta spring waters, but some isotopic enrichment of the water itself. It also permits the existence of more-oxidising conditions in the Suguta, which is reflected in the much higher sulphate concentrations compared to those of Magadi, where bacterially-mediated reduction dominates. The structural style of Magadi has allowed a concentration of trona deposits within a relatively small area. It would require tectonic and perhaps climate changes to bring this about in the Suguta Valley.
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Geochemical precursors of the 2000 eruption of Mutnovsky Volcano, Kamchatka G .M .Gavrilenko Institute of Volcanology,RAS, Petropavlovsk-Karnchatsky, Russia
M .G .Gavrilenko Moscow State University, Moscow, Russia
ABSTRACT: The Mutnovsky Volcano is one of the biggest and most active volcanoes of South Kamchatka. The most recent explosive eruption of the volcano was 1960. After that time the volcano is in a stage of weak eruption that has already been continuing for more than 40 years.The analysis of seismologic data and visual observations confirmed an eruption in March 2000. Within several hours of the eruption a steam-gas outburst reaching 2500 m was observed. It disappeared the next day. During the observations from a helicopter it was determined that the eruption took place in the northern explosive funnel of the volcano's South-Western crater which was active until middle 1950s. Hydrochemical sampling demonstrated that the eruption of March 2000 was preceded by a continuing annual increase in SOL2/ C1' and SOL2/ F- in waters of Vulkannaya River, which drains Mutnovsky's active craters. Possibly, one of the surface events led to the March 2000 Mutnovsky Volcano eruption was the glacier slide to the bottom of the North-Eastern crater in the end 1990s.
1 INTRODUCTION
lating, thus buffering and forming the fumarolic gas composition of the NE crater bottom lowtemperature fumaroles. The regime of these gas and also hydrothermal vents to a large extent depends on hydrological conditions of the crater, i.e., on intensity of water feeding and ground water table (Taran et al. 1992). SW crater was active until middle 1950s. An acid thermal lake with the diameter about 300 m and water temperature 42-46' C was located in northern part of this crater (Gavrilenko 1996). The SW and NE craters is drained by Vulkannaya River which accumulates practically all hydrothermal vents and streamlets of this craters.
The Mutnovsky Volcano is one of the biggest and most active volcanoes of South Kamchatka. Its height is 2323 m above the sea-level and it consists of 4 contiguous, successively forming stratocones with apical calderas and daughter intracaldera structures. All the volcanic mass is complicated with numerous cones of adventive eruptions (Seliangin 1993). Figure 1 shows a sketch-map of active craters of Mutnovsky Volcano: Active Funnel (AF), NorthEastern (NE) and South-Western (SW) craters. There are many fumarolic and hydrothermal vents in this craters. The AF can be considered as the upper part of an open magmatic system, representing a channel which is connected at an unknown depth to the magma body of unknown geometry. Intensive fumarolic activity (more than 200 t/day of S02) is concentrated in a large area of south-west wall of AF where a red glow was observed in the 1970s (Taran et al. 1992). Gas activity of the NE crater is also sufficiently, but it is apparently independent of that of the AF, i.e., it is determined by its own sources of heat and fluid. This can be interpreted by a cooling of though still active magma body which has a more prolonged history compared with the magma discharge vent of the AF. A local hydrothermal system is developing above this body. Inside the system highly mineralized acid boiling brines and melted sulfur are circu-
2 VOLCANO'S HISTORY AND ITS RECENT ERUPTIONS CHRONOLOGY As has been shown by consideration of the activity of Mutnovsky Volcano, it passed through its growth stage back in the Late Pleistocene. There were no outpourings of lava in the Holocene. The volcano was typified by explosive activity of a moderate nature with weak and medium strength eruptions @redominantly phreatic) predominating and separated by periods of hundreds (less frequently tens) of years. In historical times there have been not only phreatic eruptions, but also eruptions yielding juvenile material of basaltic composition.
99
volcanic structure (hydrothermally-changed volcanic rocks) indicate that there is great probability that this eruption is of phreatic character. 4 HYDROCHEMICAL DATA 4.1 River Sampling peculiarities
Figure 1, A - locality map of Mutnovsky Volcano cone and its craters, B - locality map of the North-Eastem crater floor (Donnoye Field); €3 - location of the March 2000 eruption; AF - Active Funnel; UWSS, LWSS - upper and lower watersample sites (black circles), accordingly; GF96, GF97 - 1996 and 1997 glacier fronts, accordingly,
The material of the most recent eruptions, beginning in 1904, was again resurgent. This is confirmed by written records for the eruptions of 1927-1929 and 1960. The most recent explosive eruption of the volcano was in 1960 (excepting the March 2000 eruption). However, the abnormally high evolution of heat (1 800- 1900 MW) in the period between paroxysms for volcanoes of this type indicates that the volcano is in a stage of weak eruption that has already been continuing for more than 40 years (Melekestsev et al. 1990). 3 SEISMIC EVENTS ERUPTION
AND
MARCH 2000
According to the data from a number of seismic stations, including radiotelemetric seismic stations located in the area of Mutnovsky Volcano, a powerful surface event with a tremor (1Hz and 0,l-0,9 mcds), which is still going on, took place at Mutnovsky Volcano on March 16'h, 2000, at 1856 GMT. The analysis of seismologic data and visual observations confirmed an eruption. Within several hours of the eruption a steam-gas outburst reaching 2500 m was observed. It disappeared next day. During the observations from a helicopter it was determined that eruption took place in the northern explosive funnel of the volcano's South-Western crater which was active until middle 1950s (Fig. 1A). (An acid thermal lake with the diameter about 300 m and water temperature 42-46' C was located inside of this funnel. The funnel has been filled with snow and ice for at last 10 years (Gavrilenko 1996; Gavrilenko 2000). The fact that the March 2000 eruption was not prolonged (about a day) and that the fragmental product erupted on the volcano's slopes was represented by sharp-edged pieces of a
Beginning in 1992, that is before Mutnovsky's eruption in March of 2000 and a year and a half after it, during the same season of August to September, the waters from Vulkannaya river were sampled for chemical analysis. The waters were sampled at two sites: in southern part of NE crater UWSS (upper water-sample site) and in the western part of the volcano-LWSS (lower W-S.S.) (see Fig.1). Simultaneously , the absolute discharge of chemical components dissolved in the water was measured at both sites. 4.2 Analytical methods The analytical methods applied were: titrimetry for sulfate- and chloride-anions, potentiomertry for fluoride-anion. 4.3 Results During the continuous studies of hydrochemical materials from the Mutnovsky Volcano it was observed that variations of quantity and ratios of SO<*, C1- and F- - anions from both sites of Vulkannaya River (UWSS and LWSS) during the years of 19922000 were similar. However, they had a number of differences as well (Table 1, Figs. 2,3). For example, the discharge of anions from UWSS before the eruption varied comparatively little. Only in 1998 a slight decrease of this parameter was registered for all three anions: for SOL2one and a half to two times, for F- - three times, and the discharge of C1- - fell significantly. After the eruption the discharge of anions sharply grew, tenfold (see Table 1 and Fig. 2). At LWSS the periodic variations of the discharges for C1- and F- anions are slightly similar to their variations at UWSS. Only the discharge of SOi2 anions behaved themselves differently, even ccillogically)).They grew up by 40% in 1998 and fell again in 1999, a half a year before the eruption. The discharges of main anions at UWSS and LWSS grew after the eruption but not significantly-two to five times more (see Table 1 and Fig. 2). However, the above-mentioned hydrological and hydrochemical data were not enough to predict the eruption of the volcano. In order to do so, it was necessary to utilize other hydrochemical data such as periodic variations of anions' ratios, data on general condition of the volcano, and events that took place before the eruption: for example, glacier-slide to the 100
Figure 2. Variations in time of SOi2, C1- and F- - anions discharge in the Vulkannaya River's water from 1992 to 2000. UWSS and LWSS - upper and lower water-sample sites, accordingly. Figure 3. Variations in time of SOi2/ CF and SOi2/F--ratios in the Vulkannaya River's water from 1992 to 2000. UWSS and LWSS: see on the Figure 2
bottom of NE crater in the and 1990s (see Fig. 1B). Table 1 and Figure 3 show that the ratios of s u l k to haloids in 1992-1997 were changing relatively steady: 7% to 20%. In 1998-1999 all these parameters grew significantly reaching the maximum values at UWSS: for SOi2/F- in 1998 and for SOi2/C1- in 1999. At LWSS the variation of maximums is different. It is a ((mirror image)). In 1998 the maximums were observed for S04-*/Cl. The maximum for SOi2/F-was observed in 1999. In a year and a half after the eruption, every single ratio of these anions significantly fell to the similar parameters at UWSS and LWSS (see Table l and Fig. 3). 5 DISCUSSIONS In connection with the above provided information, it becomes easy to understand the ((illogical)) and ((mirror image)) variations of hydrochemical parameters in the waters of Vulkannaya River. The following processes are to be considered: 1) falling of the discharges of main anions in waters of Vulkannaya River at UWSS before the eruption and it is characteristic for Cl-ions; 2) maximum, ((illogical))behavior of other anions, the discharge of SOL2 in 1998 and 3) the ((mirror image)) of maximum ratios of SO42/C1- and S0L2/F-over the period of time at UWSS and LWSS. All these processes are evidently connected to the glacier slide to the bottom of an active N.E. crater of the volcano. The beginning of the first process coincides with the years of the most of the glacier slide which correspondingly influenced the bottom of the N.E. crater in 1997-1998. Due to an additional pres sure from glacier masses and the increase in the level of ground waters on presurface magmatic masses, their degasification decreased but
insignificantly. It mainly affected the HCI. As a result of these events, the flow of somewhat cold ground waters to the magmatic bodies grew and further cooled down the degasification of this component. This process was most vivid at UWSS because river waters drain South-Westem crater here. The eruption took place in 2000. (see Table 1 and Fig. 2). As to the second process, the ((illogical))behavior of outlay variation of SO^-^ at LWSS compare to its behavior at UWSS, can be explained through the increase of ground waters level in the crater due to the glacier slide. As a result, the partial remobilization of soluble sulfates occurred in the lower horizons of crater-lake deposits. Our experiments with these deposits (neutral and slightly acid drawings from them) as well as some other research evidence that crater-lake sulfur deposits contain some percentage of soluble sulfates: Fe, Al, Mg, Ca, Na and other. Evidently the remobilization of SOi2ion from the crater deposits in North-East led to its discharge growth with waters of Vulkannaya River in 1998 at LWSS. It must be also noted that this process before the eruption somehow influenced the ((mirror image)) maximum ratios of anions over time (see Table 1 and Fig. 3). It is also possible that there were some other processes which did not manifest themselves so vividly.
6 CONCLUSIONS The March eruption of Mutnovsky Volcano was expected and, in a sense, predetermined. First of all, its last eruption took place 40 years ago, and from
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Sampling dates (August-September)
-------_--_-_---
o N\
Major anions discharge (in g/s) and its ratios U
F-
sop/ c1F'
SO/:' SO:-/
(CI'
+ F-)
Tf
rn
\D
z
z
m
m
z
160
m
. tm
z
cn 03
z
z
0 0
N
----
----
----
----
140
----
----
160
1000
515 18
700 23
560 13
480
----
----
620 14
850 4
650 5
1110 57
36 0.8
35
33 0.8
32
34 0.6
31 0.2
34 0.5
60 7.2
----
----
90
o\ m
200
---_-
c1-
z
180
----
sop
o m\
----
----
----
----
---_-
----
----
----
----
----
1.9 10.0
2.1 12.3
2.0 10.0
2.0 22.5
1.4 32.0
9.9 17.5
18.2 233
27.4 450
19.1 320
18.5 139
310 9.6
425 21.4
464 29.1
112 15.6
17.2
25.8
18.4
15.9
8.7
----
----
17.0 200
15.0
----
----
270 9.6
267 11.6
213
13.6
16.0
----
----
14.7 225
20.0
_------
1850s cyclic recurrence of the volcano's eruptive regime did not exceed 30-40 years (Melekestsev et al. 1990). Secondly, increase of thermal discharge of the volcano was noticed in the few years (Taran et al. 1993). Thirdly, (as shown by authors in the above paragraphs) the hydrochemical monitoring of the Vulkannaya River which drains the active craters of Mutnovsky also pointed out that in the nearest future we could expect the eruption. Fourthly, the unusual and massive glacier slide to the bottom of the active North-Eastem crater in the end of 1990s possibly triggered phreatic eruption (Gavrilenko et al. 2001). The glacier movement commenced in the fall of 1996. Its front began crawling over the northern half of the crater and damming the Vulkannaya River. As a result, a lake with the area of 2500 m2 and depth about 3 m was formed in the crater. By fall of 1998 the glacier tongue moved even further and propped against the opposite side of the North-Western volcanic crater, thereby covering most of its thermal areas. The increase of ground water level in the crater abruptly altered the hydrogeological condition of the volcano. Intracrateral deposits became more saturated with water, fumarole activity regime changed, and heat discharge of thermal areas was disturbed. It is possible that the increased the infiltration rate of ground water into the bedrock of the volcanic structure. Such alteration of the volcano's external condition (slide of a glacier into one of the active craters and subsequent alteration of hydrological situation at the volcano) at a time of its expected activation could be one of the reasons which provoked the phreatic explosion in Mutnovsky's SW crater in March, 2000 (Gavrilenko 2000). All these allows us to conclude that there is possibly a connection between the changes in hydro-
_---
----
----
----
----
----
----
----
-------
----
----
----
logical and hydrogeological environments in the active craters of Mutnovsky and its eruptions provided the dynamics and mass interchange of glaciers. A good example for this conclusion would be the coincidence of the time of the slide and its biggest influence on the bottom of the active North-Eastem crater while the glacier was full of extreme hydrochemical parameters of Vulkannaya River (see Table 1, Fig.s. 1, 2, 3). The observed phenomenon in a such vivid form was first detected and researched by the authors specializing in similar objects and partially it was reflected in this article. REFERENCES Gavrilenko, G.M. 1996. Poor-known data for the Mutnovsky Volcano Crater Lakes, Kamchatka. Abstracts. Chapman Conference: Crater Lakes, Terrestrial Degassing and Hyper-acid Fluids in the Environment. September 4-9, Crater Lake, Oregon. p. 34. Gavrilenko, G.M. 2000. Mutnovsky Volcano Awakened. Priroda. 12: 41-43 (in Russian). Gavrilenko, G.M., Zelensky, M.E. & Ja.D. Muravjov 2001. Glacier Movement at Northeast Active Crater of Mutnovsky Volcano, Kamchatka in 1996-1998: Causes and Consequences slide. Volcanology andSeismology. 2: 1-6 (in Russian). Melekestsev, I.V., Braitseva, O.A. & V.V. Ponomareva 1990. Holocene activity dynamics of Mutnovsky and Gorely Volcanoes and the volcanic risk for adjacent areas (as indicated by tephrochronological studies). Volcanology and Seismology. 9(3): 337-362. Seliangin, O.B. 1993. Mutnovsky Volcano, Kamchatka: New evidence on structure, evolution and future activity. Volcanology andSeismology. 15(1): 17-38. Taran, Yu.A., Pilipenko, V.P., Rozhkov, A.M. & E.A. Vakin 1992. A geochemical model for fumaroles of the Mutnovsky Volcano, Kamchatka. J. Volcanol. Geotherm. Res. 49: 269-283.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swefs & Zeiflinger, Lisse, ISBN 90 2651 824 2
Helium geochemistry applied to rust-mantle interaction in the Apennines (Italy) F.Italiano Istituto Nazionale di Geofisica e Vulcanologia (IJV.G.V.)- Sezione di Palermo, Italy
M .Martelli & P.M.Nuccio Dipartimento di Chimica e Fisica della Terra (C.F.TA.) - Universitci di Palermo, Italy
ABSTRACT: He isotope ratios in the volcanic areas of peninsular Italy display values 5 4.6 Ra, both in free gas and fluids included in recent lavas. These low values with respect to MOB-type mantle seem a characteristic of the local mantle rather than the result of interaction between a MOB-magma with crustal rocks. A mantle contaminated by subduction-related fluids is the most plausible hypothesis to explain the observed features. In non-volcanic areas, the relatively high 3He/4Heratios in Central Tuscany (up to 3.2 Ra) and in the Campano-Lucano Apennine (up to 2.8 Ra) are interpreted as due to crustal thinning and mantle uprising in the former case and to mantle melts intruded into the crust in the latter. Mantle fluid contribution is greatest in the perityrrhenian regions and decreases towards the periadriatic regions, coherently with the largescale geophysical properties and with the geodynamic setting of the region. 1 INTRODUCTION
In this work we collect new and edited 3He/4He data fiom the Italian peninsula, and discuss the crustmantle interaction in the volcanic and non-volcanic environments. The distribution of areas affected by mantle degassing is related to geophysical parameters that may suggest the uprising of mantle and/or of mantle melts. The geodynamic context that may reconcile the collected geochemical and geophysical distributions is also considered.
Helium isotope composition (3He/4He)in fiee or in dissolved gas is at present the clearest way to identifL the eventual presence of mantle degassing at the Earth’s surface. This is because the 3He/4Heratio of primordial mantle is about four orders of magnitude larger than typical radiogenic helium roduced in the crust. Crustal He is characterised by ?! He/4He 0.03 Ra (Ra = 3He/4He in atmosphere = 1.39~10-~), increasing 3He/4He testifies increasing mantle 2 RESULTS AND DISCUSSION contribution and the MOR mantle signature is 8 Ra (O’Nions & Oxburgh 1988). Italy presents a variety of geological 2. I Volcanicareas environments and, in this respect, the investigation of Volcanoes in peninsular Italy are concentrated in the the 3He/4He distribution appears very usehl to perityrrhenian margin, from Campania to Tuscany, understand better the crust-mantle interaction in this with the exception of Mt. Vulture, which is located region. Geodynamics of the Italian peninsula is in the inner parts of the Southern Apennines. Only mainly characterised by the subduction towards the Vesuvius, Phlegrean Fields and Ischia volcanoes, W and NW of the Ionian-Adriatic plate at least since located in the Campania region, are active. The value the early Miocene, generating both the accretionary of 613Cc02in the volcanic fluids of Campanian and wedge of the Apenninic chain and the internal Roman Province ranges between +1 and - 2 %O vs. Tyrrhenian back-arc basin (e.g. Doglioni et al. 1996). PDB (Allard et al. 1997) with the exception of Ischia The main geological elements of the Apennines are (613C~~2= -2.5 -+ -5 %o, Inguaggiato and Pecoraino particularly clear in its southern portion where, from 1998), testifLing a clear shiR from the MORB values West to East, the sequence sedimentary chain (-4 + -8 %o). The 3He/4He ratio in Ischia ranges foredeep - foreland is recognisable. Recent (< 1 Ma) between 3.5 and 4.6 Ra (Marty et al. 1994; Italiano to active potassic and ultrapotassic volcanics are et al. 2000). In Phlegrean Fields hmaroles is -2.9 Ra concentrated along the western margin of the (Italiano et al. 2000), while in Vesuvius is -2.7 Ra peninsula, from Campania to south Tuscany (e.g., (from both hmaroles and gas included in recent Peccerillo 1999).
-
-
103
of erupted rocks. However, the observed 3He/4He lavas, Graham et al. 1993). These values are among values cannot be reconciled with a metasomatism age the lowest so far measured from active volcanoes all of -0.3 Ma, because the time required to lower the over the world. 3He/4He ratio measured in fluids from volcanic areas more to the north of Campanian 3He/4Heratio from MORB values to 3 - 4 Ra may be roughly estimated in the order of 10' years. Probably, volcanoes are also well below the MOR value the time estimated with Th-U disequilibrium is (Roccamonfina and Vulsini < 0.5 Ra, Alban Hills < 2 relative to the last metasomatic event. We believe Ra, Mt. Amiata - 1 Ra, Fig. 1). that a combined study of 3He/4Heevolution with time Different processes may explain these features: a) and of Th-U disequilibrium methods could concretely magma aging with in situ 4He production; b) contribute to enlighten the evolution of the mantle in assimilation of crustal rocks during magma transfer the study area and its metasomatic processes. and storage; c) contamination of the mantle source The extinct Mt Vulture, the only volcano located by interaction with the crust or with crustal derived in the eastern part of Apennines, displays a 3He/4He fluids. In this regard, Tedesco et al. (1990), value of 5.8 Ra (Doglioni et al. 1996), indicating that observing the similarity of 3He/4He among the Campanian volcanoes (Vesuvius, Ischia, Phlegrean the source of this magmatism has a considerable geochemical difference with respect to that beneath Fields), suggest that the mantle beneath this region the Roman and Campanian Province. In particular, might have an unusually low 3He/4Heratio. Graham we may suggest a minor contamination of the mantle et al. (1993), observing the relatively uniformity of source beneath Mt. Vulture in respect to the other He, Sr, Nd and Pb isotope composition of volcanics volcanoes. This is in accord with geochemical data of of Mt. Vesuvius, suggest the presence of an enriched Vulture rocks (De Fino et al. 1986). mantle. Marty et al. (1994), discussing the potential sources of metasomatic fluids for Campanian volcanism, suggest that they may originate from a 2.2 Non volcanic areas subducted continental-type crust rather than an oceanic-type or carbonate sediments. In fact these In spite of the typical sedimentary environment and two latter components are not capable to accumulate the absence of volcanic activity, 3He/4Heratios up to the required amount of radiogenic helium. Regarding 2.84 Ra are found in mofettes located in the southern the petrologic characteristics and the 0, Sr, Nd, Pb part of Apennines (Campano-Lucano sector). isotopic signatures of volcanic rocks from Roman Studying the gas geochemistry of this region, Italian0 and Campanian provinces, Peccerillo (1 999) affirms et al. (2000) observe that the 3He/4He ratios from that the upper mantle in the study area appears to this part of the chain overlap the magmatic values have undergone repeated and long-lasting tracked by Graham et al. (1993) for the Vesuvius metasomatic events due to upper crust-derived fluids. volcano and conclude that the high mantle gas fluxes On the basis of volcanisin age and on Th-U are due to the presence of mantle melts intruded into disequilibrium method on the rocks, Villemant and the crust along tensile lithospheric faults. In the same Flehoc (1989) and Peccerillo (1999) suggest a very paper, the authors propose that melt is generated in recent (ca. 0.3 Ma) metasomatic enrichment for the mantle wedge beneath the Apenninic accretionary magmas of the Roman Province prism, due to asthenospheric currents that push Here we emphasise that the 3He/4Heratio shift by towards the East the subducting Adriatic plate. U and Th decay and 4He accumulation after magma Another area of high mantle degassing has been aging is extremely inadequate for typical primary found in the Northern Apennines, where 3He/4He mantle-derived melts. In fact, taking into account the ratios as high as 3.2 Ra have been measured low U-Th content of MO-, melts could reach the (Larderello geothermal area, Hooker et al. 1985). observed low 3He/4He ratios only hypothesising an The He signal in this area is associated with other unrealistically large degassing of mantle-type helium, peculiarities of the Central Tuscany, such as presence followed by extremely long residence time in a large of granite dikes younger than 4 Ma, very high heat magma chamber. However, for several of the flow values (> 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.
105
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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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
I .2 Approach A modified version of the chloride mass balance method was tested to determine whether it is possible to estimate recharge usiiig this method. Chemical data from rainwater, surficial salts, runoff and groundwaters were collected, and the chloride mass balance approach was used by integrating these var ous data into geochemical mass balance models using NETPATH (Plummer et al. 1991). NETPATH is useful in this regard because it can calculate the net geochemical reactions which can lead to the change in chemistry of waters from one point to the next using a mass balance approach. Calculations can take into account mineral dissolution and precipitation, mixing of water and evaporation and dilution. Since traditional chloride mass balance approaches are based on evaporation as the reason for changes in chloride concentrations along the water evolution path, it seems useful to consider mineral dissolution and precipitation as possible contributors to chemical variation. Evaporation was evaluated independently by comparing the stable isotopic compositions of the rain water and the ground water using the standard equation for Rayleigh distillation.
Figure 1. Location map of the area.
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2.3 Isotopic considerations
2 METHODS
The stable isotopes of oxygen and hydrogen were analyzed from the rain, spring and well samples in the early winter. These data were compared with published results of rainwater in the region, and the Rayleigh distillation equation was applied in order to determine the extent to which rainwater needs to evaporate to achieve the isotopic signatures seen in the ground water.
2.1 Determining end-members The most critical component of the research is the characterization of salts deposited in the surficial deposits in the area. Over 50 samples of surface deposits were collected from the drainage basin. These samples were soaked in deionized water and the resulting solutions were analyzed for Cl-, HCOj-, sod=,Mg", Ca", Na' and K'. The results were plotted on a trilinear Piper diagram and three end members were chosen. All other samples can be considered to be admixtures of the three endmembers. While it is true that these end members are not real defined mineral phases, they do repr sent a reasonable approximation of the chemical i puts caused by dissolution into water going into the shallow aquifers. Rain water was sampled at the meteorological station at Safawi. This water was compared both to the published data from Jordan as well as with labile salts at two terminal playas in the area. All three types of data agree fairly well with each other and are notably different than the salts present in the surface deposits. Surface runoff samples were collected at two locations to determine their chemical natures and to try to determine the major contributors to their solute loads. Waters from the spring and selected pits were sampled in the late summer, as well as in the winter of 2000 (January and March). These data were used to see how chemistry changed as the result of recharge events and thus determine the sources of solutes added to the water after recharge. The results were compared with results of stable isotopic analyses.
3
RESULTS
3.1 Salts The surficial deposits contain a number of different types of labile salts. Empirically, these have been divided into three end-members as shown in Table 1. It is difficult to assess these with charge balances because it is difficult to isolate carbonates present in the surficial salts from CO1 dissolved from the atmosphere during the extraction of the salts. While this may cause some consternation, it is useful to point out that the issue being looked at is the source of chloride and not the source of dissolved carbon. Moreover, the "phases" being input into the NETPATH models are not argued to be real phases, but convenient proxies for the mixtures of soluble material present in the surficial deposits. Table 1. Salts present in the surficial deposits, normalized to 1 mole chloride. Salt 1
Salt 2
Na
0.04
0.56
Ca
0.36
0.54
m
0.07 0.09
0.05 0.56
1.00 0.03 1.63
1.00
K
c1
2.2 NETPATH Modeling
NETPATH was used to simulate changes in the chemistry of the spring and ground waters caused by the recharge event. Components used were Cl-, HC03-, Sod=, Mg", Ca", Na' and K'. Phases used were the three salt end-members described earlier, calcite, gypsum, carbon dioxide and a potassium silicate. Two types of models were simulated. The first set did not take into account the salt end me bers but relied on allowing evaporation to explain changes in chloride contents in the input waters. The second set used the salt end members as well as evaporation for this purpose. Thus it is possible to assess whether chloride input into the system can be explained by evaporation, as most currently used models assume, or whether salts stored in the surficial deposits are a more important source of chloride.
SO4 HC03
0.00 4.42
Salt 3 0.01 0.28 0.03 0.04 1.00 0 . 64 1.45
The salts range from basically a calcium-choride salt (1) to Na-K-Ca-chloride (2) to a calcium-sulphatechloride salt (3). 3.2 Rain and runoff Results of analysis of rainwater and the two runoff samples analysed are shown in table 2. It is noteworthy that the two runoff samples are significantly different from each other. The second runoff sample has lower chloride content than the rainwater, rendering it useless as an input function for chloride in the groundwater. Runoff 1 is used to determine whether elevated chloride concentrations can be attributed to dissolution of salts in the surficial deposits or to evapoconcentration.
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3.4 Isotopes
NETPATH models of chemical change from rainwater to runoff 1 indicate that anywhere between 83 to 88% of the initial rainwater needs to evaporate, and a certain amount of surficial salts need to dissolve into the water, to get the needed chemical changes observed.
The results of isotopic analysis of rainwater and goundwater from the study area are presented in T ble 3. Tritium data clearly shows that the water is of modern origin. Stable isotopic data for the rainwater sample is consistent with other data from the area reported by Bajjali (1990). The relatively heavy isotopic signature of this water can be attributed to the climatic conditions in the area. Using the rainwater results as an input, the relative fraction of water which evaporated prior to recharge was calculated using the Rayleigh fractionation equation (Clark & Fritz 1997). The fractionation factor of 1.0098 was used (Mojabe 1971). The mo els suggest that the residual fraction is anywhere between 77.2 and 95.5%. This is a very high number considering that arid environments typically have recharge rates of less than 5% (Allison et al. 1994).
Table 2. Rain and two runoff samples. Concentrations in mmol/l. Safawi rain Nd
0.13
Ca Mg
0.27 0.17 0.05
K c1 SO4
Runoff 1 15.66 1.32 1.52 1.28
Runoff 2 0.83
0.59 0.26 0.34
0.34
4.80
0.14
0.12
1.47
0.48
3.3 Groundwater
Table 3. Isotope results from a rainwater sample (dated January 4) and waters from the spring and wells in the area. f is calculated using the Rayleigh distillation model.
The chemical nature of the wells and spring is different from that of the rain and runoff water in that the relative proportion of Mg++ is lower than the runoff and the relative proportion of Ca++is much lower than that of the rainwater. These distinctions are what allow the modeling of chemical changes through the hydrological cycling with a good degree of confidence. In order to better constrain the source of chloride in the groundwater, I decided to monitor the change in the chemical characteristics of the groundwater before and after a known recharge event with well constrained rainwater and runoff water properties. Changes in the chemistry were modeled using NETPATH. Waters were collected on January 17, 2000 and March 9, 2000. A runoff sample of January 17 was used and a rainwater sample from the same time. NETPATH models were simulated in two ways, the first to allow runoff water values to cause the change in chemistry and the secondto allow rai water chemistry to allow these changes. The results of the models show that runoff water as determined from the sample collected needs to be diluted in o der to fulfil1 the model. Moreover, the precipitation of gypsum is required in order to make the model work. This suggests that the runoff sample does not accurately reflect an input factor for chloride, and its precursor rainwater is a more suitable input. When rainwater was used as an input source for the model, the results clearly indicate that almost all of the chloride added to the groundwater after the recharge event can be attributed to the salts present in the surficial deposits.
Sample Date
04/01/00
Tritium
6l80 vs. SMOW
-4.04
17/01/00
2.76 5.47 6.18
17/01/00
8.43
-2.66
17/01/00
7.11
-3.16
17/01/00
-3.38 -1.53
6D vs. SMOW
-21.7 -15.4 -8.9 -15.9 -16.5
f
1.000 0.955 0.172 0.869 0.910
4 CONCLUSIONS The data collected and analyzed herein indicates that the major source of chloride to the water in the Ghussein well field is from salts stored in the surficial deposits in the area, and not from the wet precipitation that falls there. While the sources of the salts is not clear, they may be due to either dry deposition from varying sources in the region or due to the weathering of the local basaltic bedrock in the area. The implication of this is that chloride can't be considered to be a conservative tracer simply i creasing in concentration as wet deposition evaporates. While this may be true of the runoff sample collected, it is not so for the waters in the spring and in the wells. The NETPATH models clearly show that evaporation is a minor factor in chloride concentration increases.
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Given this, calculated evaporation factors for the groundwater are less than 25%, and thus over 75% of the rainfall which does not leave the area as ru off is recharged into the groundwater. Traditional chloride mass balance approaches would have yielded much higher evaporation factors. The stable isotopic data is consistent with that of the NETPATH models. Thus it is worthwhile to take into account other chloride sources in the future when attempting to estimate recharge in arid zones such as this study area. ACKNOWLEDGMENTS I would like to thank Mr. Wael Azaizeh for his help in the field and laboratory analysis. Financial and logistical support has been provided by the Badia Research and Development Programme (Jordan). Financial Support has also been provided by the International Development Research Centre (Canada) and by Yarmouk University.
REFERENCES Abu-Jaber, N., Jawad Ali, A. &K AI Qudah 1998. Use of solute and isotopic composition of groundwater to constrain the groundwater flow system of the Azraq area, Jordan. Ground Water 36: 361-365. Allison, G., Gee, G. &S Tyler 1994. Vadose-zone techniques for estimating groundwater recharge in arid and semiarid regions. Soil Sci. Soc. Am. J. 58: 6-14. Bajjali, W., 1990. Isotopic and hydrochemical characteristics of precipitation in Jordan. Msc. Thesis, University of Jordan. Amman-Jordan. Clark, I. & P Fritz 1997. Environmental isotopes in hydrogeology. Lewis Publishers, Boca Raton and New York. 328pp. Mojabe, M., 197 1 . Fractionment en oxygen-] 8 et en dueterium entre l’eau et sa vapeur. J. ChemicalPhysics 197: 14231436. Plummer, L., Prestemon, E. & D Parkhurst 1991. A interactive code (NETPATH) for modeling net geochemical reactions along a flow path. USGS Water Investigations Rep. No. 9 1-4087.
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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 2: Field-scale models of fracture networks in mineral deposits Byron R.Berger, Richard B .Wanty & Michele LTuttle U S . Geological Survey, MS 964 Denver Federal Ceizter, Deliver, CO 80225, USA
ABSTRACT: Predicting spatial properties of heterogeneous, mineralized rocks is important for many applications. Because each mineral deposit is unique, any analytical framework for studying coupled, water-rock systems must be sufficiently robust to account for variability within and between areas. Rock-water interactions are difficult to model, but a helpful approach to modeling fracture-controlled flow networks in mineralized rocks is to evaluate mineralized fractures and structural geologic systematics in and around specific mineral-deposit types. We propose that the necessary coupling of deformation, chemical transport, and heat transfer in hydrothermal systems sets constraints on fault and fracture systematics that must occur, and these systematics vary according to mineral-deposit type. We propose a set of field-based systematics at different spatial scales for porphyry-style copper and molybdenum deposits and related veins that help in field sampling design and data analysis.
1 INTRODUCTION
2 FIELD-SCALE FRACTURE MODELS
Predicting ground- and surface-water chemistry in mineralized rocks is challenging because of heterogeneity in the spatial properties (Wanty et al. 2001). Such challenges lead to implicit assumptions in field studies about scale dependence or scale invariance in hydrologic phenomena. However, variance in lithologic and structural geologic characteristics frequently contribute to a lack of correspondence between predicted and measured properties in field studies. Sensitivity analyses show that hydraulic conductivities and recharge areas are frequently the sources of the greatest uncertainty in hydrologic and hydrogeochemical modeling (Hill 2000). Thus, the effects of geologic structures on ground water flow systems are of considerable importance. The greatest heterogeneity in mineral deposits is at the microscopic and outcrop scales. At the larger scale of deposits and mining districts, the coupling of deformation, chemical transport, and heat transfer in hydrothermal systems set constraints on fault and fracture systematics that occur. The fracture systematics vary by mineral-deposit type, allowing for consistent field-scale fauldfracture models to be derived for a given type. This paper presents a fracture-network model of epizonal stockwork porphyry- and related vein-form deposits, and presents an example of a field study wherein this model is being tested.
Individual mineral deposits contain fracture networks that can vary from microscopic, short lengthscale vein-filled fractures to meters wide, long length-scale veins. Fracture density and interconnectivity can be very high as in stockworks and breccias to relatively low as along long trace-length normal faults. To serve as an analog for field-scale fluid-flow models, mineral-deposit fracture-network models must describe the spatial distribution of fracture/fault systems within a specific mineraldeposit type. Epizonal, hydrothermal mineral deposits are commonly found along strike-slip fault systems. The further development of fractures during mineralization follows a systematic pattern. Under regional tectonic compression, the strike-slip faults form typically at angles between 30-55" to the maximum principle far-field stress ( ~ 1 ) . When ichelon and parallel strike-slip faults mechanically interact, strain along them may be accommodated along linking extensional faults. Because these extensional faults are at low angles to 01, they should be hydraulically conductive. In fact, hydrothermal veins are often localized along such faults. Conversely, the presence of hydrothermal veins and alteration are direct evidence of hydraulic conductivity in the geologic past. The fracture networks that developed under
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Figure 1. Three spatial scales of fracturing in porphyry-style Mo-Cu deposit at Sierrita-Esperanza (S-E), Arizona USA. (a) Regional scale showing localization of deposit in releasing bend into right extensional stepover along right-lateral, strikeslip fault system. Contours are fracture density (cm-’). (b) Deposit scale preferred northeast orientation of veins generally parallel to far-field extension direction. (c) Sketch of hand sample showing early-stage, interconnected curved-trace fractures crosscut by straight-trace fractures (at S-E short curvedtrace network is generally absent). Data are in part after Cooper (1 973) and Titley (1990).
past tectonic regimes still exist, and are often the dominant hydraulically conductive features in the present. Thus, an understanding of the development of fractures and the systematics of their spatial extent and orientation facilitates an understanding of present-day hydrology at several spatial scales.
2.1 Porphyry-style mineral deposits Porphyry-style wall-rock alteration and ore zones may vary from 100s to thousands of millions of tons of rock. Figure 1 illustrates three scales of fracturing in the vicinity of and within these deposits using the Sierrita-Esperanza Mo-Cu deposit, Arizona, as an example. At the regional scale, strike-slip faults are the primary displacement zones (PDZ) localizing the deposits (Fig. la). Most frequently, they are within releasing bends into extensional stepovers along PDZs. The deposits form in vertically elongated cylindrical, medium- to coarse-grained porphyritic intrusive complexes. At deposit scale, the ore-bearing fracture networks form an upward-tapering, conical zone around a less intensely fractured, lower grade core, all within the cylindrical intrusive center. The ores are generally along straight fractures that tend to be preferentially elongated in the direction of the extensional faults that make up the stepover (Fig. Ib). At the microscopic and outcrop scale, the fracture networks consist of an early-stage dense, interconnected network of millimeter to centimeter length-scale curved fractures and subsequent centimeters to meters-scale, straight fractures (Fig. 1 c). Sulfide minerals are mostly along straight fractures.
Figure 2. Aerial photograph of open stopes on Mine Hill, Mahd adh Dhahab epizonal vein deposit, Saudi Arabia (from Hilpert et al. 1984). The stopes outline a “tensile-shear” fault mesh within an extensional stepover between left-lateral strike-slip faults. Extensional faults strike towards the top of the photograph and are segmented by oblique, synthetic and antithetic shear faults. The segmentation imparts the observed zigzag pattern of the stopes.
2.2 Vein-form mineral deposits Epizonal vein-form deposits frequently occur along tensile fractures. They may have strike lengths of kilometers and be up to tens of meters wide, and vary down to trace lengths of meters to hundreds of meters and widths of centimeters to a few meters. Associated alteration may affect many km3 of rock, with ore zones consisting of from 100s of thousands to millions of tons within the spatially more extensive alteration. Ore bodies are discrete entities within spatially more extensive vein networks. The ores are concentrated where the vein-controlling tensile fractures are segmented by synthetic or antithetic shear faults, producing a “tensile-shear mesh” (cf. Hill 1977) within an extensional stepover. Figure 2 shows a well-developed mesh at Mahd adh Dhahab, Saudi Arabia outlined by open stopes. Thus, vein ore bodies reflect the compartmentalization of hydrothermal fluid flow during mineralization.
3 FIELD APPLICATION At many localities worldwide, porphyry-style deposits are associated with overlapping, overlying and/or lateral vein deposits. The example below is the porphyry-style molybdenum deposit and related veins located in Redwell Basin near Mount Emmons, Crested Butte, Colorado USA (Fig. 3).
3.1 GeologicJFameworkof Redwell Basin The Redwell Basin deposits occur in a sequence of Cretaceous marine, clastic sedimentary rocks con138
sisting of silty sandstone, sandy limestone, and carbonaceous shale overlain by shale, thick bedded and massive sandstone, coal, and carbonaceous shale (Gaskill et al. 1967). The sedimentary rocks are intensely hornfelsed in the upper and lowermost reaches of Redwell Basin Creek due to intrusive activity. Composite igneous breccia pipes, consisting of rhyolite and rhyolitic intrusion breccia, crop out in the wall and floor of upper Redwell Basin and the porphyry-style mineralization is at depth beneath these pipes. Granodiorite porphyry dikes also crop out in the study area. 3.2 Porphyry- and vein-controlling faults The stockwork-veined porphyry molybdenum deposit (Sharp 1978) is localized in a zone of predominantly N5O"E-striking, left-lateral strike-slip faults in upper Redwell Basin where they intersect N30"W and north-south normal faults (Fig. 3). The northsouth normal faults continue north to a parallel in the releasing bend into the stepover. The MO deposit-
Figure 3. Geology of Redwell Basin, Colorado (from Gaskill et al. 1967). Redwell Basin Creek flows north from the cirque beneath Mt. Emmons. Porphyritic intrusions shown with letter b. Circular dashed lines show the location of Redwell Basin porphyry molybdenum deposit (A) in the uppermost part of the basin and zone of stockwork veining anomalous in molybdenum in the lowermost part of the basin. Angle drill hole location shown with crossed circle and mines as shaft symbol.
related breccia intrusion (b in Fig. 3) is elongated parallel to the stepover. Polymetallic Cu-Zn-Pb vein deposits occur to the southeast of the porphyry deposit and in the Daisy mine along the north-striking extensional faults (Fig. 3). 3.3 Hydrogeology and hydrogeochemistry
The Redwell basin is approximately 1 km wide and 2.5 km long. Surface water originates from several sources including springs and seeps, fracture systems feeding gaining stream reaches, mine workings, and artesian flow from a drill hole a few meters above Redwell Creek. The drill hole angles from the north to the south and penetrates the porphyry deposit. The average gradient of the creek is 28%. At the time of sampling, discharge varied from approximately 20 L/min to 2000 L/min.
3.3.1 Fracture-controlledJow Induration and metamorphism of the sedimentary host rocks precludes significant intergranular ground-water flow except over extremely long time periods, so the observed flow must be predominantly along fractures. This is corroborated by the groundwater chemistry that demonstrates compartmentalization and isolation of chemically distinct "packages" of water (Tuttle et al. 2000). For example, a natural spring in the bottom of the basin just west of the Daisy mine that occurs along north-south fractures north of the molybdenum deposit (Fig. 3) has a pH of 3.6 and conductivity of 200 pS and is actively precipitating ferricrete. A few meters to the south-southeast, a spring along a northwest-striking fracture system has a pH of 7.2 and conductivity of 40 pS and appears to discharge water from the unmineralized sedimentary rocks in the western wall of the canyon. To predict the current hydraulic conductivity, we estimate the present-day far-field stresses acting on the region using data in Zoback and Zoback (I 989). Our estimate implies that northeast-striking, northsouth, and northwest fractures should be hydraulically conductive. The inferred conductivities can be integrated with expected hydraulic gradients and the porphyry/vein fracture models to interpret and model the hydrogeochemistry of Redwell Basin. Mineralization in the upper reaches of the drainage should affect ground- and surface-water chemistry in the cirque area because of the elongation of stockworks and fault-veins to the north. In the intermediate and lower reaches of the basin, northweststriking fractures in the basin sidewalls should provide ground waters dominated by fracture flow from unmineralized sedimentary rocks to the northeast of the porphyry and vein deposits.
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4 CONCLUSIONS
3.3.2 Porphyry deposit-related waters Fluorine and chlorine in the mineralized porphyries together with sulfur isotopes allow distinction of mineralized ground-water sources from unmineralized sources. Fluorine is present in such minerals as fluorite and topaz in the porphyry-style mineralization, and upon leaching results in molar F-/Cl- ratios greater than 2 to as high as 9.5 in waters from mineralized rocks. Porphyry water has high concentrations of iron. Zinc accounts for 80% of the ore-metal load with Cu and Pb each accounting for 10%. 6 3 4 S ~in ~ 4the water is +2.7%0;sulfur in molybdenite in the deposit is +3.7 to +4.6%0(Stein and Hannah 1985).
In mineralized rocks, fracture flow dominates the ground-water regime. The combinations of fieldscale fracture models of specific mineral-deposit types with predictions of local hydraulic heads and regional flow patterns provide a framework within which to predict ground-water flow patterns and interpret hydrochemical data. In the instance of Redwell Basin, Colorado, the use of molar concentration ratios such as F:Cl in conjunction with sulfur isotope ratios in different styles of mineralization and unmineralized rocks provides a comprehensive model of ground water flow and an understanding of the natural geochemical state of the local area. Further, there is a basis for modeling of potential environmental effects from future mining within the basin.
3.3.3 Vein-related waters Mine drainages were sampled to characterize vein waters. Sulfate sulfur in mine workings ranges from +1 .O to -0.5%0. Mass balance calculations using sulfur isotopes indicate that mineralized waters in the upper reaches of Redwell Basin consist of 70% “porphyry water” and 30% other.
REFERENCES Cooper, J.R. 1973. Geologic map of the Twin Buttes quadrangle, southwest of Tucson, Pima County, Arizona. U.S. Geological Survey. Miscellaneous Geologic Investigations Map 1-745, Gaskill, D.L., Godwin, L.H. & F.E. Mutschler 1967. Geologic map of the OH-BE-JOYFUL quadrangle, Gunnison County, Colorado. U.S. Geological Survey. Geologic Quadrangle Map GQ-578. Hill, D.P. 1977. A model for earthquake swarms. Journal of Geophysical Research. 82: 1347-1352. Hill, M.C. 2000. Constraining models of ground-water systems using geologic information. Geological Society of America. Abstracts with Programs. 32(7): A337. Hilpert, L.S., Roberts, R.J. & G.A. Dirom 1984. Geology of Mine Hill and the underground workings, Mahd adh Dhahab mine, Kingdom of Saudi Arabia. U.S. Geological Survey. Technical Record USGS-TR-03-2. Sharp, J.E. 1978. A molybdenum mineralized breccia pipe complex, Redwell Basin, Colorado. Economic Geology. 73: 369-382. Stein, H.J. & J.L. Hannah 1985. Movement and origin of ore fluids in Climax-type systems. Geology 13: 469-474. Titley, S.R. 1990. Evolution and style of fracture permeability in intrusion-centered hydrothermal systems. In The role of fluids in crustal processes. Washington, D.C.: National Academy Press. Tuttle, M.L., Wanty, R.B. & B.R. Berger 2000. Environmental behavior of two molybdenum porphyry systems. In Geoenvironmental Analysis of Ore Deposits. Short course notes, 5“ International Conference on Acid Rock Drainage, Denver, Colorado, May 21-24. Wanty, R.B., Berger, B.R. & M.L. Tuttle 2001. Scale versus detail in water-rock investigations 1: A process-oriented framework for studies of natural systems. Proc. WRI-I0 (R. Cidu ed.). Rotterdam: Balkema. This issue. Zoback, M.L. & M.D. Zoback 1989. Tectonic stress field of the continental United States. In L.C. Pakiser and W.D. Mooney (eds.), Geophysical framework of he continental United States. Geological Society of America. Memoir 172: 523-539.
3.4 Interpretation using field-scale model As predicted by the porphyry-style deposit fracture model developed from the analysis of deposit data worldwide (B.R. Berger, unpublished data) and summarized in Section 2.1, the Redwell porphyrystyle system is localized in the releasing bend into an extensional stepover and elongated in a northerly direction. Further, the vein model presented in Section 2.2 predicts that vein mineralization should occur in the Redwell Basin along extensional faults segmented by northwest and northeast shear faults. When considered in light of predicted present-day hydraulic conductivities and hydraulic gradients, the fracture networks in mineralized rocks imply that ground and surface waters in the upper southern and southeastern parts of Redwell Basin should be dominated by fracture flow through mineralized rocks. These relations were corroborated by the hydrochemical data. Further, the hydrogeochemical data show that ground waters in the study area are compartmentalized and predominantly “porphyry waters,” “vein waters,” and “unmineralized rock waters” can be distinguished. Progressively downstream from the cirque, there is a continual input of depleted 34S water and values approach those of waters from the vein mines. This implies that vein waters mix along north-striking normal faults and northwest-striking segmentation faults with the porphyry waters. Yet farther downstream, where northwest-striking fractures conducting water from unmineralized rocks intersect the stream, the conductivity of the stream waters steadily decreases and sedimentary sulfur is introduced.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The Role of Pressure Solution in Fracture Healing: A Multi-Scale Reaction-Flow Modeling Approach O.Bildstein* & C.I.Steefe1 Lawrence Livermore National Laboratory, Livermore, California, USA (*present address: CEA Cadarache, St Paul-lez-Durance,France)
ABSTRACT: A multi-scale model is presented to investigate the healing of fractures in sandstones as a result of pressure solution (PS) mechanisms. The multiple scales include the grain-grain contact scale where the kinetic dissolution of minerals and the diffusion of dissolved species occur and the bedding scale in the vicinity of a fracture where mineral precipitation, transport, and the macroscopic stress field control the system. Using parameters from independent experiments on quartz dissolutiodprecipitation kinetics and diffusion in thin water films, the model is calibrated with experimental data from chemical compaction experiments and with observed porosity vs. depth trends in sedimentary basins. We present quantitative results highlighting the extent and time scales of fracture healing. We discuss the influence of the controlling mechanisms of PS (diffusion, precipitation or mixed-controlled), of fracture healing (diffusion or flow controlled), and the influence of the stress field close to the fracture (fracture stiffness). 1 INTRODUCTION
2 PRESSURE SOLUTION AT THE GRAIN SCALE
Pressure solution (PS) is a process coupling mechanical stress and geochemical water-rock interactions. It causes minerals to dissolve at grain-grain contacts as a result of the stress-enhanced mineral solubility (e.g. Sorby, 1863). Dissolved species then diffuse out of the contact, at which point they may (1) precipitate locally on the grain free face, andor (2) be transported over larger distances by diffusion in the pore space andor by advection. In this way, PS can cause pervasive porosity loss during burial (e.g. Weyl, 1959, Ramm, 1992). It can also provide an important source of dissolved species for localized diagenetic reactions resulting in transport from siltstones or clay layers to sandstones or transport from matrix to fracture networks. Observations show that PS is often involved in processes occurring close to faults and fractures such as stick-slip behaviors (e.g. Elliott, 1976, Angevine et al., 1982), fracture healing and vein formation (e.g. Carrio-Schaffhauser & Gaviglio, 1990, Peacock & Sanderson, 1995). We investigate here the effect of pressure solution on fracture healing in sandstones. Of particular interest is the role of the matrix compaction as a source of dissolved silica diffusing into fractures or veins (Elliott, 1973, Fisher & Brantley, 1992, Fisher et al., 1995).
2.1 Phenomenology There is still some controversy about the mechanisms involved in pressure solution. Two different models are generally invoked to explain PS: free face pressure solution (FFPS, e.g. Raj, 1982), and water film diffusion (WFD). The WFD models use different assumptions about the nature of the contact surface, e.g., planar surface, Weyl, 1959, Rutter, 1976, island-channel structure, Elliott, 1973, stress corrosion microcracks, Gratz, 1991). We chose to use the water film diffusion model with a planar surface contact. This mechanism has been invoked in several experimental studies to explain the observed displacement rate of an indenter and aggregate compaction rate (e.g. Gratier & Guiguet, 1986, Schuitjens, 1991). In this model, PS operates in three sequential steps: (1) dissolution of mineral at the grain-grain contact, (2) diffusion of dissolved silica in the water film between the grains, and (3) precipitation on the free face in contact with fluid (Fig. 1). The driving force for PS is the gradient in chemical potential which is due to a difference in stress between the grain-grain contact and the pore. Theoretically, the overall rate of PS may be controlled by any one of the three steps.
141
and the diffusion of aqueous silica in thin films (Farver & Yund, 1991) (Table 1). The water film thickness was inferred from both experiments (e.g. Horn et al. 1989) and theoretical work (Renard et al., 1999), and depends strongly on the effective stress (the difference between normal stress and pore pressure). Table 1. Kinetics and diffusion parameters measured in independent experiments and used in the reference case. Kinetic constant k(25"C) = 4.10-14mol/m2/s E,, = 89 kJ/mol
Figure 1. The water film diffusion model of pressure solution consisting in three sequential steps: dissolution and diffusion in the grain-grain contact, and precipitation in the pore space.
Diffusion flux 0 . 5 n m < 6 < lOnm 5.1O-I4 m2/s< D(25"C) < 2.1 O'9 m2/s Ead= 15 kJ/mol
2.2 Mathematical model The model assumes a rate law for mineral dissolution and precipitation based on transition state theory, and radial diffusion in the viscous water film at the contact (Bildstein & Steefel, 2000). At any point in the contact, rates are calculated taking into account the solubility of the mineral as a function of the normal stress. Quartz precipitation on the grain free face occurs at the pore pressure. The diffusion flux in the contact requires special treatment since diffusion has to be written in terms of the gradient in chemical potential:
Using these parameters, the model results provided a good match with experiments conducted over a wide range of temperature (150 to 360°C) and effective stress (15 to 50 MPa) (Bildstein & Steefel, 2000, in prep.). A major conclusion of these simulations is that pressure solution is not diffusion-controlled except in the experiments conducted by Gratier & Guiguet (1986) at high temperature (360"C), high effective stress (50 MPa) and with alkaline solutions. PS in experiments representative of diagenetic conditions (Schuitjens, 1991, Dewers & Hajash, 1995) shows a mixed control by diffusion and precipitation. Despite the good agreement between the model and the experiments using the parameters in Table 1, the comparison with the observed evolution of porosity as a function of depth in sedimentary basins shows that the kinetic constants used in the model are too high (reference curve in Fig. 2). This fact has been reported in other studies and leads us to consider an inhibition factor for dissolution (e.g. Dove, 1995) and precipitation (Renard et al., 1999). Some uncertainties remain as to whether the inhibition factors are related to chemical or physical processes and to what extent texture influences the strain rate (Bildstein & Steefel, in prep.). Nevertheless, we will use the inhibition factors from case 2 (Fig. 2) later in this paper, because they show the best agreement with the observations in sedimentary basins.
where Jd = diffusion flux; pf= density of water; D = diffusion coefficient; ai = activity of dissolved silica; and pj = chemical potential of dissolved silica. Following the Kamb-Paterson equilibrium assumption (Paterson, 1973), the chemical potential can be written:
where po = chemical potential at reference pressure Pyel; oc = normal stress at the contact; and Vj = partial molar volume of aqueous silica. These expressions are combined in the conservation equation for dissolved silica at the grain scale and solved on a radial grid using a finite volume formulation. At this point, the system consisting of the grain-grain contact and the adjacent pore space is considered to be a closed system. The textural parameicrs of the grain, represented by truncated spheres, are updated every time step in order to follow the evolution of the grain radius and truncature, the surface area of the contact, and the porosity.
3 PRESSURE SOLUTION, FLOW AND MASS TRANSPORT AT LARGER SCALES 3.1 Conceptual and mathematical model The pressure-solution module described above was coupled to CRUNCH (Steefel, in prep.) for the calculation of flow and transport by advection and diffusion in the porous media. Two coupled transport equations must be solved: one at the grain scale and the other at the macroscale. An additional source
2.3 Model calibration To build a reference case, we use parameters determined in independent experiments on the kinetics of quartz dissolution (Dove, 1994, Tester et al., 1994)
42
Figure 3. Matrix block symmetry with discrete fracture and computational grid (the actual number of grid cells is 100). Figure 2. Simulated and observed evolution of porosity as a function of burial depth. The reference curve uses k, = k, = 4.10-14mol/m2/s (no inhibition factor). The other curves use k, = 4.10-'* mol/m2/s, and respectively: (1) k, = 4.10-14mol/m2/s, (2) kd = 1.10-I6mOl/m2/S, (3) k, = 2.10-l' mOl/m2/S, (4) kd = mol/m2/s. 5.10-'*mol/m2/s, (5) k, = 1.
term was added to the macroscopic reaction-transport equation in order to take into account the contribution of PS (Bildstein & Steefel, in prep.):
a
-(4Ci)= at
R,,+R,,-V*(uCi-DVCj)
(3)
where 4 = porosity; U = fluid velocity; Ci = concentration of dissolved silica; R = precipitation rate of quartz in the pore space; and;lps = source term representing the flux of silica out of the grain contacts. At the grain scale, a source term representing the pore fluid transport was added to the conservation equation of aqueous silica in the pore space. These two equations are coupled using a sequential iterative approach. Finally, the fluid flow due to matrix compaction (i.e. porosity reduction) is calculated from the equation for the conservation of water. 3.2 Application to fiacture healing To model fracture healing or vein filling, we used a grid with the first grid cell representing a discrete fracture (Fig. 3). Flow of water out of the system is allowed only in the fracture where permeability is high. We assume that the matrix block is bounded by fractures at each end, thus producing a no-flux symmetry boundary condition in the middle of the block. In addition to the kinetic parameters of Figure 2, case 2, we use an initial grain diameter of 0.12 mm, a fracture aperture of 100 pm, and a fracture roughness factor of 5. All simulations are carried out at 150°C. The pore pressure in the fracture is fixed to 35 MPa and the 45 cm block matrix is initially subjected to a uniform, uniaxial, vertical compaction with a normal stress of 69 MPa. Under these conditions, the fracture heals in 400,000 years (Fig. 4).
The dominant mode of transport of aqueous silica from the matrix to the fracture is via diffusion since matrix compaction results in very low flow rates (maximum values of 5.10-6&year close to the fracture). The overall healing process is controlled by quartz precipitation and its duration is directly proportional to the roughness factor, and more generally to the kinetic constant and surface area in the fracture. As a result, the pore water and the fracture water is highly supersaturated with respect to quartz (saturation concentration x 4), and decreases very slowly with time. Putnis et al. (1995) suggested that nucleation required a higher supersaturation in small pores than in open fractures. The lower supersaturation required in open fractures would lead to higher rates of precipitation there than in the matrix and thus to even faster rates of fracture healing than obtained in Figure 4. Among the important parameters is the fracture normal stiffness, which expresses the resistance to displacements (e.g. closure due to normal stress). If we assume a low, constant fracture stiffness then a gradient in the normal stress develops in the matrix resulting in lower stress close to the fracture. This results in a slightly higher compaction rate towards the center of the block and a relative preservation of porosity close to the fracture (see Fig. 5 and field observations in Carrio-Schaffhauser & Gaviglio, 1990), without affecting the rate of fracture healing. 4 CONCLUSIONS Under most basin conditions, simulations presented in this work suggest that quartz precipitation in the pore space appears to be the rate-limiting step for pressure-solution. The model shows that silica diffusion from the matrix to fractures occurs at a length scale of only several centimeters and provides an important source of silica, enough to heal fractures in a 400,000 years time scale. The model also suggests that the compactional flow does not influence significantly the rate of fracture healing, regardless of the spacing between fractures.
143
Figure 5. Porosity profile and stress field in the matrix after 10,000 years of compaction plotted versus distance from the fracture.
Figure 4. Evolution of grain parameters in the matrix (immediately adjacent to the fracture) and the fracture aperture.
ACKNOWLEDGMENT
Fisher, D.M., Brantley, S.L, Everett, M. & J. Dzvonik 1995. Cyclic fluid flow through a regionally extensive fracture network within the Kodiak accretionary prism. J. Geophys. Res. lOO(B7): 12,881-12,894. Gratier, J.P. & R. Guiguet 1986. Experimental pressure solution-deposition on quartz grains: the crucial effect of the nature of fluid. J. Struct. Geol. 8(8): 845-856. Gratz, A.J. 199 1. Solution-transfer compaction of quartzites: progress toward a rate law. Geology 19: 90 1-904. Horn, R.G., Smith, D.T. & W. Haller 1989. Surface forces and viscosity of water measured between silica sheets. Chem. Phys. Lett. 162: 404-408. Paterson, M.S. 1973. Non-hydrostatic thermodynamics and its geologic applications. Rev. Geophys. Space Phys.. 1l(2): 355-389. Peacock, D.C.P. & D.J. Sanderson 1995. Pull-aparts, shear fracture, and pressure solution. Tectonophys. 241 : 1- 13 Putnis, A., Prieto, M., & L. Fernandez-Diaz 1995. Fluid supersaturation and crystallization in porous media. Geol. Mag. 132(1): 1-13. Raj, R. 1982. Creep in polycrystalline aggregates by matter transport through a liquid phase. J. Geophys. Res. 87(B6): 473 1-4739. Ramm, M. 1992. Porosity-depth trends in reservoir snadstones: theoretical models related to Jurassic sandstons offshore Norway. Mar. Pet. Geol. 9: 553-567. Renard, F., Park, A., Ortoleva, P. & J.P. Gratier 1999. An integrated model for transitional pressure solution in sandstones. Tectonophys 3 12(2-4): 97-1 15. Rutter, E.H. 1976. The kinetics of rock deformation by pressure solution. Phil. Tram Roy. Soc. Lond A.283: 203-2 19. Schuitjens, P.M.T. 1991. Experimental compaction of quartz sand at low effective stress. J. Geol. Soc. London 148: 527539. Sorby, H.C. 1863. On the direct correlation of mechanical and chemical forces. Phil. Trans. Roy. Soc. Lond. 12: 538-550. Steefel C.I., in prep. User's manual for CRUNCH: a software for modeling multicomponent reactive transport, Lawrence Livermore National Laboratory, California. Tester, J.W., Worley, W.G., Robinson, B.A., Grigsby, C.O. & J.R. Feerer 1994. Correlating quartz dissolution kinetics in pre water from 25 to 625°C. Geochim. Cosmo. Acta 58( 11): 2407-2420. Weyl, P.K. 1959. Pressure solution and force of crystallization, a phenomenological theory. J. Geophys. Res. 64: 20012025.
We thank Bill Glassley (LLNL) for his scientific interest and his support, and Joe Morris (LLNL) for providing a code to calculate the stress field close to fractures. This work was performed under the auspices of the U.S. Department of Energy by University of California Lawrence Livermore National Laboratory under contract No. W-7405-Eng-48. REFERENCES Angevine, C.L., Turcotte, D.L. & M.D. Furnish 1982. Pressure solution lithification for the stick-slip behavior of faults. Tectonics l(2): 151-160. Bildstein, 0. & C.I. Steefel 2000. Modeling pressure solution at the grain scale. AGU Spring Meeting, Washington D.C., May 30- June 3 . (www.Ilnl.gov/ees/pressure-solution. html) Bildstein, 0. & C.I. Steefel in prep. Modeling pressure solution: from the grain scale to the bedding scale. Carrio-Schaffhauser, E. & P. Gaviglio 1990. Pressure solution and cementation stimulated by faulting in limestones. J. Struct. Geol. 12(8): 987-994. Dewers, T.A. & A. Hajash 1995. Rate laws for water-assisted compaction and stress-induced water-rock interaction in sandstones, J. Geophys. Res. lOO(B7): 13,093-13,112. Dove, P.M. 1994. The dissolution kinetics of quartz in sodium chloride solutions at 25-300°C. Am. J. Sci. 294: 665-712. Dove, P.M. 1995. Kinetic and thermodynamic controls on silica reactivity in weathering environments, In A.F White & S.L. Brantley (eds.), Chemical Weathering Rates of Silica Minerals; Reviews in Mineralogy 3 1 : 235-290. Elliott, D. 1973. Diffusion flow laws in metamorphic rocks. Geol. Soc. Am. Bull. 84: 2645-2664. Elliott, D. 1976. The energy balance and deformation mechanisms of thrust sheets. Roy. Soc. Lond. Trans. A-283: 2893 12. Farver, J.R. & R.A. Yund 1991. Measurement of oxygen grain-boundary diffusion in natural, fine-grained, quartz aggregates. Geochim. Cosmo. Acta 55(6): 1597- 1607. Fisher, D.M. & Brantley, S.L. 1992. Models of quartz overgrowth and vein formation: deformation and episodic fluid flow in an ancient subduction zone. J. Geophys. Res. 97(B 13): 20,043-20,06 1.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Reactions governing the chemistry of waters interacting with serpentinites J.Bruni, M.Canepa, F.Cipolli, L.Marini, G.Ottonello & M.Vetuschi Zuccolini Deparment for the Study of the Territory and its Resources, University of Genoa, Italy
G . Chiodini Osservatorio Vesuviano,Naples, Italy
R.Cioni Institute of Geochronology and Isotopic Geochernistry,CNR, Pisa,Italy
A .Longinelli Department of Earth Sciences, University of Parma, Italy
ABSTRACT: The irreversible water-rock mass transfer accompanying the progressive evolution from rainwaters to neutral, Mg-HC03 waters and to mature, high-pH, Ca-OH waters was simulated through reaction path modeling, involving bulk dissolution of a serpentinite and precipitation of gibbsite, goethite, calcite, hydromagnesite, kaolinite, a montmorillonite solid mixture, a saponite solid mixture, sepiolite, and serpentine. The simulation includes two subsequent steps: Mg-HC03 waters are generated in the first step, under open CO2 conditions, whereas Ca-OH waters are produced in the second step, under closed-system conditions. In the second step, calcite precipitation depletes the aqueous solution in C species, allowing the concurrent increase in Ca and the attainment of the Ca-OH hydrochemical facies. 1 INTRODUCTION
2 SAMPLING AND ANALYSES
In a recent survey of the spring waters of the Genova province (Italy), many neutral Mg-HC03 springs and some high-pH, Ca-OH springs were found in association with ultramafic rocks, variably affected by serpentinization. All the springs are of meteoric origin as indicated by water isotopes and dissolved N2 and Ar. The association of these chemical types of waters with metamorphosed ultramafites is known (e.g., Barnes et al., 1978 and references therein; Pfeifer, 1977), but the processes generating Ca-OH waters are not fully understood. How do Ca-rich waters originate through interaction of meteoric water with MgO-rich, CaO-poor rocks, such as serpentinites and ultramafites? Why do these waters have such unusually high pH values? Besides, total dissolved carbonate concentration, Got, was not reported in previous analyses of Ca-OH waters with pH values of 11-12. In view of these unanswered questions and because of the possible use of high pH, CaOH waters in the sequestration of C02, we decided to study the chemistry of these springs. As first goal, here we investigate the water-rock interaction processes originating these peculiar waters.
Outlet temperature, pH, Eh, and TDS were measured in the field. Water was filtered through 0.45 pm membranes. A filtered portion was taken, and conc. HN03 was added to a separate portion. To measure Ctot of Ca-OH waters, these were sampled in evacuated bottles, containing conc. HC1 in excess with respect to alkalinity, leaving a head space of -XI of the bottle volume above the liquid. In this way, all the carbonate species are converted to C02, which accumulates in the head space. Ctotis obtained through mass balance calculations, measuring the Pc02, the volumes of head-space gas and sampled water, and assuming equilibrium distribution of CO2 between the gas and liquid phases. All the water samples were analyzed for: (1) Na, K, Mg, Ca, AI, Cr, Mn, Fe, Ni, Cu, Zn and Pb by either AAS or ICP-MS; (2) C1, S04, NO3 by IC; (3) Si02, NO2, N H 4 , Po4, by visible spectrophotometry; (4) alkalinity by acidimetric titration. Dissolved N2, Ar, and CH4, were measured in Ca-OH waters. Moreover, the lsO/’60 and D/’H ratios were also determined in selected springs. 3 WATER CHEMISTRY
Neutral Mg-HC03 springs have temperature of 5.320.6”C, pH of 7.0-8.5, PCo2 of 10”.5-10-2.5bar, and measured Eh of 150-250 mV 145
The high-pH, Ca-OH springs have temperature of 10.5-23"C, pH of 10.1-11.9, and Eh (computed on the basis of the S04'-/ HS- redox couple), of -390 to -516 mV. These waters have low Mg contents (0.008-0.2 mg/kg), and low Ctot (13.7 to 50.6 mol/kg), reflecting very low PCOZvalues, from 10to 10-'o.6bar. Sample C-1 1 exhibits Na(Mg)-OH composition, Mg of 5.8 mg/kg, Ctotof 334 pmolkg and Pc02 of 10-6.05 bar. Both chemical types of waters originate through low-temperature interaction of meteoric waters with serpentinites and ultramafites, but high-pH waters appear to be both more evolved and related to deeper aquifers than Mg-rich waters. Moreover, the evolution of meteoric waters towards the Mg-rich facies takes place under open-system conditions with respect to COz, whereas the subsequent shift towards Na(Mg)-OH and Ca-OH facies occurs in a system closed to CO:! exchanges (Pfeifer, 1977).
"
4 REACTION PATH MODELING To test this conceptual geochemical model, the irreversible exchanges between meteoric waters and a serpentinite were simulated by means of the Software Package EQ3/6, version 7.2 (Wolery and Daveler, 1992). Calculations were carried out assuming bulk rock dissolution, i.e., without any constraint on mineralogy and on dissolution rates of primary solid phases. The COM thermodynamic database was used, after including a serpentine mineral, whose hydrolysis reaction is: serpentine + 6 H+ = 2 SiOz(aq)+ 3 Mgz++ 5 HzO (1) and the temperature dependence of the thermodynamic constant of reaction (1) is (T in OC): = 40 - 0.1371 12 x T log Kscrpentine + 0.000187528xT2 + 5.68471E-07 x T3
(2)
The thermodynamic properties of this serpentine mineral are similar to those of the fine-grained, impure chrysotile-asbestos of Pfeifer (1977).
In a second step, the serpentinite was dissolved in the Mg-HCO3 water generated in the first step for log@ of -1.62, i.e., at incipient saturation with calcite. Temperature was kept at 12°C (25°C in a separate run), POZwas lowered to 10-73.6bar, and PCOZwas not fixed externally, to reproduce the conditions of the relatively deep circuits of high-pH waters. 4.2 The solid product phases Based on the general knowledge on the minerals forming through weathering and low-temperature alteration of serpentinites and ultramafites, the following solid phases were allowed to precipitate: brucite, gibbsite, goethite, kaolinite, sepiolite, serpentine, calcite, hydromagnesite, nesquehonite, an ideal solid mixture of Mg-, Ca-, Na-, and Kmontmorillonite, an ideal solid mixture of Mg-, Ca-, Na-, and K-saponite, and an ideal solid mixture of 14A-clinochlore and 14A-daphnite. The solid phases precipitated during open-system dissolution of 1 mole of serpentinite (under PCOZof 10-2.5bar) are, in order of appearance, goethite, gibbsite, kaolinite, montmorillonite, saponite, sepiolite, calcite, and hydromagnesite. Chlorites do not form due to low availability of Al. The most important product minerals are goethite and sepiolite, which start to precipitate at log(<) of -6.02 and -1.97, respectively. Gibbsite is transformed in kaolinite, kaolinite is converted to montmorillonite, and montmorillonite reacts producing saponite, which appears at log(<) of -2.19, and is the only stable product of this reaction sequence. Calcite and hydromagnesite are produced for log(<) > -1.5. The most important product phases generated through closed-system dissolution of 1 mole of serpentinite are serpentine and goethite. Calcite, sepiolite, and saponite are produced in smaller, similar amounts. Hydromagnesite has ephemeral existence. The solid mixtures of montmorillonites and saponites are usually dominated by the Mgendmembers.
4.3 The role of calcite precipitation during closedsystem dissolution of serpentinite
4.1 The external constraints In a first step, the local meteoric water (Marini et al., 2000) was reacted with serpentinite at temperature of 12"C, PCOZof 10-3.5bar (10-2.5bar in a separate run), under open-system conditions, and PO*of 1041.5 bar, to reproduce the average T-Pcoz-redox conditions prevailing at shallow depths, where MgHC03 waters circulate.
Calcite precipitation starts at log(5) of -4.14, but initially it has no effect on Ca and Ctotmolalities, since the moles of precipitated calcite, ncalcite, are much less than mCa and mctot(Figure 1). At log(<) of -2.8, ncalc;teequals mca and, therefore, mca decreases significantly for -3
Figure 1. Computed moles of calcite precipitated during dissolution of serpentinite, in a system closed to exchange of COz gas, and related changes in Ca and C,,, molalities.
Figure 2. Plot of Ctot vs. pH showing both analytical data and theoretical trends for open- and closed-system dissolution of serpentinite. Symbols are as follows: diamonds = Mg-HC03 waters, triangles = high-pH, Ca-OH waters.
dissolution of serpentinite, mca reaches a sort of plateau due to a balance between inputs of Ca (serpentinite dissolution, in spite of its relatively low CaO content) and outputs (calcite precipitation). At log({) of -0.52, ncalcite equals mctor. Further serpentinite dissolution, coupled with calcite precipitation, causes a remarkable decrease in mctot, since C is not supplied to the system. This decrement in the total molality of carbonate species (mainly co32-) is accompanied by a concurrent increase in mca, as the aqueous solution is in equilibrium with calcite.
log (aHco3-) - pW space, thus explaining the slope of theoretical reaction paths. This slope of +l holds true for waters interacting, under open CO2 conditions and at constant Pc02 and temperature, with any lithotype and not only with calcite and dolomite, as already observed. Most Mg-HC03 waters plot between the two theoretical paths in Figure 2. During serpentinite dissolution under closed CO2 conditions, an abrupt decrease in theoretical mctotis determined by calcite precipitation, for log(<)>-0.52 (pH > 10.2), at 12OC, and for log(<)>-1.50 (PH > 9.48), at 25°C. Analytical pH and mctotof high-pH waters plot along the theoretical trends. Calcium. During serpentinite dissolution under closed CO2 conditions, theoretical rnca and pH increase, until attainment of calcite saturation (Figure 3). A decrease in Ca concentration, concurrent with the increase in C,,, concentration (see above), takes place upon further dissolution of serpentinite, under saturation with calcite. Most MgHC03 waters plot between the two theoretical trends or somewhat above them. During serpentinite dissolution under closed CO2 conditions, calcite precipitation brings about an increment in theoretical Ca concentration for log(<) > -0.52 (pH > 10.2), at 12"C, and for log(5) > -1.50 (pH > 9.48), at 25°C. Analytical Ca concentrations and pH values of high-pH waters are consistent with theoretical data. Magnesium. Under open-system serpentinite dissolution, the logarithm of Mg concentration is expected to increase with pH following linear trends of slope +l, at least for pH > 7.44, i.e, when HC03- is the dominant anion in the aqueous
4.4 The log(concentration) vs. pHplots for C,,, Mg, and Ca
Total carbonate. Two theoretical reaction paths for open-system serpentinite dissolution, under Pc02 of 10-2.5and 10-3.5bar, are drawn in the C,,, vs. pH plot (Figure 2). At pH > 7.44 (i.e., one log unit above the -pKH2C03 at 12"C), where mctotis practically equal to r n H C 0 3 - , both log(mctot)-pH functions approach straight lines of slope +l. The slope of these linear relationships can be justified referring to the heterogeneous equilibrium:
-
whose equilibrium constant can be written as:
-
When both temperature and Pc02 are kept constant, equation (4) defines a straight line of slope +1 in the 147
5 CONCLUSIONS
Figure 3. Plots of Ca vs. pH showing both analytical data and theoretical trends for open- and closed-system dissolution of serpentinite. Symbols as in Figure 2.
solution. The reason for this is that the concentration of HC03 (in equivalent unit) is balanced by a corresponding concentration of Mg. Again, most Mg-HC03 waters plot between the two theoretical reaction paths.
The irreversible water-rock mass exchanges during the progressive evolution from rainwaters to neutral, Mg-HC03 waters and to mature, high-pH, Ca-OH waters was simulated through reaction path modeling, involving bulk dissolution of a local serpentinite. The computed concentrations of Ca, Mg, Ctot,and SiO2 (not shown due to space reasons), and pH values, are fully comparable with analytical data, indicating that the computed irreversible water-rock mass transfer is a realistic simulation of the natural process. During serpentinite dissolution in the absence of C sources, calcite precipitation depletes the aqueous solution in C species, bringing about the concurrent increment in Ca concentrations and the evolution of water composition towards a Ca-OH facies. As soon as Ca-OH waters reach the surface, they absorb atmospheric CO2, because of their high pH values, triggering calcite deposition. The occurrence of this process suggests that these high-pH waters might be used in the sequestration of industrial CO2 for preventing environmental impact to the atmosphere.
REFERENCES
Figure 4. Plots of Mg vs. pH showing both analytical data and theoretical trends for open- and closed-system dissolution of serpentinite. Symbols as in Figure 2.
Barnes I, O'Neil J.R. & J.J. Trescases 1978. Present-day serpentinization in New Caledonia, Oman and Yugoslavia. Geochim. Cosmochim. Acta, 42: 144-145. Marini L., Canepa M., Cipolli F., Ottonello G. & M. Vetuschi Zuccolini 2000. Use of stream sediment chemistry to predict trace element chemistry of groundwater. A case study from the Bisagno valley (Genoa, Italy). J. Hydrol., 24 1, 194-220. Pfeifer H.-R. 1977. A model for fluids in metamorphosed ultramafic rocks. Observations at surface and subsurface conditions (high pH spring waters). Schweiz. Mineral. Petrogr. Mitt., 57: 361-396. Wolery T. & Daveler S.A. 1992. EQ6, A computerprogramfor reaction path modeling of aqueous geochemical systems: Theoretical manual, user's guide, and related documentation (version 7.0). Report UCRL-MA- 1 10662 PT IV. Lawrence Livermore National Laboratory.
In the dissolution of serpentinite under closed CO2 conditions, theoretical Mg concentration does not change significantly until equilibrium with serpentine is attained for log(<) of -2.03 (pH of 9.48), at 12"C, and for log(<) of -2.77 (pH of 9.47), at 25°C. Serpentine precipitation causes instead a remarkable decrease in Mg concentration. Analytical Mg contents of high-pH waters are in substantial agreement with theoretical values.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Does plagioclase control the composition of groundwater in the crystalline basement? Kurt Bucher institute of Mineralogy, Albertstr. 23b, 0-79104 Freiburg, Germany
Ingrid Stober Geological Survey Albertstr. 5, D-79104 Freiburg, Germany
ABSTRACT: Two chemically distinct types of groundwater are present are the fracture porosity of the Black Forest crystalline basement (granites and gneisses). The first type is a C02-rich Ca-HCO, water which is used by the extensive mineral water industry of the Black Forest area. The second type is a Na-C1 thermal water often used in thermal spas, e.g. Baden-Baden. Composition data of waters from the Black Forest area suggest that the following secondary phases are formed during the water-rock reaction in the basement: quartz, kaolinite, chlorite, smectite, calcite, dolomite and most importantly stilbite (and other Ca-zeolites). Groundwater composition from crystalline basement indicates that the dissolution of the major minerals of basement rocks, feldspar and mica, dominate the chemical evolution of water in the granitoid rock matrix and it is the major source of cations. One prominent feature of many basement waters is their relatively low pH (6-7). This suggests that silicate mineral hydrolysis cannot explain the high solute concentration alone. Zeolite-formation consumes water from an initially low-TDS fluid leaving behind a residual brine. The process does not affect pH. Depending on temperature, laumontite, heulandite and stilbite are the major Ca-zeolites formed by the plagioclase-water reaction. Zeolites have been reported from fractures and clefts in granites from a large number of localities including deep research bore holes and Alpine fissure clefts.
1 INTRODUCTION
Water in granite and gneiss of the upper continental crust shows a limited number of characteristic composition patterns. Two distinct patterns (Figure 1) for moderatly mineralized groundwater of the Black Forest basement have been reported by Stober and Bucher (1999a). The common rock-forming minerals of granites, feldspar and biotite, are possible prime sources for the major cations (Na, Ca, Mg, K). The origin of the anions is less obvious. The major cations in deep groudwater, Ca and Na, are main constituents of plagioclase feldspar, the most abundant mineral in the continental crust. How does plagioclase control the composition of groundwater?
present study, we recalculated the data 111 terms of component activities by using the code PHREEQE (Parkhurst et al., 1980) and displayed the data on activity diagrams using the data of Berman ( 1988) for the solids and Helgeson et al. ( 198 1 ) for the aqueous species. Zeolite data were provided by de Capitani (pers. comm.). 3 STABILITY RELATIONSHIPS
The Ca-Na relationships are shown on Figure 2. A distinct separation of thermal and mineral waters appears on the diagram and the waters define a continuous trend, there is no offset between the groups or within the groups. The data can be fitted to one continuous linear trend, however, they also can be segmented into two linear trends with different slopes for thermal and mineral waters, respectively. At 25"C, plagioclase dissolution i n mineral water forms kaolinite and all thermal waters are consistent with stilbite formation. At 5 0 ° C a Ca-smectite field appears on Figure 2b, however, thermal waters will produce stilbite from plagioclase dissolution.
2 WATER COMPOSITION DATA The composition of the water in the basemerit of the Black Forest area has been described by Stober & Bucher (1999a) and Bucher & Stober (2000) and its major element load has been interpreted in terms of conventional hydrolysis reactions. For the
149
Figure 1. Scholler diagrams showing typical composition patterns of mineral waters (left) and thermal waters (right) of the Black Forest crystalline basement. Many waters are oversaturted with respect to quartz and all waters with SIqtz>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 << 1:
Figure 1. Contours in MPa of force-of-crystallization stress generated by A replacing B, for several mineral pairs. For k$kA>>l as in “quartz in calcite”, the force-of-xln and growth rate are smallest (eqs. 10,8a). For k$k,<
Thus, quartz or chert concretions growing in limestone are predicted not to fracture or deform the host limestone, but to replace it. 4 GROWTH-DRIVEN VEINS
Under the maximum normal stress difference, mechanisms of accommodation of growth of crystal A other than dissolution may take over or dominate - such as fracturing or deformation of the host rock. On the other hand, under the minimum stress, the growth of A becomes comparable to unconstrained crystal growth in a fluid medium: growth-driven fracturing and/or deformation of the host rock are now unlikely. The common case of replacement between quartz and calcite reflects well the asymmetry inherent in equations (9,lO); see Figure 1. For rate constants for quartz and calcite of 4 x lO-*l and 10‘l6 s/cm, respectively (Fletcher & Merino, 2000), it is clear that where quartz (= mineral A) replaces calcite (= mineral B), kAkB << 1 and the force of crystallization is minimal, Eq (10). But where calcite replaces quartz, kA/kB >> 1 and the force of crystallization is greatest, Eq (9). For a small supersaturation of, say, l&, = 5 with respect to calcite, Eq (9) together with ApA= RT~~SZA gives a normal stress difference of 1040 bars (104 MFa) for the replacement of quartz by calcite at 25OC. For the same supersaturation with respect to quartz, Eq (10) yields a minute force of crystallization of order lO-’ bars (106 MFa) during the replacement of calcite by quartz.
As mentioned in the Introduction, another potential consequence of crystal growth in rocks is fracturing, which has a crucial role in the feedbacks described in Section 2 to account for earthquake triggering and ore-body self-localization. But a less dramatic consequence of the fracturing that results from crack propagation driven by stresses around dispersed growing crystals or crystal aggregates, or from mineral growth starting in small preexisting fractures, is the focusing of mineral precipitation in the form of veins. Following Taber (1917) we suppose that at least some vein opening is driven directly by growth: the crystal growth forces the host rock aside as it takes place. These veins are the ones we consider here. (Whether a vein passively fills a previously open fracture or is accommodated by the host rock creep driven by growth, may have a lot to do with the crystalline texture, equant or fibrous, of the vein mineral.) To model growth-driven veins we imagine each one to be penny-shaped (radius c, thickness w) and to be embedded in a spherical representative volume element (RVE) of radius by itself embedded in a large rock sphere of radius R (the “mineralized zone,” M Z ) which expands by viscous creep to accommodate the formation of the vein. The normal stresses outside and inside this large sphere are 00 163
and (TR, respectively. The normal stress just inside the RVE, radius b, is o n . See Fig. 2. We apply here, with some modification, Walder and Hallet’s (1985) treatment of the growth of an ice lens in a rock driven by a difference in the chemical potential of pore water. We use the following relations: a) the rate of opening, w, of the crack is governed by this version of Eq. (7) dw/dt = 2k*[ApA - (00
-
on)VoA]
(11)
b) the accommodation of the vein by viscous expansion of the MZ, viscosity q, is
d(c2w)/dt = (9b3/8q)(oo- OK)
(12)
c) We suppose that the aspect ratio w/c remains constant and that crack growth requires a constant normal stress difference S:
Figure 2. Model system for veins in a septarian concretion. Penny-shaped veins with aperture w and radius c grow within representative volume elements, RVE, of radius b, themselves packed in a mineralized zone, M Z , of radius R.
Writing 00 - = (00 - 0,) - (GR- 0,) = (00- 0,) S, and eliminating CTO - OK, we solve for 00 - on, obtaining 00 - on= ( 1 + M)-’[S + MRT(I~Q)/VO~]
REFERENCES ( 15)
Substituting (15) into the kinetic relation, (1 l), the rate of vein growth becomes dw/dt = (1 + M)-’2kA[RTlnQ- SVoA]
(17)
This equation summarizes the dependence of growth rate on stress, supersaturation, rock viscosity, kinetic rate constant, and vein dimensions. For a quartz vein as an example, taking S = 5 MPa, T = 373K, VoqUam = (60/2.65) cm3/mole, and kA = 4 x 10‘21s/cm,one obtains a growth rate of only 0.005 cmka. The factor M may be estimated for septarian concretions, which range in size from 10 cin to several meters and contain internal arrays of calcite veins. See Fig. 2. For a meter-size septaria with veins having radii c = 5 cm separated by distances of, say, b x 10 cm, and using the rate constant kA = 1.19 x 10-l6 s/cm for calcite, and 6 = 0.1 to reflect the relatively small aspect ratios of veins in this setting, we obtain
M x 3 x 1O-l8q (with the viscosity in Pas)
supersaturation goes into mechanical work rather than growth kinetics.
(18)
For M to be significant in Eq. 15, the effective viscosity of the sediment would have to be q = 1OI8 Pas or more. This would lead to a reduction in growth rate of about (l+M)-‘ = ?4that in the mechanically unconstrained value (Eq. 17). This example, if typical, suggests that much of the “locally available” 164
Bruton, C.J. & H.C. Helgeson 1983. Calculation of the chemical and thermodynamic consequences of differences between fluid and geostatic pressure in hydrothermal systems. Amer. J. Science 283A: 540-588. Dahlen, F. A. 1992. Metamorphism of nonhydrostatically stressed rocks. Amer. J. Science 292: 184-198. Fletcher, R.C. & E. Merino 1997. Rheology and kinetics of replacement. In 7‘h Annual V.M. Goldschmidt Conference, Lunar & Planetary Institute Contribution No. 921, p. 72. Fletcher, R.C. & E. Merino 2000. Mineral growth in rocks: Kinetics and rheology, replacement and veins. Geochim. Cosmochim. Acta, submitted. Kamb, W.B. 1959. Theory of preferred crystal orientation developed by crystallization under stress. J. Geology 67: 153170. Maliva, R.G. & R. Siever 1988. Diagenetic replacement controlled by force of crystallization. Geology 16: 688-691. Merino, E., Nahon. D. & Y. Wang 1993. Kinetics and mass transfer of pseudomorphic replacement: Application to replacement of parent minerals , . _ by AI, Fe, and Mn oxides during weathering. Amer. J. Science 293: 135-155. Merino, E. & T. Dewers 1998. Implications of replacement for reaction-transport modeling. J. Hydrol. 209: 137-146. Misch, P. 1971. Porphyroblasts and ‘crystallization force’: Some textural criteria. Geol. Soc. Amer. Bull 82: 245-25 1. Nahon, D. & E. Merino 1997. Pseudomorphic replacement in tropical weathering: . . .kinetic-rheological origin. Amer. J. Science 297: 393-417. Steefel, C & A. Lasaga 1994. A coupled model for transport of multiple chemical species and precipitatioddissolution reactions. Amer. J. Sci. 294:529-592. Taber, S. 1917. Pressure phenomena accompanying the growth of crystals. Proc, Nat. Acad. Sciences 3: 297-302. Walder, J. & B. Hallet 1985. A theoretical model of the fracture of rock during freezing. Geol. Soc. Amer. Bull. 96: 336346.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Large-scale hydrothermal dolomitization in the Southern Cantabrian Zone (NWSpain) M.Gaspanirri & T.Bechstadt Geologisches Institut, INF 234, Universitat Heidelberg, Germany
M .Boni Dipartimento di Scienze della Terra, Universitci di Napoli, Italy
ABSTRACT: Epigenetic replacive and cavity filling saddle dolomites, frequently forming zebra structures, are widespread in the Cantabrian Zone of NW Spain, especially in Carboniferous carbonates. Over 250 rock samples of both dolomite and parent limestone have been collected in a region of approximately 500 km’,along 12 different sections through Carboniferous rocks. Petrography,CL, stable 0-and C- isotopes and chemical analysis have been carried out so far. The dolomitization consists of subsequent replacive and void filling stages and is associated with an important increase of porosity and significant Fe enrichment. Later cement generations (commonly calcite) are found as well. The isotope data point to a rock-dominated fluid flow reaching burial temperatures at the time of dolomitization.Extensional phases during the late Variscan compression as well as the Permian post-orogenic extension could be invoked as a probable time of dolomitization. A later (Jurassic?) age cannot be ruled out, however. 1 INTRODUCTION
covered by Devonian rocks with stratigraphic gaps. This succession is overlain by a thick pile (up to 2800 m) of Carboniferous carbonates and some clastics.
In the Cantabrian Zone, the foreland of the Hercynian belt in the Iberian Peninsula, extensive occurrences of epigenetic burial dolomite, located in different thrust units (Somiedo-Correcillas, Sobia-Bod6nYPicos de Europa, Esla, and Ponga Units; Fig. 1) have been reported by different authors (GomCz-Fernhdez et d . 2000, Grimmer et al. 2000). This study concentrates on the Boddn Unit with the aim of investigating geochemistry, emplacement mechanisms, relative timing of dolomitization, as well as geodynamic implications.
Table 1. Main charcteristics of the dolomitized formations in the Bod6n tectonic Unit Formation
Age
Thickness
Lithotype
0-1000 m
Light grey massive bioclastic limestone
Westfalian A
Valdeteja Namurian B
Barcaliente
NamurianA
200-350 m
Alba
Dinantian
10-30 m
2 GEOLOGICAL SETTING The main structural features of the Cantabrian Zone are thin-skinned thrust-sheets with faults and folds, showing a vergence towards its central part (PerGzEstaiin et al., 1990). The Bod6n Unit comprises three allochthonous nappes. It is bound by the Leon Line (an E-W oriented strike-slip fault) in the North and by the Somiedo-CorrecillasUnit in the South. Nappe emplacement began in the Westfalian B with the thrust surfaces being mainly located in Cambrian sediments. The Bod6n unit consists of a relatively complete Cambrian to Silurian mainly siliciclastic succession,
Black (organicrich) micritic and laminated limestone Encrinitic pink and red nodular limestones
3 FIELD OBSERVATIONS The dolomitization mainly affects Lower Carboniferous limestones of the Alba, Barcaliente and Valdeteja Fms. gable 1). It affects only rarely the underlying Cambrian to Devonian sediments, while being absent in the Westfalian and Stefanian lithologies.
165
Figure 1. Geological sketch map of the Cantabrian Zone showing the different tectonic units (after Pertz-EstaGn et al. 1990).
Figure 2. C - 0 isotopic analysis plot of identified dolomites, calcite and undolomitized Carboniferous limestones.
166
The morphology of the dolomite bodies is generally irregular, often showing remnants of non dolomitized limestones. Limestone-dolomite contacts are very sharp and cut both stratification and sedimentary structures. They are neither clearly controlled by stratigraphy nor by tectonics. The dolomite often forms banded fabrics which resemble the ,,zebra structures" described from many other epigenetic dolomites (Fontbot6 et al. 1983, Arne et al. 1989, Nielsen et al. 1998, Boni et al. 2000). The banded fabrics have a bilateral symmetry, resulting from the repetition of mm-scale dark grey and light grey to white dolomite sheets and are spatially associated with pre-existing microfissures. The dolomitization is closely related with secondary porosity development,strongly different from the tight precursor limestones: cavities, linear to roundish in shape, are fairly widespread. In some cases, however, late cement generations, mainly represented by calcite,completely fill the cavities. The cavities are usually unconformable to the often strongly inclined bedding of the host rocks, which clearly underwent deformation previous to dolomitization. Although a constant orientation of the sheetlike cavities has not been recorded, they are often horizontally aligned where dolomitization affected well bedded or laminated sediments of the Alba and Barcaliente Fms., and randomly distributed in the replaced massive limestones of the Valdeteja Fm. This evidence suggests a control by stratification/ lamination / microfissures on the developmentof the cavities.
zoned. The most common type of zonation shows only one darker red and impurity-free rim, parallel to the growth surfaces. The calcite C shows a very bright orange and usually unzoned CL,whereas the unreplaced limestone has a slightly less bright orange one.
4.2 Stable isotopes The 6l80values are similar for all dolomite phases (A, B1, B2; Fig. 2) and show an ample spread (-3.5 to -1 1%0PDB). It mirrors the equally large field of the precursor limestones (+1 to -6.5%0 PDB) but are somewhat lighter. The values are consistent with a burial diagenetic field. In the same sample, a depletion in 6l80is observed from generation A to B2. The 613C values for dolomites (+2 to +5%0 PDB) are inherited from the Carboniferous limestones and are close to the marine values of this time (Grossmann, 1994). The late calcite phase has the most depleted 6l80 values (-8 to -16%0 PDB): it seems to have been deposited by hotter fluids. 4.3 Magnetic behaviour Magnetic susceptibility measurements at room temperature performed for the limestone showed a diamagnetic behaviour, whereas the first generation of dolomite A evidenced a paramagnetic behaviour. These changes in magnetic behaviour are possibly due to replacement of Mg2+by Fe2+in the crystalline lattice during dolomitization (Shogenova, 1999).
4 RESULTS 4.4 Microprobe analysis
4.1 Petrography and Cathodoluminescence (CL) More than 40 thin-sections were observed for petrographic characterization and staining was applied to many of them. Petrography of a well developed ,,zebra-dolomite" fabric shows (from outer to inner layers) medium crystalline, non-planar, replacive dolomite A, followed by medium to coarse crystalline, non-planar, inclusion-rich, replacive saddle dolomite B 1. These generations are postdated by coarse crystalline, non-planar, void-filling saddle dolomite B2; the latter is inclusion-rich in the core and inclusion-free at the outer rim. The last phase is a coarse crystalline, blocky calcite C, filling completely the cavities. Both dolomite A and B1 show a uniform dull red CL, are unzoned and characterized by numerous bright red-orange luminescent spots, probably due to the presence of mineral impurities derived from remnants of the preceding limestone. Dolomite B2 is characterized by the same dull red colour, with the exception of the most external part of the saddle scimitar-like crystals, which are always
The analyses (EDS) were performed along profiles perpendicular to the zebra-fabric. All dolomite phases are slightly calcium enriched (average value of 54.08% CaCO,); this is consistent with their saddle nature (Radke et al., 1980). A relatively constant iron content (1.14% FeCO,) is observed in dolomite A and B 1; it increases rapidly (2.88% FeCO,) in correspondence with the most external rim of the B2 crystals, having a darker CL,thus confirming the role of iron as quencher of luminescence. Iron contents in the parent limestone are significantly lower (0.27% FeCO,). Mn in the dolomites has a very irregular trend but its content is surprisingly high at the very dull luminescing rim of B2 crystals, also having the highest Fe content. This, apparently in contrast with the well established role of Mn as activator of CL,confirms the positive correlation between Mn and Fe and the not simple proportionality between intensity of CL and Mn concentrationin dolomite (Piersen, 1981).
167
5 DISCUSSION
REFEXENCES
A higher Fe content in dolomite relative to limestone due to replacement of Mg2+ by Fe2+ during dolomitization has been qualitatively established (staining, magnetic properties, CL),and confirmed by microprobe analysis. The dolomitizing fluid apparently migrated through the rock along discontinuities, controlled firstly by stratificatiodlaminationjoints, and then by small microfissures, which constrained the orientation of the sheet-like cavities (Wallace et al. 1994). Different stages of dolomitization can be recognised: replacementldissolution of precursor carbonates with formation of the dolomite A, gradual transition to the dolomite B 1 and precipitation of the cement B2. The increasing crystal size from A to B2 could be due to more space availability for crystal growth, whereas the inclusion-free outer rim of the B2 phase could point to a slowing down of the crystal precipitation. The 6l80 depletion from dolomite A to B2, belonging to the same zebra-fabric, points to a temperature increase of the dolomitizing fluid during its evolution, or to its contamination by a more depleted solution. The second hypothesis is quite unlikely, because no contamination can be observed, affecting the chemistry and the CL of the subsequent dolomite generations. Further information on crystallisation history, kinetics and the types of fluids is expected from future fluid inclusion and Raman studies. The close relation between the isotopic values of dolomites and parent limestone argues for a rockdominated dolomitizingfluid-flow.This was probably set into motion at a time of crustal thinning, increased geothermal gradient and extensionalregime. In the Picos de Europa Unit (Fig. 1) similar dolomites show a clear spatial relationship with faults, crosscutting the Variscan thrusts. This has lead GomCz-Fernindez et al. (2000) to conclude that dolomitization took place in the late Carboniferous-Permianor even later, possibly related with the Jurassic thermal event (compare also Meyer et al. 2000). Although in the studied unit we could not find the same clear relationships between dolomites and late tectonic structures, field and isotope similarities between the dolomites of the two areas let us assume a similar timing of dolomitization. However, even if an emplacement of the burial dolomitization events in Northern Spain during the Permian post-orogenic extensional phase seems to be more likely, a later emplacement cannot be ruled out at the moment.
168
Arne, D.C. & S.A. Kissin 1989. The significance of ,,diagenetic crystallization rhythmites" at the Nanisivik PbZn-Ag deposit, Baffin Island, Canada. Mineralium Deposita 24: 230-232. Boni, M., Parente G., Bechstaedt T., De Vivo B. & A. Iannace 2000. Hydrothermal dolomites in SW Sardinia (Italy): evidence for a widespread late-Variscan fluid flow event." Sedimentary Geology 131: 181-200. Fontbot6, L. & G.C. Amstutz 1983. Facies and sequence analysis of diagenetic crystallization rhythmites in stratabound Pb-Zn-@a-F-) deposits in the Triassic of Central and Southern Europe. In H.J. Schneider (ed):Mineral Deposits of the Alps and of the Alpine Epoch in Europe: 347-358. GBmez-Fernandez, F., Both R.A., Mangas, J. & A. Arribas 2000. Metallogenesis of Zn-Pb Carbonate-Hosted mineralization in the southwestern region of the Picos Cf: Europa (Central Northern Spain) province: geologic, fluid inclusion, and stable isotope studies. Economic geology 95: 19-40. Grimmer, J.Q.W, Bakker, R.J., &eh, S. & T. Bechstaedt 2000. Dolomitization and brecciation along fault zones in the Cantabrian mountains. Journal of geochemical Exploration 69-70: 153-158. Grossman, E.L. 1994. The carbon and oxygen isotope record during the evolution of Pangea: Carboniferous to Triassic. Geological Society of America, Special Paper 288: 207228. Meyer, M., Brockamp, Q., Clauer, N., Renk, A. & M. Zuther 2000. Further evidence of a Jurassic mineralizing event in central Europe: UAr dating in hydrothermal alteration and fluid inclusion systematics in wall rocks of the Wersteige 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 the Dinantian zebra dolomites south of the Brabant-Wales Massif, Belgium. Sedimentology, 45: 727743. Pkrez-Estatin, A. & F. Bastida 1990. Cantabrian Zone. In: PreMesozoic geology of Iberia, Dallmeyer R.D. & Martinez G.E. Eds., Springer-Verlag, Berlin: 55-66. Pierson, B.J. 1981. The control of cathodoluminescence in dolomite by iron and manganese. Sedimentology 28: 601610. Radke, B.M. & R.L. Mathis 1980. On the formation and occurrence of saddle dolomite. Journal of Sedimentary Petrology 50(4): 1149-1168. Shogenova, A. 1999. The influence of dolomitization on the magnetic properties of Lower Palaeozoic carbonate rocks in Estonia. In: Paleomagnetism and diagenesis in sediments, Geological Society London, 151: 167-180. Wallace, M.W., Both R.A., Morales-Ruano, S., Fenoll HachAli, P & T. Lees 1994. Zebra textures from carbonatehosted sulfide deposits: sheet cavity networks produced by fracture and solution enlargement. Economic Geology 89: 1183-1191.
Water-RockInteraction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Geochemical Modelling of Groundwater Quality Changes during Aquifer Storage and Recovery (ASR) in the dual porosity Chalk aquifer, England I.Gaus, P.Shand & 1.N.Gale British Geological Survey, Wallingford, Oxon, UK
J .Eastwood Wessex Water Services, Wessex House, Passage Street, Bristol, UK
ABSTRACT: Aquifer storage and recovery (ASR) is a technique whereby water injected into an aquifer during periods of surplus (winter) is recovered from the same borehole when demands are high (summer). ASR cycle-testing trials in a dual porosity Chalk aquifer have been successfully modelled for non-reactive components. This has allowed geochemical reaction modelling between injected water, native groundwater and the aquifer minerals. In order to quantify these reactions, a geochemical model was developed and calibrated using PHREEQC 2.0, which is able to incorporate the physical characteristics of an ASR-scheme (e.g. radial flow, diffusive mixing as a consequence of dual porosity). The geochemical processes controlling fluoride, which is high in the native groundwater, were studied in detail. Different mechanisms (e.g. limited mineral availability, reaction kinetics) causing fluoride concentrations above that expected from dual porosity mixing were investigated. Such modelling can be used as a tool to optimise ASR cycle-testing and operation. 1 INTRODUCTION
Aquifer storage and recovery is an artificial recharge technique where potable water is stored in the aquifer adjacent to boreholes that are used for both injection and abstraction. It is being investigated in the UK because of its intrinsic environmental acceptability. An ASR-scheme can only be economically viable if a high percentage of the water initially injected is abstracted at a later date. Physical mechanisms for loss of injected fresh water include dispersion (hydrodynamic mixing and molecular diffusion), lateral groundwater flow and density stratification. In some dual porosity media (e.g. the Chalk aquifer in the UK) molecular diffusion is the dominant mixing process. The quality of the recovered water is also determined by geochemical reactions between the injected and native groundwater and reactions with the aquifer material itself. These reactions can be both detrimental and beneficial to the quality of the recovered water. In this study, physical and geochemical modelling were used to investigate changes in the recovered water quality of an ASR-trial in the confined Chalk at Lytchett Minster in southern England. The trials were undertaken by Wessex Water plc. Recovered water from the trials contained fluoride concentrations exceeding those predicted solely from mixing with the native groundwater (Williams et al. 2000). Identification of the mechanisms causing the elevated fluoride concentrations in the recovered water
is key to predicting concentrations in future ASRcycles and therefore the viability of the ASRscheme. 2 PHYSICAL MODELLING OF ASR IN THE CHALK
A physical model to quantify the role of dual porosity effects in an ASR scheme was developed and calibrated using experimental data from the extensive cycle-testing programme in the Chalk aquifer at the Lytchett Minster site (Gaus et al. 2000). The modelling package used was SWIFT/486, which is capable of modelling radial flow and dual porosity media. The system modelled comprised a cylinder of fractured aquifer centered round a well. Initially, the aquifer (both fracture and matrix) was full of native groundwater. The responses of the model to a typical annual cycle are shown in Fig. I. The curve (Fig. 1A) indicates the relative component of the native water measured at the well during a typical cycle. During the injection phase (Fig. 1B) the concentration in the well is the same as the injected water concentration (in this case zero). If the well is left standing (Fig. 1C) after injection stops, the concentration in the well gradually increases due to the effect of diffusive exchange between the injected water in the fractures and the native water in the matrix. During the recovery phase (Fig. 1D) the concentration in the well increases as more native water from 169
Figure 1. Modelled response of a dual-porosity aquifer to a typical injection-stand-recovery cycle.
further within the aquifer is drawn into the well. This is the key phase of the cycle as predictions can be compared to measured data. Once recovery stops, the native water component decreases due to diffusive mixing with the partly freshened water in the matrix.
tween the injected and the native water. Cycles varied considerably in length of time of injection and recovery periods and rates. While the injected water was of potable quality, the native groundwater was not potable due to high concentrations of fluoride (3.6 mg 1-’) and iron (0.34 mg/l). The SWIFT modelling showed that the dual porosity nature of the Chalk resulted in significant diffusive mixing which rapidly led to a large component of the native water in the recovered water. Due to the large volume of fluoride stored in the native groundwater within the matrix, the fluoride concentration in the recovered water will only decrease slowly during subsequent cycles. Furthermore, it was discovered that the high fluoride concentrations could be a consequence of processes other than diffusive mixing in a dual porosity aquifer. Fluoride concentrations in the recovered water were higher than those expected from simple mixing alone and it was concluded that the additional fluoride (approximately 10%) was possibly caused by the dissolution of fluorite in the aquifer. However, the recovered water does not reach saturation with respect to fluorite. An indication of the excess fluoride in the recovered water is shown in Fig. 2 for selected cycles (cycles 1, 7, 8 and 9). Chloride (which shows conservative behaviour) was used to calculate the amount of mixing between injection and native water and is shown on the x-axis. The ratio between mixing based on fluoride to that based on chloride is plotted on the y-axis. If fluoride also behaved conservatively, the fitted line would be horizontal (intersecting the y-axis at 1) indicating that both mixing estimates are the same. The fact that this ratio is much higher than 1 at the beginning of each recovery cycle, and reduces during the recovery phase, indicates that processes other than simple mixing determine the fluoride concentration. One can observe that where the main component of the recovered water is injection water, the fluoride content is much higher than would be expected from simple mixing. With increasing percentage of native water in the mixture, excess fluoride due to geochemical reactions will influence the fluoride mixing ratio to a lesser degree because of the high fluoride concentration present in the native groundwater. Geochemical influences on fluoride concentrations are therefore most obvious during the early part of the recovery cycle, in mixtures dominated by injection water. 4 MODELLING FLUORIDE
3 FLUORIDE AT THE ASR SITE
A total of 9 injection and recovery cycles were carried out at the trial site in order to measure the water quality changes under different injection/abstraction regimes and to start to build up a “buffer zone” be-
A 1-dimensional geochemical model was developed to model the ASR-cycles (incorporating both dual porosity and radial flow). The geochemical model package chosen was PHREEQC 2 (Parkhurst & Appelo 1999). Physical parameters used and details of the modelled ASR-cycles are shown in Table 1. 170
Table 1. Physical parameters for the PHREEQC-2 model. Parameter
Value
Fracture dispersivity Matrix porosity Fracture porosity Diffusivity Injectiordrecovery rate ASR-cycle specifications
No. of ASR-cycles
10 m
0.3 0.0; 10- m /s 45m3/d per fracture 4 months injection, 2.6 months standing, 2.6 months recovery, 2.6 months standing 2
fracture surfaces, was confirmed in some samples. Furthermore, analysis of chalk pore waters indicated that fluoride concentrations in the matrix may be higher than in pumped groundwater. Based on these findings, three different chemical mechanisms were modelled, each one able to cause an increase in fluoride in the recovered water: 1) the fluoride concentration in the matrix is higher than in the fractures: the recovered water will reflect the difference in fluoride concentrations between the fracture water and the matrix water in the Chalk with no geochemical reactions occurring; 2) fluorite is available for dissolution in the matrix and in the fractures, but only a limited amount is present: the concentration of fluoride in the recovered water will be limited by the amount of fluorite in the sediment available for dissolution; 3) fluorite is available in sufficient quantities for the fluoride concentration to reach saturation but the reaction kinetics is slow: the rate of fluorite dissolution determines the fluoride concentration in the recovered water (for the dissolution kinetics, a simple reaction mechanism was assumed). Detaiis for the modelling runs and the geochemical parameters are listed in Table 2. Table 2. Description of PHREEQC 2 model runs for modelling fluoride concentration in the recovered water. Model Run
Description
Sim Mixing Behaviour of fluoride is conservative Recovered water calcite/fluorite saturated Saturation High Matr 1 Fluoride concentration in matrix 20% higher than in fractures High Matr 2 Fluoride concentration in matrix 40% higher than in fractures Rate of dissolution= k* (1 -(Ca'') (P)' /K) Kinetics 1 k = 10-l' Kinetics 2 Rate of dissolution= k* ( 1-(Ca'') (F-)' /K) k = 2" 1O-l' Rate of dissolution= k* (1-(Ca") (F-)2/K) Kinetics 3 k = 4*10'"' Lim AV 1 10 mg fluorite/ kg sediment can dissolve Lim AV 2 30 mg fluorhe/ kg sediment can dissolve 50 mg fluorite/kg sediment can dissolve Lim AV 3
Figure 2. Observed excess of apparent mixing based on fluoritekhlorite mixing ratio for observations during cycles 1 (upper), 7, 8 and 9 (lower) of the Chalk ASR-trial.
Flow modelling results were calibrated using the SWIFT physical model described previously (Gaus et al., 2000). Only if the modelled ASR-cycles were at the operational time scale (e.g. 1 year) could a good m,itch between the results of both models be obtained. The c:a!ibrated geochemical model was subsequently used to model the evolution of the fluoride concentrations in the recovered water. Mineralogical and hydrochemical investigations were carried out on chalk cores in order to determine potential sources of fluoride in the confined Chalk aquifer. The presence of fluorite, particularly along 171
-Differences in fluoride concentration between the native water in the matrix and the fractures leads to a constant ratio between fluoride and chloride mixing during the recovery phase (Fig. 3A). -Limited availability of fluorite leads to an excess of fluoride concentrations at the beginning o f the recovery cycle. However, this excess is limited to a ratio of maximum 1.5 (Fig. 3B) and is expected to decrease further in subsequent ASR-cycles because the fluorite will be dissolved and flushed out during recovery. -Kinetically determined fluoride concentrations show an excess in fluoride during the beginning of the recovery cycle, reducing to approximately 1 at the end of the recovery cycle. Depending on the rate constant, the initial relative mixing ratio fluoride/chloride can vary between 1 and >5 (Fig. 3C). When comparing these results with the data from the field trial (Fig. 2) it can be concluded that differences in fluoride concentrations in the native water between the matrix and the fiactures cannot explain the observed fluoride pattern during the different cycles from the field trial. Also, limited availability of fluorite for dissolution can only partly explain the observations because the field data do not indicate a decrease in slope of the fitted line in subsequent cycles. Therefore, it is concluded that the increase in fluoride is controlled by the dissolution kinetics of fluorite. If this is the case, a decrease in fluoride in subsequent cycles is unlikely to occur until available fluorite is exhausted and mechanism 2 becomes the predominant process. However, the mount of fluoride removed in solution during ASR-cycles is likely to be small compared with that present within the solid phases in the aquifer. Therefore, the viability of ASR under such conditions is limited unless additional measures are taken such as blending or removal of fluoride from solution. ACKNOWLEDGEMENTS
Figure 3. Modelled excess of apparent mixing based on fluoridekhloride mixing ratio assuming that the excess of fluoride is caused by: (A) a difference in fluoride concentration between the matrix and the fractures; (B) dissolution of fluorite while fluorite has a limited availability in the Chalk aquifer; (C) dissolution of fluorite conditioned by slow reaction kinetics,
5 DISCUSSION AND RESULTS
Fig. 3 indicates the modelling results for the three different cases. It shows the relative increase in fluoride with respect to chloride during the recovery in year 2 of two consecutive one year ASR-cycles. The following interpretations can be made.
This work has been funded jointly by a Foresight LrNK award and the UK Water Industry Research Ltd as a project called ASR-UK. Wessex Water Services are gratefully acknowledged for permission to use data from this trial site. This paper is published with the permission of the Director of the British Geological Survey.
REFERENCES Gaus I., Williams A.T. & P. Shand 2000. Physical and chemical modelling (SWIFT-PHREEQC) of British aquifers for aquifer storage and recovery purposes. BGS Technical Report WD/00/08. Parkhurst, D.L. & C.A.J. Appelo 1999. User’s guide to PHREEQC (version 2) - A computer program for speciation, reaction-path, l-D transport, and inverse geochemical calculations. U S . Geol. Survey Water Resources Inv. Rept. Williams A.T., Gaus I. & I.N. Gale 2000. The impacts of dual porosity aquifers on aquifer storage recovery (ASR) schemes. Proceedings IAH 2000. Cape Town.
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Wafer-Rock lnferacfion 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Molecular Dynamics Simulation of the Uranyl Ion Near Quartz Surfaces J.A .Greathouse Department of Chemistry, St. Lawrence University, Canton, New York, USA
G.Bemis Department of Chemistry, St. Lawrence University, Canton, New York, USA
R.T.Pabalan Centerfor Nuclear Waste Regulatory Analyses, San Antonio, Texas, USA
ABSTRACT: Molecular simulation techniques were used to study the structure and dynamics of the aqueous uranyl ion near the quartz (010) surface. The potential parameters used in the study for water, quartz, and other aqueous ions are based on the CVFF forcefield, but for the uranyl ion, uranium was assigned a charge of 1-2.5 e and each oxygen was assigned a charge of -0.25 e. These partial charges, along with the accompanying short-range potential parameters, produce uranyl-water and uranyl-carbonate complexes in aqueous solution with geometries that are in excellent agreement with available experimental data. Molecular dynamics simulations were performed in the constant (NVT) ensemble to investigate the uranyl ion's sorption characteristics near the quartz surface. The simulation supercell contains approximately 32 quartz unit cells and 300 water molecules. The adsorbed uranyl ion retains a five-fold solvation shell, with the carbonate ion readily replacing two water molecules. The 0-U-0 axis aligns perpendicular to a singly protonated surface, and simulations involving a partially deprotonated surface resulted in inner-sphere surface complexes.
1 INTRODUCTION
Sorption processes are potentially important mechanisms for retarding radionuclide migration from nuclear waste repositories to the accessible environment. The effectiveness of sorption processes and their dependence on the characteristics of mineral sorbents and the chemistry of radionuclide-bearing water typically are determined through sorption experiments. In this study, molecular simulation techniques were used to provide a molecular-level description of sorption phenomena that complement data from sorption experiments. 2 SIMULATION METHODS All simulations were performed using the Open Force Field modules of Cerius-2 (Molecular Simulations, Inc., San Diego, California). Potential parameters were based on the consistent-valence force field (CVFF) using the Lennard-Jones form of the shortrange nonbonded interactions. CVFF parameters were used without modification for water, and slight modifications were made in the partial charges for the carbonate ion. Potential parameters for the uranyl ion and quartz atoms were taken from previous studies (Gilbaud & Wipff 1996, Pabalan & Lupkowski 1997). Bond stretch and angle bending parameters were included for all relevant species, except that the
quartz surface was held rigid throughout the calculations. We used the Crystal Builder module of Cerius-2 to create a 32-unit cell slab of a-quartz and the Surface Builder module to create the (010) surface. The simulation supercell was created with x- and ydimensions of 19.68 and 21.65 A,respectively, corresponding to four unit cells of quartz in each direction. The depth of the quartz slab was approximately 8 A,or approximately 2 unit cells. The z-dimension of the supercell was fixed at 32 A to allow room for water molecules and aqueous species. Before beginning a molecular dynamics simulation, the local water structure around the ions and the quartz surface was optimized for at least 500 steps using the Minimizer module. After minimization, the uranyl ion had already formed its primary solvation shell. Molecular dynamics simulations were performed in the constant (NVT) ensemble with the NosC-Hoover thermostat (Allen & Tildesley 1987). Temperature was set to 300 K with a 0.1 ps relaxation time. A cutoff of 9 A was used for short-range interactions and Coulombic interactions were treated using the three-dimensional Ewald sum (Allen & Tildesley 1987) with a reciprocal-space cutoff of 0.5 A-'.Total simulation times varied from 50-200 ps with a timestep of 0.001 ps. Table 1 provides supercell compositions and parameters for the seven simulations performed in this study.
173
Table 1. Composition and parameters for molecular dynamics simulations. Simulation Number of number Surface composition Time / ps waters 1 305 224 71 300 2 Singly protonated 3 Singly protonated 100 293 4 Partially deprotonated 129 297 5 Partially deprotonated 100 296 Partially deprotonated 6 50 300 7 Partially deprotonated 109 300
Aaueous species 1 uo*?+
Initial U-surface distance I A
-
1 UOZ2+ 1UOF 1 uo22+ 1 u o a 2 + ,1 co32I u o Z 2 + 2, ~ 0 ~ ' -
-
8.0 6.1 2.7 4.0 4.0
Table 2. Results of molecular dynamics simulations. Simulation D,, / 10-" m2.s D,, / I o-" m 2 s (nearest third j" (middle t h number Uranyl solvation complexa U 4 distance / Ab 1 U02(H20jci2+ 2.49 15.6" 2 7.3 13.1 3 U02(H20)62' 2.50 6.3 15.2 4 LJO2(H20)62+ 2.5 1 6.2 13.3 5 U02(H20)s(0)2' 2.50 6.3 13.4 6 U02(Hz0)dC03)(0) 2.4 1 6.7 14.7 7 fJ0z(Hz0)dC03)(0)2 2.36 4.1 12.0 aUranyl complexes in simulations 5-7 include surface oxygen atoms, denoted by (0).Charges due to (0)are not included. bU-O distances were obtained from the first peak in U-0 rdfs. "Calculated for the third of water molecules nearest the quartz surface. 'Calculated for water molecules in the middle third. "Calculated for all water molecules.
3 RESULTS
water molecules are oriented such that their Czv rotational axes are bisected by the bipyramidal plane. Next, we tested our model of the quartz-water interface in simulation 2, which included 300 waters and a singly protonated quartz (010) surface. A snapshot of the simulation supercell during the equilibrated portion of the run is shown in Figure 2. The z-axis in this supercell was held constant at 32 A, which provides ample separation for water far from the surface to exhibit bulk-like properties. The 0-0rdf peaks are located at 2.79 and 4.67 A,which compare well with simulation results for pure water (data not shown). The second 0-0 peak at 4.67 A is not consistent with bulk water, and we believe that disruption in long-range order in water is due to the quartz surface. Nevertheless, we are confident that we have developed a workable, three-dimensional model of the quartz-solution interface.
Table 2 shows the results of the simulations. Simulations 1 and 2 were conducted in order to validate the nonbonded interaction parameters and model for the quartz-solution interface, and simulations 3-7 were conducted to investigate uranyl solvation and surface complexes near the quartz (010) surface. The quartz slab remained rigid during the simulations except for the surface H atoms, which were allowed to move relative to adjacent oxygen atoms. 3.1 Validation of interaction parameters and quartz model In order to test the Coulomb and van der Waals parameters for UO?+ in bulk solution, simulation 1 was carried out with 200 water molecules and 1 U0z2+ ion in a box with edge lengths of 18.2 A.The equilibrium U022+ solvation shell consisted of five water molecules in a bipyramidal geometry with uranyl oxygen atoms in the axial positions. In Figure 1, we see U-0 and U-H radial distribution functions (rdfs) averaged over the equilibrium portion of simulation 1. The U 4 rdf contains two peaks, one corresponding to the primary solvation shell at 2.49 A and a second broader peak at 4.85 A.This average U-water-0 distance of 2.49 A is in excellent agreement with the experimental value of 2.42 A, obtained from proton NMR and X-ray diffraction experiments (Aberg et al. 1983). The first sharp U-H peak occurs at 3.21 A,indicating that the solvating
Figure 1. U-O radial distribution function (solid line) and UI3 radial distribution function (dashed line) for U02*+ in bulk water (simulation 1).
174
For simulations 4-7, we created a partially deprotonated region of the quartz surface by moving six adjacent surface hydrogen atoms to nearby surface oxygen atoms. While a completely deprotonated quartz surface is unlikely in natural aqueous environments, areas of deprotonated oxygen atoins are possible. For simulation 4, the uranyl was again inserted approximately 6 8, from the quartz surface and, as with simulation 3, the partial charges of lower surface H and 0 atoms were adjusted in each case to maintain local electroneutrality. The uranyl ion again formed an outer-sphere complex with the surface, as seen in Figure 4, indicating that a region of deprotonated surface has little effect on uranyl surface complexation. Simulation 5 was carried out with the initial uranyl placement much closer to the surface (2.7 A), which resulted in an inner-sphere surface complex. In addition to the U-surface-0 interaction, two solvating waters also formed hydrogen bonds with protonated surface sites. In this case, a surface oxygen replaced a water molecule in the uranyl solvation shell, but the U-0 rdf peaks remained at 2.51 and 4.78 A,which are very close to rdf peaks for the outer-sphere complex. The effect of the carbonate ion on uranyl surface coinplexation was investigated in simulation 6-7. Again, the initial placement of the uranyl ion was critical in the type of surface complex observed. In simulation 6, the uranyl ion was placed approximately 8 8, from the surface, and the carbonate ion was placed near the uranyl ion. As this system reached equilibrium, the carbonate ion quickly replaced two solvating waters, forming the U02(H20)3(COj) complex, and the U-0 rdf peaks, which now include carbonate 0 atoms, appear at
Figure 2. Equilibrium snapshot in the (z,x) plane of water near a singly protonated quartz surface (simulation 2). Si atoms are light gray, 0 atoms are dark gray, and H atoms are white.
3.2 Uranyl-quartz surface complexes For simulation 3, the initial configuration consisted of a uranyl ion approximately 6 A from a singly protonated surface. Several surface protons 011 the lower quartz surface were removed, and the lower surface H and 0 charges were adjusted to maintain electroneutrality for the entire supercell. Specifically, the lower surface H atoms were assigned a charge of +0.30 while the lower (z = 0) surface 0 atoms were assigned a charge of -0.7725. This slight difference in charge of H and 0 atoms offsets the additional +2 charge created with the uranyl ion.
Figure 3. Orientation of the uranyl 0 - U 4 vector relative to the quartz surface normal (simulation 2). A value of 90 degrees indicates that U0: is perpendicular to the surface.
The equilibrated structure indicates that the uranyl ion maintains its five-fold water solvation shell and forms an outer-sphere surface complex with quartz. As seen in Figure 3, however, the uranyl ion niaintains an i n which the o-u-o axis is nearly perpendicular to the surface normal.
Figure 4. Equilibrium snapshot of uranyl solvation complex, U02(H20)b2+, near a partially deprotonated quartz surface (simulation 4). Only a portion of the quartz surface is shown for clarity. Shading for Si, 0 and H is identical to Figure 2 .
175
2.41 and 4.66 8,. However, this species never approached the quartz surface. For several siinulations afterwards, the initial placement of the uranyl ion was moved closer to the surface, and an initial uranyl-surface separation of 4 8, was required to obtain an inner-sphere surface complex (data not shown). For simulation 7, we introduced two carbonate ions, and the uranyl ion was placed 4 8, from the partially deprotonated surface. After a carbonate ion replaced two water molecules in the uranyl solvation shell, the U02(H20)3(C03) complex proceeded to form a bidentate, inner-sphere surface complex, U02(H20)2(C03)(02). The second carbonate ion never entered the primary solvation shell, as seen in an equilibrium snapshot (Fig. 5). One surface oxygen atom is always closer to the uranium atom than the other, and the U-0 rdf peak at 2.36 8, (minor peaks at 2.81 8, and 3.38 8,) clearly indicate that both surface oxygen a t o m should be considered to be in the primary solvation shell. Time evolution of the two U-surface-0 distances are plotted in Figure 6. The surface oxygen a t o m are held fixed throughout the simulation, so the narrow range of these distances is evidence of the limited motion of the uranyl ion.
Additionally, D,, values of approximately 15 x IO-"' in2 s-', which is identical to that of bulk CVFF water, indicate that water in the middle region displays similar transport properties to bulk water.
Figure 6. Time evolution of U-surface 0 distance for two surface 0 atoms involved in the uranyl inner-sphere surface coinplex (simulation 7).
ACKNOWLEDGMENTS This work was funded by the New York Science Education Program (G. Bemis) and the U S . Nuclear Regulatory Commission (NRC) under Contract Number NRC-02-97-009 (J.Greathouse and RPabalan). This paper does not necessarily reflect the views or regulatory position of the NRC.
REFERENCES Aberg, M., D. Ferri, J. Glaser & I. Grenthe 1983. Structure ofthe dioxyuranium(V1) ion in aqueous solution-An x-ray diffraction and ' H NMR study. Inorg. Chern. 22:2986. Allen. M.P. & D.J. Tildesley 1987. Conzpirter Simulutioi7 ofLiqziids. Oxford, UK: Clarendon Press. Gilbaud, 1.' & G. Wipff 1996. Force field representation of thc UO~" cation from free energy MD simulations iii water. Tests on its 18-crown-6 and NO3- adducts, and on its calyx[6]arene6- and CMPO complexes. <J. Mol. Struct. (THEOCHEW 36635. Pabalan, R.T. & M. Lupkowski 1997. Molecular dynamic.^ Sinidation of Radioniiclide Soiption on A4inei-a1 Suifuces. SWRI-IRD-20-9875. San Antonio, Texas: Southwest Research Institute.
Figure 5 . Equilibrium snapshot of uranyl solvation complex, U02(H20)2(C03)(0)2(large spheres), near a partially deprotonated quartz surface (simulation 7). The quartz surface and nonsolvating species are indicated in stick format. The two large spheres on the surface are deprotonated surface oxygen atoms in the uranyl solvation shell. Shading for Si, 0 and H is identical to Figure 2.
Table 2 also includes calculated water selfdiffusion coefficients, D,,(Allen & Tildesley 1987), in two distinct regions near the quartz surface. Water molecules within a distance of 7.8 8, from the (0 10) surface are considered to lie within the nearest third, while water molecules between 7.8 and 15.6 8, constitute the middle third. In each simulation, D,,, for water in the middle third of the simulation supercell was nearly double that of water near the surface. 176
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Relative importance of physical and geochemical processes affecting solute distributions in a clay aquitard Glenn A.Harrington & Andrew J.Love Department for Water Resources, Adelaide, South Australia, Australia
Andrew L.Herczeg CSIRO Land and Water,Adelaide, South Australia, Australia
ABSTRACT: A solute transport and water-rock interaction model is developed to determine the relative importance of physical (advection), physico-chemical (diffusion) and geochemical (dissolutiodprecipitation and ion exchange) processes during the evolution of chemical profiles within a 40 m thick aquitard in the Tertiary Otway Basin of South Australia. Pore fluids exhibit relatively high concentrations of major cations, sulfate and chloride compared with concentrations in both the overlying and underlying regional aquifers, while bicarbonate and dissolved silica concentrations are lower. Modelling results indicate that diffusion and very low rates of advective transport can adequately explain the observed sodium, chloride and dissolved silica profiles. All other ionic profiles appear to have evolved more by water-rock interactions rather than advection or diffusion since the time of recharge, however uncertainties associated with estimating initial concentrations preclude quantitative mass transfer modelling. A 1-D advection-diffusion model was developed by Love et al. (1996) to simulate the observed chloride distribution in three vertical profiles collected from the aquitard. After setting boundary and initial conditions, the authors varied aquitard tortuosity and vertical porewater velocity at each site until a suitable match with the observed data was obtained. At the first site, diffusion was found to be the dominant solute transport mechanism. A combination of advection and diffusion was found to be controlling the solute distributions at two other sites located further down the regional hydraulic gradient. Concentrations of major cations and chloride in the aquitard pore fluids are much higher than in the aquifers, and are therefore decreasing with time due to diffusion of salts out of the aquitard. Conversely, bicarbonate and dissolved silica concentrations in the aquitard are lower than in the aquifers, thereby creating a potential for diffusion of these ions from the aquifers into the aquitard. This paper investigates whether or not the existing advection-diffusion model (Love et al. 1996) can adequately explain the vertical distributions of cations and anions (other than Cl) at Site 1, where diffusive fluxes dominate over advective transport. The importance of waterrock interactions for controlling solute distributions in the aquitard is currently not known.
1 INTRODUCTION Thick, low-permeability geological formations have become a target for increased research over the last decade due to their potential as repositories for toxic or radioactive waste (Nordstrom et al. 1989), for reconstructing past climatic or geologic conditions (Hendry & Wassenaar 1999) and for identifying mechanisms of inter-aquifer mixing (Brown et al. 2000). However, very little is known about the processes controlling the distribution of dissolved ions in such formations. The only recent work of this kind is that of Hendry & Wassenaar (2000) which reported the complex evolution of major ion profiles in glacial tills of southern Canada. The Otway Basin in South Australia contains two regionally extensive Tertiary aquifer systems that are separated by a thick (10 - 50 m) clay and mar1 aquitard. The aquitard is responsible for the confined conditions currently observed in the lower Dilwyn Sand Aquifer (DSA), and serves as a semipermeable barrier to the transmission of water and solutes to or from the overlying Gambier Limestone Aquifer (GLA). The groundwater flow regime and origin of salts in the two aquifer systems are well established (Love et al. 1993), however further research is required to quantify rates of inter-aquifer leakage and to identify the mechanisms which control the migration of solutes through the aquitard.
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2 METHODS
3 RESULTS AND DISCUSSION
The US Geological Survey’s finite-difference groundwater flow model MODFLOW, coupled with the solute transport code MT3D (Zheng 1990) was selected for all model simulations due to the user friendly interface provided by Visual MODFLOW and its capability of handling advective-diffusive transport. Initially, the porewater chloride profile simulated by Love et al. (1996) had to be reproduced using the same input and calibration data. This included an “initial” chloride concentration (50,000 years ago) of 7,500 mg/L, aquitard tortuosity z of 0.01, and mean porewater velocity v of 2.7 x 10-* m.d-’. The porewater velocity was converted into an equivalent vertical hydraulic conductivity K,, of 3.1 x 10-7 m.d-* using a head difference of 1 m and an aquitard thickness of 38 m. Any deviations between observed and modelled Cl concentrations were calculated and applied to all observed ion concentrations. For example, suppose at a particular depth within the aquitard that the modelled C1 concentration is 10 YOlower than the observed value. In this case, observed Cl and other ionic concentrations would all have to be decreased by 10% (ie. the correction is such that modelled and observed C1 concentrations are equal at any given depth). Simulations of all corrected ionic profiles were then performed with the same hydraulic data used for the chloride calibration. The most difficult component of this modelling exercise is to estimate the “initial concentrations”, ie. the concentrations of major ions in the water which recharged the aquitard. The stable isotope composition of pore water extracted from the aquitard is indicative of a meteoric origin (Love et al. 1996). Therefore, a shallow groundwater sample from the unconfined aquifer, which is also known to have a meteoric origin, was used as the basis for calculating initial concentrations. The major ion concentrations of the saline groundwater sample were then scaled up to a C1 concentration of 7,500 mg/L to match the initial C1 concentration determined by Love et al. (1996). However, the initial Ca and HCO3 concentrations in the aquitard would have been significantly different from those of the concentrated groundwater sample due to equilibration with the overlying Gambier Limestone Aquifer. To overcome this issue, the evaporated groundwater sample was equilibrated with calcite at 25OC using PHREEQC to provide equilibrium Ca and HC03 concentrations at the time of recharge to the aquitard.
Initial concentrations together with observed and modelled ionic profiles are presented in Figure 1 (a-h). The simulated C1 profile (Fig. l(a)) provides a reasonable match to the observed data and is in agreement with the result obtained by Love et al. (1996). Thus we can continue to use MODFLOW and MT3D for the simulation of other ionic distributions. With the exception of sodium, all other major ions show significant deviations between observed and modelled concentrations. These differences clearly demonstrate that waterrock interactions (in addition to dissolution of calcite prior to recharge) have had a strong influence on the present-day observed chemical profiles. Observed Ca, Mg and total S concentrations in porewater samples are generally higher than the advection-diffusion model predicts, which suggests addition of these species since the aquitard was recharged. Conversely, observed K and HC03 concentrations are lower than the modelled results, indicating possible removal of these ions. The aquitard consists predominantly of clay and glauconitic marl. Therefore, potential water-rock interactions may include cation exchange, silicate hydrolysis and carbonate precipitation or dissolution. It is most likely that the K has been adsorbed onto clays due to its high affinity for such minerals. The elevated Ca and S concentrations may have been derived through dissolution of gypsum within the soil zone during recharge of the aquifer. The initial chloride concentration of 7,500 mg/L determined by Love et al. (1996) suggests that the climate was more arid in the past (present day C1 concentrations in the unconfined aquifer rarely exceed 1,000 mg/L). Many parts of the Australian arid zone currently contain gypsum in the soil zone; hence this mechanism for the addition of Ca and SO4 is possible. However, on an equivalents basis, substantially more S appears to have been added compared to Ca, especially at a depth of 20-27 m where the Ca excess is relatively small. Furthermore, dissolution of gypsum does not account for excess Mg and removal of HCO3-. Another possibility derives from the assumption that the aquitard was deposited in a restricted marine basin where anoxic conditions prevailed with consequent precipitation of pyrite (FeS2). Subsequent penetration of oxygen (either by diffusion or recharge of oxygen bearing waters from the overlying aquifer) ma have oxidized the pyrite, which in turn releases H and dissolves Mg-calcite within the sediments.
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Figure 1 (a-h). Observed (e), corrected (0)and modelled (-) ionic profiles through the clay and mar1 aquitard at Site 1 in the Otway Basin, South Australia. Dotted lines (---)represent estimated initial concentrations in the aquitard 50,000 yrs before present, which were determined by evaporating a saline groundwater sample from the Gambier Limestone aquifer up to a C1 concentration of 7,500 mg/L and subsequently equilibrating the solution with calcite at 25OC using PHREEQC.
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International Atomic Energy Agency, Vienna. 73-86. Nordstrom, D.K., Olsson, T., Carlsson, L. & P. Fritz 1989. Introduction to the hydrogeochemical investigations within the International Stripa Project. Geochim. et Cosmochim. Acts 53: 1717-1726. Zheng, C. 1990. MT3D: A modular three-dimensional transport model for simulation of advection, dispersion and chemical reactions of contaminants in groundwater systems. Report to the US Environmental Protection Agency, Ada, OK, 170 pp.
This would yield high amounts of dissolved Mg and Ca, especially if there is subsequent re-precipitation of some CaC03. Similarly, there would be release of sulfate and buffering of pH by the carbonate mineral reactions. The apparent discrepancies between modelled and observed ionic profiles may also reflect uncertainties in our choice of initial concentrations. We have also tried using a C1 concentration of 7,500 mg/L and other ion concentrations determined from seawater iodC1 ratios; as well as an artificially concentrated rainfall sample from Kybybolite, a town located 50 lun east of the study site. Both of these scenarios produced similar results to those presented above, although the relative magnitude of “excess” alkaline earths was sensitive to the choice of initial conditions. Therefore, we believe that quantitative mass-transfer models for determining the origin and transport of solutes in aquitards is limited by the accuracy of the estimates of initial ionic composition of the recharge water.
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4 CONCLUSIONS Vertical distributions of dissolved C1 concentrations in a deep, low-permeability aquitard in southern Australia have previously been simulated using a simple advection-diffusion model. We have shown that most major cation distributions, as well as bicarbonate and total sulphur concentrations, cannot adequately be explained using this model, and hence water-rock interactions have played a significant role in their evolution. Unfortunately, it is not possible to quantify the effects of advectiondiffusion and water-rock interactions on the observed profiles due to large uncertainties in the choice of initial ion concentrations in the aquitard. REFERENCES Brown, K.G., Love, A.J. & G.A. Harrington (2000). Vertical groundwater recharge to the Dilwyn Confined Aquifer, South East, South Australia. Department for Water Resources SA, Report Book 2000-00044. Hendry, M.J. & L.I. Wassenaar 1999. Implications of the distribution of 6D in pore waters for groundwater flow and the timing of geologic events in a thick aquitard system. Water Resour. Res. 35(6): 1751-1760. Hendry, M.J. & L.I. Wassenaar 2000. Controls on the distribution of major ions in pore waters of a thick surficial aquitard. Water Resour. Res. 36(2): 503-5 13. Love, A.J., Herczeg, A.L., Armstrong, D., Stadter F. & E. Mazor 1993. Groundwater flow regime within the Gambier Embayment of the Otway Basin, Australia: evidence from hydraulics and hydrochemistry. J. Hydrol. 143: 297-338. Love, A.J., Herczeg, A.L. & G.R. Walker 1996. Transport of water and solutes across a regional aquitard inferred from porewater deuterium and chloride profiles, Otway Basin, Australia. Isotopes in water resources management. 1.
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Chemistry-transport coupled modelling of the Asp0 groundwater system (Sweden) since the last glaciation W.Kloppmann, D .Thikry, C .Kervkvan, A.Bourguignon, P.Nkgre1 & JCasanova BRGM, BP 6009, F-45060 Orlians Cedex 2 , France
ABSTRACT: Based on the specific coupled modelling approach developed at BRGM, a reaction-transport model of the Aspo groundwater system (Sweden) has been developed; it simulates a continuous scenario from 12.000 BP to the natural modern situation. Examples of the results (evolution of Na, "0, calcite) are presented. Na, Ca, C1 and stable isotopes are mainly related to mixing between the different mixing endmembers subsequently entering the system (glacial meltwater, Baltic Sea stages, meteoric water), calcite contents, pH and silica are determined by water rock interaction.
1 INTRODUCTION The long term prediction of the geochemical evolution of groundwater bodies within potential host formations is of particular concern for site assessment in the field of nuclear waste disposal (Boulton et al. in press). Evaluating the impact of changes in climatic/hydrological surface conditions on the groundwater quality has to make use of a combined approach of flow and transport modelling taking into account chemical reactions. An adapted methodology of chemistry-transport coupled modelling based on Specific Chemical Simulators (SCS) has been developed in the last few years (Kervevan et al. 1998). It was applied to the site of Aspo (Sweden) for which a detailed site investigation yielded huge quantities of geological, hydrogeological and geochemical data from the crystalline basement up to depths of >loo0 m (Laaksoharju et al., 1999a). The paleohydrological/hydrochemical evolution since the last glacial maximum is particularly complex due to the presence of deep seated brines, meltwater injection, several phases of immersion of Asp0 island by the subsequent stages of the Baltic Sea and a final emergence of the site allowing mixing with meteoric waters at the surface (Laaksoharju et al. 1999a). In the following, we describe, as an example for the application of the SCS approach, the attempt to reconstitute the geochemical and flow conditions at Aspo from the last deglaciation to the modern natural conditions prevailing before the construction of the underground laboratory.
Figure 1. Situation of the study site
2 FLOW AND TRANSPORT MODELLING The modelled zone is a square of 8 km length and extends to a depth of 1600 m below the present day sea level. All geometric (elevation) and hydraulic data (3 components of conductivity, hydraulic head, salinity) were provided by U. Svensson (1997). A compromise had to be found between computing time, spatial and temporal resolution of the transient state flow modelling and the complexity of the geochemical model. The final configuration adopted is a horizontal irregular grid of 26 x 26 cells and a vertical grid of 1 1 layers from 3 to 300 m of thickness. The hydrodynamic code used is the finite difference model MARTHE (Thikry, 1993a, b) taking into account an unsaturated zone, density effects, direc181
tional hydraulic conductivity (in fractures) and multi-component transfer integrating advection and mechanical dispersion. The initial concentrations in chemical species have been considered as constant in the whole domain depending only on whether a cell is considered as crossed by fractures or as composed of matrix only. The “ 0 signature is -8.9%0in the initial dense brine; -20.45%0 in the initial mixed brine and -32.0 %O in the initial glacial meltwater. The sea level, the concentrations and the stable isotope composition of the Baltic varied considerably with time. The boundary conditions were accordingly adjusted every 1000 year period. Salinity and hydraulic head calculations were performed 320 times for 1000 years with a frequency decreasing from the beginning to the end of each 1000 year stress period. The hydrochemical calculations were performed every 1000 years. A test with twice as much time steps yielded no significant difference for the concentrations of aquatic species.
and two cation exchange reactions (K and Na versus Ca). 3.2 Chemical boundary conditions All calculations are based on the reference waters defined by Laaksoharju et al. (1999b) as the groundwater samples closest to the hypothetic pure mixing endmembers. In order to characterise the geochemical system in terms of major species to be taken into account, we performed preliminary EQ 3/6 (Wollery, 1992) calculations of chemical speciation and saturation states of all the reference waters. The species whose relative concentration was less than 10-*with respect to the major species were not considered in the SCS. The different inputs of marine and continental waters into the system during the past 12,000 years were defined and taken into account in the coupled simulations. In order to finalize our selection of aqueous species actually taken into account in the S C S , EQ3/6 mixing calculations were also performed to simulate chemical compositiQns, speciation and saturation states of the different Baltic-Sea stages. The three predominant lithologies encountered in the four main boreholes and in the tunnel are the Aspo diorite, the Smgland granite and the Finegrained granite (Mazurek et al., 1995 ; Stanfors et al. 1999). Extensive studies characterise the mineralogy of the fracture fillings (Mazurek et al., 1995; Tullborg, 1989; Wikman et al., 1988 ; Landstrom et al., 1993). A selection of the occurring minerals, taking into account the constraints related to computing, was integrated into the SCS.
3 GEOCHEMICAL MODELLING
3.1 Development of the “PAGEPA ” Specific Chemical Simulator (SCS) The general approach chosen by BRGM is to design a dedicated geochemical code (so called “SCS”) developed specifically for each particular application that only takes into account the relevant processes to be considered (Kervkvan et al. 1998). One of the advantages of this approach is a significant improvement in computer efficiency that in turn allows relatively complex 3D coupled modelling. This approach thus requires us to develop a new SCS for each new application. The combined use of ALLAN 3.1 and the NEPTUNIX 4 software package automatically generates the coded part of the simulators, as long as the processes involved in the physical system being modelled can be described by a set of algebraic and/or ordinary differential equations. NEPTUNIX4 is a general solver which automatically generates a FORTRAN code solving the selected set of equations. ALLAN 3.1 functionalities include graphical interface, description language, translator for NEPTUNIX 4, graphical assembling of models (including the generation of coupling equations), exploitation of the simulators and postprocessing of the generated data. The SCS developed for the Aspo site includes 24 aqueous reactions of 12 elements @ C, 0, -ICa, , Al, Si, Mg, Na, C1, K, Fe, S,,), the ‘*O isotope, 7 precipitation/dissolution reactions (calcite, kaolinite, clinochlore, goethite, quartz, K-feldspar, anorthite)
3.3 Chemical evolution As a general observation we can state that certain parameters, the conservative species like C1, but also Na and bivalent cations, reflect essentially the mixing processes in the course of the subsequent stages of glacial/meteoric/marine waters injection whereas others (Si, pH, mineral precipitation) are closely related to the geometry of the reservoir and in particular of major fractures. Figure 2 shows, as example of the modelling results, the distribution of the Na concentrations of the 6”O and of calcite for 11,000, 6000 and 1000 years BP. All block diagrams show E-W sections cutting through Aspo Island. Na: The patterns of the Na distribution are, much like those of Ca, closely related to the chlorinity. The initial scenario at 12,000 BP represents the stratification of glacial meltwater over deep brines. The fluctuations between 12,000 and 10,000 BP can be attributed to density effects. The influence of Yoldia sea becomes visible at the 9000 BP stage by increasing salinity of the surface layers. Ancylus 182
lake waters (8000 BP) dilute the superficial strata and the Litorina stage increases again the the concentration in the surface waters. From 5000 BP onward, the progressive invasion of the system by low salinity meteoric waters from the West is driven by the increasing hydraulic gradient due to lowering of the sea level. The Baltic 2 stage introduces a second low Na plume from 1000 BP to present. Low chlorinities of the modern Baltic Sea (since 1000 BP) and of meteoric water determine the present day conditions in the surface layers. We observe a residual local salinity high near the Asp0 island where Baltic 1 (or Litorina sea) water is conserved between the intrusion of meteoric water and modern Baltic Sea water. Calcite: According to the defined initial conditions, calcite is limited to major fractures at 12,000 BP. Calcite precipitation in the fissure porosity of the matrix starts from the surface during the Baltic ice lake stage (Fig. 2). During the whole evolution of the site, we observe a rather continuous progress of a horizontal “calcite precipitation front” moving from the surface to depth. Obviously, calcite precipitation is fairly independent from the subsequent marine/meteoric entries into the system and from the major ion concentration distribution. From 3000 BP to present, calcite contents in the surface layers reach values comparable with the initial concentrations defined for the major fractures. Silica: Si concentrations do not vary significantly during the evolution of the site. This is only partly due to the fact that Si variations in the initial fluids entering the system are small. The crucial factor seems to be a buffering of the concentrations by mineral dissolution/precipitation (according to the chemical model, silica concentrations are controlled by quartz). pH: The total range of pH variations is relatively restrained (around 1 pH unit). The pH is definitely not an indicator of mixing but it’s distribution is closely related to mineral equilibria. The pH distribution is, much like silica concentrations and calcite contents, determined by the geometry of major fractures. In the fractures we observe the highest values around 9, which is in the uppermost part of the range of measured field values. 6I8O: Initially, the system is invaded by isotopically extremely depleted glacial meltwater, which mixes at depth with the more enriched brines (Fig. 2). The pattern of isotopic contents till 10,000 BP remains rather constant and is very similar to the sodiudchloride distribution (Fig 2). Both tracers can in fact be regarded as conservative under the conditions prevailing in the system. The impact of the I80-enriched Yoldia sea changes profoundly the isotopic contents of the system and the influence of the Yoldia water reaches depths of >900 m. At 8000 BP, the isotopic composition of the system is ho-
mogenised up to the bottom of the block where slightly enriched brine remains. The impact of isotopically enriched LitorinaSea (6”O = -5.5%0)becomes predominant from 7000 to 5000 BP and the isotopic composition of the Baltic Sea is constant up to present day conditions around -6%0. Meteoric water which is depleted (6l80 = -10.5%0) with respect to Litorina seawater progressively invades the system from the continent. Present day conditions are somewhat similar to chloride distribution but contrasts of the isotopic composition are much lower than for chlorinity.
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4 CONCLUSIONS
The coupled hydrodynamic-SCS modelling approach has been successfully applied to a complex setting of a fractured coastal aquifer influenced by glacial melt water, meteoric water and successive stages of brackish to saline seawater intrusions over 12,000 y The number of chemical reactions taken into account (in both space and time) are at present mainly limited by the constraints of computing capacity. Some simplifications of the natural system that were adopted in the present study, as neglecting redox processes though they seem to play a particular important role in this site (Banwart, 1995), could be overcome in a near future considering the fast evolution of computer capacities. Acknowledgements: The work was conducted under the PAGEPA project (F14W-CT96-0038), a European research project financed- through the EU 4‘hFramework Program. REFERENCES Banwart S. (ed.) 1995. The redox experiment in block scale. Final reporting of results from the three year project. SKB HRL Prog. Rep. (25-95-06}.Stockholm. Boulton G., Gustafson, Schelkes K., Casnova J. & L. Moren in press. Pakohydrology and Geoforecasting for Performance Assessment in Geosphere Repositories for Radioactive Waste Disposal (PAGEPA). Final Report. KervCvan, C., ThiCry, D. & P. Baranger 1998. SCS: Specific Chemical Siinulators dedicated to chemistry-transport coupled modelling Part I11 - Coupling of SCS with the hydrotransport modelling software MARTHE. Goldschrnidt Corgerence Toulouse Sept. 1998. Mineralogical Magazine 62A: 773-114. Laaksoharju M, Tullborg E.-L., Wikberg P., Wallin B. & J. Smellie 1999a. Hydrochemical conditions and evolution at the Asp0 HRL, Sweden. Appl. Geochern. 14: 835-859. Laaksoharju M., Skarman C. & E. Skarman 1999b. Multivariate mixing and mass balance (M3) calculations, a new tool for decoding hydrogeochemical information. Appl. Geochem., 14: 861-871. Landstrom O., Studsvik E. & E.L. Tullborg 1993. Results from a geochemical study of zone NE-I, based on samples from the Asp0 tunnel and drillcore KAS 16 (395 to 451M). SKB HRL Prog. Rep. (25-93-01}.Stockholm.
Figure 2. E-W section through the study area (lateral extension: 8 km x 8km, heith of the block: 1800 m): Distribution of solid calcite, dissolved Na and 6’’O at l 1,000, 6000and 1000 BP. Mazurek M., Bossart P.& T. Eliasson 1995. Classification and characterization of waterconducting features at Aspo: results of phase I investigations. SKB HRL Prog. Rep. (25-9503). Stockholm. Stanfors R., RhCn I., Tullborg E.-L. & P. Wikberg 1999. Overviev of the geological and hydrogeological conditions of the Asp0 hard rock laboratory site. Appl. Geochenz. 14: 8 19-834. Svensson, U. 1997. A regional analysis of groundwater flow and salinity distribution in the Asp0 area. SKB Technical Report 97-09. Stockholm. ThiCry D. 1993a. ModClisation des aquiferes complexes - Prise en compte de la zone non saturCe et de la salinitb. Calcul des intervalles de confiance. ii’ydrogkologie 4: 325-336. ThiCry, D. 1993b): Tridimensional and multi-layer modelling of transfers in unsaturated porous medium. GEOCONFINE Syniposiunz international G h l o g i e et confinement des cl&-
chets toxiques. Montpellier, juin 1993: 467-472. Rotterdam: Balkema. Tullborg, E-L. 1989. Fracture fillings in the drillcores KASOSKAS08 from Asp& Southeastern Sweden. SKB HRL Prog. Rep. (25-89-16).Stockholm. Wikman, H., Kornfalt K.A., Sgu L., Riad L., Munier R. & E.V. Tullborg 1988. Detailed investi,gationof the drillcores KAS 02, KAS 03 and KAS 04 on Asp0 island and KLX 01 at Laxemar. SKB HRL Prog. Rep. (25-88-1I ) . Stockholm Wollery T.J. 1992. EQ3/6, A Software Package for Geochemical of Aqueous Systems : Package overview and installation Guide (Version 7.0). Report UCRL-MA-110662-PT-I. Livermore, California: Lawrence Livermore National Laboratory.
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The Maqarin natural analogue study of a cement-buffered hyperalkaline groundwater plume: structural model and flow systems U .K .Mader, M .AdIer & V.Langer Rock- Water Interaction Group, University of Bern, Bern, Switzerland
P.Degnan, A .E .Milodowski & J .A.T. Smellie UK Nirex Ltd, Harwell, UK; British Geological Survey, Keyworth, UK; Conterra AB, Uppsala, Sweden
E.Salameh & H.N.Khoury University of Jordan, Amman, Jordan
L.Y.Griffault & L.Trotignon ANDRA, ChBtenay-Malabry, France; CEA, Cadarache, France
ABSTRACT: Active hyperalkaline groundwater systems at Maqarin (Jordan) are derived from a re-hydrated source rock of metamorphic origin formed by combustion of bituminous clay biomicrite. The effects of the hyperalkaline waters (pH to 12.8) on host sedimentary rocks provide natural analogues for the interaction of alkaline pore waters from degradation of underground cementitious repositories for the disposal of radioactive wastes. Outcrop, borehole and tunnel mapping indicate high-density fracture zones dominated by two sets with cross-cutting orientations that also control the movement of hyperalkaline waters. Hydraulic transmissivities determined by shut-in borehole testing range from IO-" to 10-l2m2/s for unaltered wall rock, cemented brecciated source rock, and conductive features, except for a few major inflow zones. Preliminary ages for rock alteration are around 100 ka.
1 INTRODUCTION The Maqarin area is located 25 km East of the Jordan Rift Valley, 16 km North of Irbid, and South of the Yarmouk River bordering Syria (Fig. 1). The former Maqarin Station along the abandoned Hijaz railway is located just west of the confluence of Wadi a1 Harir and Wadi as Shallala.
using repository concepts that include the use of cement and concrete in their design (Smellie et al. 2001). The Maqarin site has been under investigation for about 10 years by an international consortium in collaboration with the University of Jordan in Amman. This paper attempts to integrate field observations on the extent of the high pH plume with boundary conditions for fracture flow, hydraulic properties, and structural constraints. 2 GEOLOGICAL SETTING, METAMORPHIC AND HYPERALKALINE FEATURES
Figure 1. Location map of the Maqarin area. Elevation contour lines are shown on Jordanian territory only.
Active systems of very high-pH groundwaters at Maqarin represent a natural analogue par excellence for investigating the long-term effects of hyperalkaline pore waters from cement degradation on surrounding low permeability rocks. Such long-term effects are of importance in performance assessments for the geologicical disposal of radioactive waste
The deeply incised Yarmouk river valley exposes 300-400 m of Upper Cretaceous-Lower Tertiary carbonates overlain by Pliocene-Pleistocene basalts (3-5 Ma). The Amman Formation (not exposed; confined artesian regional aquifer) is overlain by the Bituminous Mar1 Formation (ca. 200m thick; bituminous clay biomicrite; aquiclude) and the Chalky Limestone Formation (regional aquifer), with strata generally dipping gently to the North (Khoury & Nassir 1982, Khoury et al. 1992, Smellie 1998, Alexander & Smellie 2000). Incision of the Yarmouk valley is constrained to be younger than the capping basalts (3-5 Ma) and older than the youngest valleyfilling basalts (0.35-0.48 Ma) near the mouth of the Yarmouk, found only a little above the present river level (Mor 1993). An undated basalt capping an ancient river terrace at 180 m.a.s.1. located immediately above the study area (Fig. 1, 'basalt plateau') is 185
thought to be older than the events leading to the production of hyperalkaline waters (Smellie 1998). This basalt may be related to the 0.8 Ma old ‘Yarmouk’ basalt dated by Mor (1993). Dischagging seepages of hyperalkaline waters occur at the Western Springs and Eastern Springs localities (characterized by travertines) and in dripping zones within several adits excavated to facilitate evaluation of the planned Jordan-Syria Unity Dam (Fig. 1). It is inferred that ingress of oxygen through fracture systems and spontaneous combustion of bituminous limestone created several bodies of high-temperature metamorphic rocks with relict mineral assemblages reminiscent of products from a clinker kiln. Subsequent re-hydration and recarbonatisation led to the formation of cement-like bodies. Leaching of such natural cements produces Ca-hydroxide (portlandite) buffered leachates guided along preferential flow-paths through unaltered limestone to discharge, or to inflow zones within the adits. An age of 107 ka on one hyperalkaline flow system (U-Th disequilibrium dating by ICP-MS, unpublished) might hint at a younger age for the Maqarin systems than the incision of the Yarmouk, in agreement with some earlier data (Alexander 1992). A deep adit located at 110 m.a.s.1. (A6; Fig. 1, Fig. 2) within the uppermost section of bituminous limestone allows access to two, possibly connected, re-hydrated / re-carbonated metamorphic bodies (‘cement zones’), enveloped by relatively sharp transition zones: 4m thick laterally and 0.7-2m thick at the base. The boundary at the top is not well constrained due to intense alteration, brecciation and interference with possible pre-existing weathering phenomena. The site is capped by a 70-80m thick stack of basalt flows with baked palaeo-soiland bedrock exposed in some places. A variety of hill slope deposits (scree, cemented and uncemented debris flows) and alluvial deposits record a complex history of valley incision and associated erosion. Largescale land slide events referred to by earlier work could not be confirmed by ground truthing and air
photo interpretation. However, there is extensive evidence that multiple land-slipping of the valley sides is an important process locally and that it is due to over-steepening of the valley sides andor tectonic activity. 3 STRUCTURAL FRAMEWORK Maqarin is located in a region of relatively complex tectonic activity, although apart from pervasive fracturing, the site itself is surprisingly unaffected by intense deformation. The regional tectonics are dominated by the sinistral movement of the Arabian plate relative to North Africa, with superimposed active extension all along the margin. Four major and discrete phases of deformation may have affected the area: (1) compressional deformation associated with transpressional stresses along strike-slip zones forming NE-SW trending open folds (Oligocene-Eocene ?), (2) pre-Miocene E-W extension related to the Red Sea rifting, (3) transverse fault movement due to the more rapid movement of the Arabian plate (< 20 Ma), and (4) possible local extension due to the transfer of strain off the Red SeaJordan River rift valley, marked by an E-W zone of volcanism and earthquake epicentres in southern Syria (< 5 Ma). Based on an interpretation of satellite imagery and quantitative fracture mapping, a sequence of structural elements was established and put in relationship to the metamorphic event and hyperalkaline systems. Beginning with the oldest: 0
0
The hingeline of a broad anticlinal structure trending SW-NE forms a structural high at the top of the bituminous limestone just east of the section depicted in Fig. 2 (may be related to the event (1) above). There are no obvious parasitic mesoscale structures related to this oldest element. A regional sub-vertical N-S trending fracture system with only minor or no associated dis-
Figure 2. Geological section along the trace of Adit A6 (Fig. 1). Numbers are referred to in the text. Y.R. = Yarmouk River.
186
other hand near-constant flow rates observed over several years and seasons within Adit A6, and the largely near-detection tritium contents of the hyperalkaline waters, are suggestive of regional rather than local recharge sources. A n additional complication is due to the perched nature of the groundwater table at shallow depths of Adit A6, and evidence from bore hole monitoring indicates that locally the unsaturated zone extends to the elevation of the adit. Inferred minimal path lengths of hyperalkaline fluid conducting fracture systems are in excess of 100 m, but possibly much longer both at Western Springs and the area of Adit A6. Diamond drilling to 6m depth in Adit A6 and subsequent hydrotesting was performed in 1999 and 2000 in an attempt to characterize and quantify the local scale hyperalkaline flow systems. The subhorizontal wells located in unmetamorphosed bituminous limestone (with or without hyperalkaline fluid conducting features) and in the re-hydrated metamorphic zone were fitted with hydraulic packers and pressure data loggers. Shut-in pressure pulse decay tests were initiated by a hand pump. Data evaluation was performed by type-curve matching of H/H" vs. log(t) plots (H: hydraulic head) using a simple radial and homogeneous model. Unaltered but jointed bituminous limestone has a transmissivity of 2-4.10-" m2/s. Measurements by a constant rate injection test (affected radius of 0.2-0.4m) yield transmissivities of almost two orders of magnitude larger for the same interval, suggesting that skin effects might affect the much shorter pulse tests. Bituminous limestone intersected by small hyperalkaline water conducting veins yield T-values of 5-10.10-'2 m2/s, and a section of brecciated and cemented metamorphic zone of 2.10-'* m2/s. Locally, much higher transmissivities do occur with rates of inflow exceeding 100 ml/min. Three major dripping zones are exposed in Adit A6 (Fig. 2, locations 2, 3, 4). Beyond the exposed cement zone in Adit A6 (location 5 ) the bituminous limestone is dry and, to date, no reliable formation water could be sampled. A new set of pressure-pulse data at a spatial resolution of 45 cm is presently being evaluated and does show marked differences between undisturbed, jointed and veined sections that can be correlated with a structural log based on core mapping and bore hole video logging. The portion of the metamorphic zone located within the bituminous limestone lithology is hosted by very low permeability rock. Hyperalkaline water conducting features are fracture controlled and generally also of low transmissivity, with few exceptions. The flow paths are rather complex based on the observed inability to obtain cross-hole responses even in bore hole sections that are separated by less than lm. At the exposed base of one of the metamorphic zones (Fig. 2, location 5 ) features with a very low conductivity (almost dry) are oriented parallel to the bedding immediately underlying meta-
placement. Locally, sections of intense fracturing occur, and these also appear to control the intensity and distribution of metamorphism, at least in the upper portions of metamorphic zones. This fracture system also defines the orientation of one of the hyperalkaline fluid conducting features. A second system of steep fractures trending NNW-SSE is associated with only minor displacement, and also controls fluid flow in the area of Adit A6. Similarly oriented fractures also weakly cut through alluvial terrace sediments as well as underlying strata, indicating that the regional tectonic regime is still active locally. Shallow listric mesoscale faults and open planar fractures that formed parallel to the valley slopes are geologically recent. They are interpreted to generally post-date the metamorphism as they affect cemented hill slope deposits that contain admixtures of metamorphosed and unmetamorphosed Chalky Limestone. Collapse structures possibly related to karstic features occur locally. The observed features appear to be generally parallel with the N-S fracture system and they may have started as faults, as a remnant cemented tectonic breccia margin was observed in one location. At this location the void space, up to 3m wide, is infilled with a chaotic assemblage of poorly sorted breccia comprising baked Chalky Limestone. The karstic features therefore post-date metamorphism, but the age of fault initiation is as yet unknown. There is also evidence that at least some karst features are older than the (undated) basalt immediately overlying the area. Although the earlier two structural systems pre-date metamorphism, they also show evidence of multiple reactivation at later times, albeit with only minor motion. This is in accord with some observations of multiply zoned vein fillings along the hyperalkaline flow paths which have been interpreted to infer repeated reactivation of flow systems triggered by tectonic movement (Milodowski et al. 2001).
4 HYDRAULIC PROPERTIES AND FLOW SYSTEMS The regional hydrogeologic situation is complicated by topography, karstification of the overlying Chalky Limestone Formation, and locally by brecciation (related to metamorphism, karst development and/or slope collapse) and weathering features. This has resulted in a range of hydraulic conductivities of 10-5to 10-7m / s for altered rock (including 'cement zone') and <10-7 m / s for unaltered rock, derived from well testing during site investigation of the proposed Unity Dam site (Smellie 1998). On the 187
morphic rocks, in contrast to dominantly steep fracture zones observed at locations 1 and 2. The characteristics of the exposed base of the metamorphic zone are remarkably similar to those observed on larger fossil hyperalkaline systems in Central Jordan (Rassineux et al. 2001).
ACKNOWLEDGEMENT
5 WATER-ROCK INTERACTION AND REACTIVE TRANSPORT MODELLING
REFERENCES
One objective of obtaining quantitative data on fluid conducting systems is their use in modelling of coupled processes of transport and mineral reactions. The detailed understanding of the geometry and hydraulics of the fracture controlled hyperalkaline flow systems is not yet far advanced. Nevertheless, insight from constrained geochemical modeling may already have pointed at some of the key issues detailed below. The chemistry of the hyperalkaline waters are discussed in Khoury et al. (1989, Smellie (1998), Alexander & Smellie (2000), and details of vein mineralization are provided by Milodowski et al. (2001), Linklater (1998), and Smellie (1998). The effects on the rock matrix adjacent to flow systems by diffusion of hyperalkaline fluid is reported in (Smellie 1998) and is a focus of presently ongoing invetsigations. Based on a data set available at that time, Steefel & Lichtner (1994, 1998) discussed the tendency of self-sealing of the rock matrix as a possible result of dolomite dissolution paired with formation of secondary calcite. The dual porosity model allowed for advective-dispersive transport along a fracture and diffusive transport in the adjacent rock matrix. The authors are able to relate the mineral alteration profile in the rock matrix resulting from reactive diffusion to the sequence of mineralization in the fracture. One crucial parameter to distinguish between a tendency to seal the fracture before or after sealing of the adjacent matrix are the relative reaction rate constants in the matrix and in the fracture, respectively. Cementation by secondary calcite proposed by the same mechanism had also been observed by Adler et al. (1999, 2001) in rock core infiltration experiments with a carbonate-bearing hard clay rock and artificial portlandite-buffered hyperalkaline fluid. The thrust of present and any future studies will be the quantitative recording of field data and their integration in a spatial 3-D data base as a point of departure for evaluating time scales, rates and geometric constraints on water-rock interaction, and its consequences for solute transport in the context of repository performance assessment criteria.
The authors acknowledge United Kingdom Nirex Limited, NAGRA, SKB, UK Environment Agency, Ontario Hydro, ANDRA and CEA for funding and support.
Adler, M., Mader U.K. & H.N. Waber 1999. High-pH alteration of argillaceous rocks: an experimental study. Swiss Bulletin of Mineralogy and Petrology 79: 445-454. Adler, M., Mader U.K. & H.N. Waber 2001. Core infiltration experiment investigating high-pH alteration of lowpermeability argillaceous rock at 30 "C. Proc. WRI-I0 (R. Cidu ed), Rotterdam: Balkema. This issue. Alexander, W.R. (ed.) 1992. A natural analogue study of the Maqarin hyperalkaline groundwaters. I. Source term description and thermodynamic database testing. Nagra Tech. Rep. NTB 91- 10. Wettingen, Switzerland: Nagra. Alexander, W.R. & J.A.T. Smellie 2000. The Maqarin Natural Analogue Project (1989-1998). Proceedings of the Eighth CEC Natural Analogue Working Group (NAWG) Meeting. September 1999, Strasbourg, France. In press. Khoury, H.N. & S. Nassir. 1982. High temperature mineralisation in the bituminous limestone in the Maqarin area northern Jordan. N. Jb. Miner. Abh. 144: 192-213. Khoury, H.N., Salameh, E. & 0. Abdul-Jaber 1985. Characteristics of an unusual highly alkaline water from the Maqarin area, north Jordan. J. Hydrol. 8 I : 79-8 1. Khoury, H.N., Salameh, E., Clark, I.D., Fritz, P., Bajjali, W., Milodowski, A.E., Cave, M.R. & W.R. Alexander 1992. A natural analogue of high pH waters from the Maqarin area of northern Jordan. I: Introduction to the site. J. Geochem. Explor. 46: 117-132. Linklater, C.M. (ed.) 1998. A natural analogue study of cement buffered, hyperalkaline groundwaters and their interaction with a repository host rock 11. Nirex Science Report, S/98/003. Harwell, UK: Nirex. Milodowski, A.E., Hyslop, E.K., Khoury, H.N., Hughes, C.R., Mader, U.K., Griffault, L.Y. & L. Trotignon 2001. Mineralogical alteration by hyperalkaline groundwater in northern Jordan. Proc. WRI-10 (R. Cidu ed), Rotterdam: Balkema. This issue. Mor, D. 1993. A time-table for the Levant Volcanic Province, according to K-Ar dating in the Golan Heights. J. Afr. Earth. Sci. 16: 223-234. Rassineux, F., Parneix, J.C., Griffault, L.Y., Smellie, J.A.T., Trotignon, L., Raynal, J., Khoury, H., & F. Mercier 2001. Mineralogical evolution of clay-bearing rock during alkaline alteration (Khushaym Matruk, Central Jordan). Proc. WRI-10 (R. Cidu ed), Rotterdam: Balkema. This issue. Smellie, J.A.T. (ed.). 1998. MAQARIN: natural analogue study: Phase 111. SKB Tech. Rep.. TR-98-04. Stockholm, Sweden: SKB. Smellie, J.A.T., Alexander, W.R., Degnan, P., Griffault, L.Y., Mader, U.K. & L. Trotignon 2001. The role of the Jordan natural analogue studies in the performance assessment of cementitious repositories for radioactive wastes. Proc. WRI-10 (R. Cidu ed), Rotterdam: Balkema. This issue. Steefel, C.I. & P.C. Lichtner 1994. Diffusion and reaction in rock matrix bordering a hyperalkaline fluid-filled fracture. Geochim. Cosmochim. Acta 58: 3595-3612. Steefel, C.I. & P.C. Lichtner 1998. Multicomponent reactive transport in discrete fractures I: Controls on reaction front geometry. Journal of Hydrology 209: 186-199.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Water quality changes during aquifer storage recovery in limestone-silicate aquifer material J.Mirecki Department of Geology, College of Charleston, Charleston, South Carolina, USA 29424
M.D.Petkewich, K.J.Conlon & B .G.Campbell U S . Geological Survey, 720 Gracern Road Suite 129, Columbia, South Carolina, USA 29210-7651
ABSTRACT: The U.S. Geological Survey is investigating the potential for implementation of several Aquifer Storage Recovery systems across the Charleston, South Carolina, peninsula. A pilot study, conducted between 1993 and 1995, showed recovery efficiencies ranging between 38 and 61 percent during nine Aquifer Storage Recovery cycles. Although the pilot study confirmed that Aquifer Storage Recovery methods were feasible in the aquifers beneath Charleston, questions were raised regarding the reactions of injected water with aquifer material, and mixing of native aquifer water during long-term storage conditions. A second, more complex Aquifer Storage Recovery study was initiated in 1998 to evaluate the geochemical and hydrologic effects of storing potable water for 1 to 6 months. Preliminary results from Aquifer Storage Recovery tests with 1month storage periods indicate recovery efficiencies up to 81 percent. A decrease in the transport time from the production well to the observation well was observed, indicating a probable increase in aquifer permeability. Geochemical modeling of water-quality data collected at depth in observation wells will quantify the dominant geochemical reactions taking place during Aquifer Storage Recovery cycling. Results from the study will indicate whether long-term storage of potable water will provide the city of Charleston, South Carolina, with adequate water for emergency use.
1 INTRODUCTION
Quaternary age; and the Santee Limestone/Black Mingo (SL/BM) aquifer, within limestones of Eocene age (Santee Limestone of the Cooper Group) and interbedded, bioturbated, muddy limestones, gray mudstones, and siltstones (Black Mingo Group) of Paleocene age (Campbell et al. 1997). The SL/BM aquifer is confined, and is separated from the overlying surficial aquifer by approximately 102 meters of consolidated calcarenite and calcilutite, known locally as the Cooper marl. Injection, storage, and recovery of treated drinking water occur within two permeable zones near the contact of the Santee Limestone and Black Mingo Group sediments. The upper production zone (UPZ) occurs at depths of 115-123 m. Permeability in this zone is characterized by karstic solution-type, centimeter-sized openings. The lower production zone (LPZ) occurs at depths of 129-132 m. Permeability in this zone is characterized by fractures within semi-consolidated sandstone and inter-bedded limestone. Core samples of aquifer material at these permeable zones were obtained in cores obtained by the U.S. Geological Survey (USGS) from drill holes CHN-733 (at the pilot site), and CHN-800 (Calhoun Park core) located approximately 400 m northwest from the Charleston ASR site.
The primary source of potable water for the city of Charleston, South Carolina, is treated surface water from the Edisto and Back Rivers. Although the Charleston Commissioners of Public Works (CCPW) has a treatment capacity that far exceeds normal demand, there is a concern that demand may exceed delivery capacity if the water distribution system is disrupted during a hurricane, earthquake, or hard freeze. The densely populated peninsular section of Charleston (12,475; 1990 census) is particularly vulnerable, due to deteriorating cast-iron distribution lines. For this reason, CCPW in conjunction with the U.S. Geological Survey (USGS) is evaluating the feasibility of several Aquifer Storage Recovery (ASR) systems across the Charleston peninsula, primarily for the storage and intermittent recovery of an emergency potable drinking water supply (Fig. 1).
2 HYDROGEOLOGY Local hydrostratigraphy consists of two aquifers: the surficial aquifer, within unconsolidated sands of 189
from west to east toward the City of Charleston. These waters typically are “black water streams” characterized by high dissolved organic carbon concentration (20 to 50 mglL), acidic (PH 5), and having low carbonate alkalinity (less than 50 mg/L as CaC03), dissolved chloride (less than 30 mg/L) and specific conductance values (less than 300 microsiemendcm). Raw surface water is treated at the plant by successive sand pack and carbon filtration, then chlorination. After solids are removed, carbonate alkalinity is increased with the addition of lime, thus raising the pH. Chloramines are added toward the end of the process to minimize the formation of trihalomethane precursors. Orthophosphate is added to prevent lead leaching from older water mains and household plumbing. Figure 1. Map showing locations of the pilot ASR site (19931995) and larger Charleston ASR site (1998-present), in Charleston, South Carolina.
The number and arrangement of injection and monitoring wells was expanded at the Charleston ASR site based on experience at the pilot ASR site. The Charleston ASR site consists of a single injection well surrounded by three variably spaced monitoring wells (Fig. 2). The 40.6 cm diameter injection well (CHN-812) is cased with stainless steel to the top of the UPZ (105 m), then narrows to 20 cm diameter open-hole to the base of the well (440ft, 135 m). There are three screened intervals in the open-hole portion of CHN-8 12, corresponding to depths of the UPZ and,LPZ defined by geophysical logs. A 25-horsepower pump is installed to a depth of 97 m in CHN-812. Monitoring wells (CHN-809, 810; and -8 11) are cased with PVC to the upper contact of the Cooper Group confining unit (approximately 34 m). Monitoring wells CHN-809, CHN-8 10, and CHN-8 11 were installed at distances of 23.4, 37.5, and 150 m respectively, from the production well, specifically to allow characterization of aquifer hydraulic properties, and also to monitor injected water movement and water-quality changes that occur during ASR tests. Water quality differs significantly between end members, represented by treated drinking water from a water supply main connected to the injection well CHN-812, and native aquifer water of the SL/BM aquifer These compositions will be discussed in detail in the following section.
Ground-water from both permeable zones of the SL/BM aquifer shows brackish water-quality characteristics. Ground waters have moderately high carbonate alkalinity (300 to 500 mg/L as CaCO3), dissolved chloride (up to 1,900 mg/L) and specific conductance values as high as 8,000 microsiemendcm. SL/BM ground water samples at the Charleston ASR site degas hydrogen sulfide when recovered; we are quantifying these concentrations at present to assess redox state of the aquifer. 4 RECOVERED WATER QUALITY Interpretations of recovered water quality and geochemistry obtained from the ASR pilot site defined the characteristic effects of water-rock interactions and mixing of ground waters (Mirecki et al. 1998). Interpretations of water quality and geochemical changes at the Charleston ASR site are in progress. Many of the geochemical reactions that were quantified at the pilot site, are also prominent at the Charleston ASR site. Water quality changes are interpreted in context of regulatory guidelines, primarily the Maximum Contaminant Levels (MCLs) established for several solutes by the US Environmental Protection Agency (USEPA 2000). These solutes include chloride (MCL is 250 mg/L) and sulfate (MCL is 250 mg/L). Trace metals such as copper and lead are below detection in recovered waters. Nutrients such nitrate, nitrite, and phosphate are below their respective MCLs. Chloride measurements allow calculation of recovery efficiency, defined as: Recovery Efficiency
3 SOURCE WATER QUALITY Treated drinking water injected into the SL/BM aquifer during ASR testing is piped from a water main connected directly to the injection well (CHN-812). The drinking water source is surface water from the Bushy Park Reservoir and Edisto Rivers, which flow
=
Injected water volume < 250 m d L C1total volume of water injected
Recovery efficiencies ranged between 38 and 61 percent during nine ASR cycles at the ASR Pilot Site (Campbell et al. 1997). Preliminary results from tests at the Charieston ASR site having 1month storage periods indicate recovery efficiencies up to 81 percent. 190
Figure 2. Map showing arrangement of injection well (CHN-812) and monitoring wells (CHN-809, -810, -81 1) at the Charleston ASR site.
Geochemical changes during ASR testing are interpreted using USGS geochemical modeling codes such as NETPATH (Plummer et al. 1994) and PHREEQC (Parkhurst 1995). Dominant geochemical reactions that proceed during storage were quantified by an inverse modeling approach: solute concentrations from injected water represent an initial condition, and water recovered after storage represents a final condition. Mass transfer between water and rock for major and trace dissolved constituents can be calculated using NETPATH. This approach allows us to determine the dominant geochemical controls on water quality that occur during storage in the limestone/clastic SL/BM aquifer. At the pilot ASR site, carbonate dissolution and increasing ionic strength are the major reactions that occur during storage (Mirecki et al. 1998). Carbonate dissolution is estimated to have enhanced aquifer permeability approximately 1 to 3 percent based on mass transfer calculations. More detailed water quality and geochemical interpretations, specifically related to changing aquifer redox condition are in progress at the Charleston ASR site.
below the MCL for chloride at well CHN-809. Breakthrough curves are defined using specific conductance trends from probes placed at the upper and lower permeable zones, supplemented with waterquality data from ground-water samples collected weekly at depth using a Bennett pump. The duration of storage is 1-month, 3-months, or 6-months, during which water-quality samples are collected biweekly. Injected water is recovered at an approximate pumping rate of 8.2 L/s, and continues until samples show chloride concentration and specific conductance values equal to pre-test conditions. Water-quality samples are collected only from the production well head during the recovery stage. 6 ANALYSIS OF DATA Prior to injection of treated drinking water, an aquifer test was conducted to determine the hydraulic characteristics of the SL/BM aquifer. Aquifer-test data were collected from the production well and several observation wells screened or open to the SLBM aquifer, and a well screened in the overlying confining unit. Aquifer-test results from the Charleston ASR site are similar to results from the ASR pilot site (Campbell et al. 1997). Analytical and numerical methods indicate an aquifer transmissivity of 37 m2/day, storage coefficient of 4.3 x 10-5, and anisotropy in permeability with the maximum on a bearing of 116 degrees being 6 times that in the orthogonal direction. Analytical and numerical results also indicated leakage to the aquifer from the overlying confining unit. Aquifer testing after completion of all the Phase 11 cycles will indicate whether ASR testing has increased SLBM aquifer permeability due to mineral dissolution.
5 OPERATION OF THE SITE The following describes a typical ASR cycle during testing at the Charleston ASR site (now ongoing). Each ASR cycle consists of an injection, storage, and recovery period. The length of the injection phase (and hence volume of injected water) is determined by the breakthrough of ‘fresh’ (low chloride concentration) water at the proximal observation well (CHN-809). Treated drinking water is injected at an approximate rate of 0.7 L/s. Injection proceeds until the chloride concentration decreases 191
As of December 2000, two complete ASR cycles (with 1-month storage periods) and the injection phase of a 3-month-storage cycle have been completed. During the second and third ASR tests, chloride concentration decreased to the MCL more rapidly (29 days) during breakthrough at well CHN-809 than the first ASR test (78 days). Faster breakthrough following the first injection may indicate that some fresh water residual has remained in the aquifer or that permeability has been enhanced by mineral dissolution. This decrease in time of travel also was observed during the pilot ASR project (Mirecki et al. 1998). Increases in recovery efficiencies with successive ASR tests are typical of ASR development (Pyne 1995). Recovery efficiencies are relatively higher during the Phase I1 investigation compared to the pilot study. Whether the higher efficiencies at the Phase I1 site are due to greater volumes of injected water, lower injection rates, differences in the design of the production wells (open-hole well construction at the pilot site), or different storage periods has yet to be determined.
7 CONCLUSIONS Results from the Charleston ASR project will address the unanswered questions from the pilot study. Using improved sample-collection methods, the USGS is capable of quantifying water-quality changes that occur during storage, and estimating when equilibrium between rock and ground water occurs. Potential injection and recovery rates will be estimated and used to determine whether or not these rates increase with successive ASR testing due to increased permeability. Phase I1 investigation results will indicate whether SL/BM aquifer properties are enhanced or degraded during long-term storage of treated drinking water. ACKNOWLEDGEMENTS The authors acknowledge the Charleston Commissioners of Public Works for continued support of ASR research and applications in the coastal plain of South Carolina.
REFERENCES Campbell, B.G., Conlon, K.J., Mirecki, J.E. & M.D. Petkewich 1997. Evaluation of aquifer storage recovery in the Santee Limestone/Black Mingo aquifer near Charleston, South Carolina, 1993-95: U S . Geological Survey WaterResources Investigations Report 96-4283, 89 p. Charleston Commissioners of Public Works 2000. Monthly Drinking Water Quality Report for September 2000. Accessed on December 18,2000 at: http://www.charlestoncpw.com/monthlywaterquality. htm.
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Mirecki, J.E., Campbell, B.G., Conlon, K.J. & M.D. Petkewich 1998. Solute changes during aquifer storage recovery in a limestone/clastic aquifer. Ground Water 36(6): 394-403. Parkhurst, D.L. 1995, User’s guide to PHREEQC-A computer program for speciation, reaction-path, advectivetransport, and inverse geochemical calculations: U.S. Geological Survey Water-Resources Investigations Report 95-4227, 143 p. Also available at: http://water.usgs.gov/software/geochemical. html. Plummer, L.N., Prestemon, E.C. & D.L. Parkhurst 1994. An interactive code (NETPATH) for modeling NET geochemical reactions along a flow PATH, version 2.0. U.S. Geological Survey Water-Resources Investigations Report 944169, 130 p. Also available at: http://water.usgs.gov/software/geochemical.html. Pyne, R.D.G. 1995. Groundwater recharge and wells. A Guide to Aquifer Storage Recovery. Lewis Publishers. US Environmental Protection Agency, 2000. Current Drinking Water Standards. Accessed on December 20,2000 at: http://www.epa.gov/safewater/consumer/mcl.pdf.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Analytical model for deep well injection of cold brine into a hot aquifer A.F.Moench & Y.K.Kharaka U S . Geological Survey, Menlo Park, California, USA
ABSTRACT: An analytical model was used to predict temperature and chemical concentrations that result from the constant injection of cold, hypersaline brine in a hot, deep-seated limestone aquifer. Due to the presence of high concentrations of sulfate and calcium in the brine there is concern that precipitation of anhydrite will occur rendering the aquifer unusable for the planned disposal of the brine that presently contaminates an important surface water supply. The model assumes the aquifer to be a homogeneous, densely fractured system composed of intersecting, highly permeable fractures separated by low-permeability blocks. Based on selected values of aquifer dimensions, and realistic hydraulic, thermal, and chemical properties for the rock and fluid, results show possible locations of the thermal and chemical fronts and/or breakthrough curves. From these results it may be possible, using geochemical modeling, to infer when or where in the aquifer the precipitation of anhydrite is likely to occur and to make recommendations for alleviation of problems related to aquifer plugging. 1 BACKGROUND The Colorado River and its tributaries annually provide nearly 20 billion m3 of water for municipal and industrial use of more than 20 million people and for irrigation of about 1.2 million ha of land in seven western States in USA. Anthropogenic activities, primarily irrigated farming, have caused the salinity of the Colorado River water to be among the highest of the World’s major rivers. Prior to significant human activities in the Colorado River basin the salinity of pristine water was estimated to be 334 mg/L; currently the salinity at Hoover Dam is about 700 mg/L, and has ranged to values close to 1,000 mg/L (US. Department of Interior 1999). In order to limit the damaging impacts of high salinity to farming and other water uses in the USA and the Republic of Mexico, the US Federal government and the Colorado River basin States authorized several salinity control projects. High salinity in the lower reaches of the Colorado River is due to some extent to the seepage of hypersaline brine, with a salinity of 250,000 mg/L, from an unconfined aquifer adjacent to one of its tributaries in Colorado, USA, the Dolores River. To alleviate the problem, which results from the addition of -200,000 tons/yr of salt to the Colorado River, the Bureau of Reclamation is pumping brine (ultimately at a rate of -2000 m3/d) from 11 wells situated along the banks of the Dolores, piping it to
one of the deepest (total depth of 4.9 km) disposal wells in the world, and injecting it into the Mississippian Leadville Limestone at a depth of 4.3 km (U.S. Bureau of Reclamation 1978). The cold (10°C) brine is injected into a hot (-120°C), fractured-limestone aquifer. Because the brine contains high concentrations of sulfate (about 6,000 mg/L) and calcium (-1,400 mg/L), geochemical modeling indicates that the brine is at saturation with anhydrite at 10°C and highly supersaturated at reservoir conditions. Geochemical modeling indicates, and laboratory water-rock interaction experiments (Rosenbauer et al. 1992) confirm, that large amounts (5,000-10,000 kg/d, depending on injection rate) of anhydrite, which has reverse solubility, will precipitate from the brine, reducing the fracture permeability in the aquifer. In addition, it is known that the formation water of the Leadville Limestone is highly incompatible with the brine at downhole conditions leading to additional anhydrite precipitation if the two fluids come in contact. If magnesium in the brine dolomitizes the calcite of the aquifer, anhydrite precipitation could then increase by up to a factor of three (Kharaka et al. 1997). It has been demonstrated (Kharaka et al. 1997) that it is feasible to remove a large (>95%) percentage of the dissolved SO4 prior to injection by use of nanofiltration membranes. Experiments are currently being conducted whereby brine concentrations are 193
reduced by dilution with Dolores River water before injection (-70% brine and 30% river water). The remaining sulfate could continue to be a problem if mixing with formation water and/or dolomitization occurs and may ultimately render the aquifer unusable for brine disposal. Drilling of the injection well was completed in 1988. The stratigraphic chart for the well shows the top of the Leadville Limestone to be located at a depth of 4293 m. It has a thickness of 127 m over which perforations for fluid injection were distributed. The formation is described as locally vuggy, fractured dolomitic limestone with an effective porosity of less than 6 percent. The permeability is enhanced by wide-spread fracturing that also tends to make the reservoir permeability anisotropic in the horizontal plane. Formations that overlie and underlie the Leadville Limestone have relatively low permeability and therefore serve as confining layers. A large volume of water and brine has been injected into the formation since operations began in 1991. The injection rate increased substantially after 1995, apparently resulting in the desired aquifer cooling but also resulting in the production of thousands of small and a few moderate (up to magnitude 4.3) earthquakes, probably as a consequence of elevated fluid pressures and fault lubrication. For the most part the earthquakes are of such small magnitude as to be imperceptible to people, but a dozen or so events have been large enough to be felt by local residents. In this paper, as a preliminary assessment, the thermal response of the hot Leadville Limestone to the constant injection of cold brine is provided under idealized but nevertheless realistic conditions. Using a simple analytical model and assumed fluid and rock properties, theoretical temperature and chemical fronts in the aquifer at specified times from the start of injection, and theoretical temperature and chemical breakthrough curves at a specified distance from the injection well are obtained. The ultimate objective is to infer where in the aquifer anhydrite precipitation in significant amounts is likely to occur.
hydraulic properties of the fracture system are the result of chemical processes, thermal stresses, and tectonic processes. Figure 1 shows a schematic diagram of the physical system. Within the confines of an infinite aquifer of constant thickness, heat and chemicals are advected horizontally through the fracture network and simultaneously diffuse into or out of the adjacent blocks. The diffusion of chemicals in the blocks occurs in the liquid phase and is limited by the effective block porosity. The diffusion of heat in the blocks occurs in both the liquid and solid phases but is dominated by the solid phase.
Figure 1. Schematic diagram of a double-porosity aquifer of thickness, H, with a steady-state flow field established around an injection well. Flow velocity due to constant injection, Q, is a function of radial distance, r. The model was developed under the following general assumptions: 1. An injection well of finite diameter fully penetrates a horizontal, confined, double-porosity aquifer of constant thickness and of infinite radial extent. 2. The aquifer is densely fractured and composed of equal-sized cubical blocks that are approximated as spheres for the sake of mathematical tractability. 3. A steady-state flow field, which is radially divergent, one-dimensional, and axially symmetric with respect to the injection well, is established instantaneously in the fracture system with the onset of injection. 4. The hydraulic and thermal properties of the formation are homogeneous and isotropic and the fluid properties do not vary with space or time. 5. The upper and lower boundaries of the confined aquifer are thermally insulated. 6. The chemicals are non-reactive and non-sorptive.
2 MATHEMATICAL MODEL The analytical model used in this paper is a modification of a model for radial dispersion in a double-porosity aquifer published by Moench (1987). In this model the term “double-porosity” refers to the fractured-rock mass as two interacting, overlapping continua: 1. a continuum of lowpermeability, primary porosity blocks and 2. a continuum of high-permeability, secondary porosity fractures. The porosity and permeability of the primary porosity blocks are largely the result of depositional and lithification processes. The 194
Advection and longitudinal mechanical dispersion dominate the transport of heat and chemicals in the fracture system. D i f k i o n dominates the movement of heat and chemicals in the porous blocks in accordance with the laws of Fourier and Fick, respectively. Mechanical dispersion is linearly related to velocity and is therefore a function of radial position.
and
where T and C are the local temperature and concentration, respectively, in the aquifer. The subscript inj refers to the injection temperature and concentration. The subscript aq refers to the initial aquifer temperature and concentration. Results of computations (see Figure 2) show[s] the thermal and chemical breakthrough curves for two values of fracture porosity as would be observed at a distance of 50 m from the injection well. Based on these results it takes as long as 10 years to reduce the aquifer temperature (at a distance of 50 m) to the brine injection temperature, whereas the resident aquifer fluid is replaced by the injected brine at this distance within 20 or 200 days from the start of injection, depending on the value of the chosen fracture porosity.
Moench (1987) used the method of Laplace transformation to solve the controlling differential equation (advection-dispersion equation for plane radial flow) in combination with appropriate boundary and initial conditions. A Laplace transform solution was obtained and was evaluated by numerical inversion using the well-known Stehfest (1970) algorithm. 3 RESULTS & DISCUSSION
The Laplace transform solution was evaluated for this paper using the parameters indicated in Table 1. Table 1. Aquifer parameters required by the model. Property Value injection rate 2000 m’/day aquifer thickness 120 m fracture aperture 0.01 m block porosity 0.005 1 m2/day free-water chemical diffusion 8 . 6 ~O-’ coefficient 8 . 6 1~0-2m2/day thermal diffusivity 0.6 ratio of the heat capacity of limestone to the heat capacity of brine see Figures 2 & 3 fracture porosity see Figures 2 & 3 aquifer dispersivity -
~~~
The only differences between the model of Moench (1987) and that used in this paper are that the “fracture skin”, a coating of material on the blocks that provides resistance to the diffusion of chemicals into the blocks, is not included (for simplicity) and, for the heat component, the effective diffusion coefficient for chemicals in the block system (D’ in Moench 1987), becomes the coefficient of thermal diffusivity, which is the thermal conductivity divided by heat capacity of the rocWfluid system. Based on the analytical model and the parameters given in Table 1 the location of breakthrough curves and the thermal and chemical fronts in the aquifer are given in Figures 2 and 3. The dimensionless temperature TDand concentration CDare defined as:
Figure 2. Thermal and chemical breakthrough curves at a distance of 50 m from the injection well. ( ais~ the longitudinal dispersivity for the fracture system and $f is the fracture porosity.) Results in Figure 2 show that the chemical fronts lead the thermal fronts by amounts that depend upon the chosen fracture porosity. Also, the difference in the location of the thermal fronts for the two fracture porosities is imperceptible. The explanation for these results lies in the fact that the thermal diffusivity of limestone is about five orders of magnitude greater than the effective chemical diffusion coefficient of the porous blocks, which is the free-water chemical diffusion coefficient multiplied by the block porosity.
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CONCLUSIONS
Figure 3. Thermal and chemical fronts after 100 days of injection.(al is the longitudinal dispersivity for the fracture system and +f is the fracture porosity). Sensitivity analyses show differences in the results if the injection rate, duration, or other parameters shown in Table 1 are modified. The eight assumptions listed above are reasonable first order characterizations. Beyond that, perhaps the hardest assumptions to justify are 4,5 , and 6, which can introduce significant uncertainties into the computed results. The aquifer is known to be heterogeneous and anisotropic (assumption 4) but without additional near-well data these properties are impossible to quantify. The upper and lower boundaries of the confined aquifer are treated as thermally insulated (assumption 5), allowing no reheating of the cooled parts of the aquifer from the overlying or underlying confining layers even when water injection is stopped. Conduction of heat from the upper and lower confining layers should result in a larger separation between the chemical and thermal breakthrough curves and fronts shown in Figures 2 & 3. That the chemicals involved are reactive (assumption 6) and likely to result in aquifer plugging is well known (Kharaka et al. 1997). Results indicate that the bulk of the injected brine would be at temperatures much higher than 10°C. These have profound implications to the injection program. Geochemical modeling indicates that the unmodified brine is at equilibrium with anhydrite at 1O°C, and that anhydrite precipitation increases essentially linearly with temperature, between 10 and 120°C (Figure 1 in Kharaka et al. 1997). Some lag time between brine heating and anhydrite precipitation may occur, especially at the lower temperatures, but eventually thermodynamic equilibrium will be achieved after mineral precipitation. These results indicate that sulfate removal or brine dilution before injection are necessary to prevent mineral precipitation and aquifer damage.
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Breakthrough curves, at a given distance from the injection well, illustrate that it takes considerably longer to cool the aquifer to brine injection temperature than to replace the resident fluid in the aquifer. The thermal fronts in the aquifer lag behind the chemical fronts by an amount that depends upon the fracture porosity (the smaller the fracture porosity the greater the lag) and the position of the thermal fronts is independent of the fracture porosity. One can infer that if brine is injected into the aquifer after a prolonged period of cold-water injection, the brine front will overtake the thermal front and, depending on the chemical saturation conditions, anhydrite will likely precipitate making the aquifer unusable for brine disposal. REFERENCES Kharaka, Y.K., Ambats, G., Thordsen, J.J. & Davis, R.A. 1997. Deep well injection of brine from Paradox Valley, Colorado: Potential major precipitation problems remediated by nanofiltration. Water Resour. Res. 33: 10131020. Moench, A.F. 1987. Radial dispersion in a double-porosity system with fracture skin. Proceedings of the Twelfth Workshop on Geotherrnal Reservoir Engineering, Stanford University, Stanford, California 20-22 January 1987: 125129. Rosenbauer, R.J., Bischoff, J.L., & Kharaka, Y.K. 1992. Geochemical effects of deep-well injection of the Paradox Valley brine into Paleozoic carbonate rock, Colorado, U.S.A. Appl. Geochem. 7: 273-286. Stehfest, H. 1970. Numerical inversion of Laplace transforms. Cornrnun. ACM. 13:47-49. U.S. Burea of Reclamation. 1978. Paradox Valley Unit definite plan report. Appendix B Hydrosalinity. BOR, Denver, CO. 173PP: U.S. Department of Interior. 1999. Quality of water Colorado River Basin. BOR Prog. Rep. No. 19.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, iS6N 90 2651 824 2
Examination of the Effect of Uncertainty in Thennodynamic and Kinetic Data on Computer Simulations of Complex Systems Craig H .Moore Craig H . Moore Research and Applications, Bloomington, Indiana, USA
ABSTRACT: Many researchers hesitate to take on computer modeling and simulations projects because of a lack of good data. This is especially true when kinetic data are required for the project. Computer simulations of complex systems using a large number and range of thermodynamic and kinetic constants suggest that, for a number of applications, the lack of good data is not always good reason to not perform a study. 1 INTRODUCTION Many times geochemists are faced with the prospect of having to simulate or model systems for which thermodynamic and kinetic data are not well known or must be estimated. This is especially true in solving “real world problems” often found in projects for industry or government. Although it is wise to be skeptical about the validity of the result of such simulations, many times the system are not that sensitive to the choices of values for the constants or are sensitive to only a few of them. Therefore it may be unwise to reject a project simply because of some poor data.
2 EXAMPLE As an example of this, let us look at the result of coupled flowheaction modeling of the dissolution of an immature sandstone by a hot aqueous fluid. The minerals in the sandstone and their initial concentrations (volume percent of whole rock), for the purposes of the simulations were: quartz (12.3); oligoclase (2.58); K-feldspar (10.4); basaltic rock fragments (5.55), rhyolitic rock fragments (5.51); chert (5.84); siderite (29.9); illite (0.23); kaolinite (0.04); dioctahedral smectite (0.67); trioctahedral smectite (5.9). Other minerals that were considered in the simulation were: muscovite, biotite, calcite, dolomite, tnoctahedral smectite; septechlorite; ripidolite, Caanalcite, K-clinoptilolite, amorphous silica, and albite. Their initial concentrations were zero. The physical situation was the injection of hot aqueous fluid into a formation via an injection well.
For this system, 70 different simulations were run. Each run consisted of varying some equiIibrium constants andor rate constants by as much as two orders of magnitude, in each direction, and also varying some grain sizes and fluid velocity. The results of these simulations were then compared with a simulation that was considered to be a ‘base line’ simulation. Many different types of comparisons can be made and here I show just a few examples. The lithology whose composition was given above is an example of a system where parameter sensitivity is very high and yet still not so high, it can be argued, as to render the results not usehl. The following graph shows a comparison of the magnitudes of changes in mineral amounts as compared to the base line simulation. I have termed these changes ‘excursions’. Figure 1 shows the excursions for all minerals that had a non-zero concentration in any of the simulations. As an example of the interpretation of these data, consider the mineral albite. In 34 of the 70 simulations the total amount of albite present in the system varied by less than or equal to 10 percent; in 11 of the simulations the amount varied between 11 and 50 percent; in 25 of the simulation the amounts varied by greater than 50 percent of the total in the base line simulation. For the mineral siderite, however, its total amount in all the simulations varied by less than 10 percent. Of those minerals where a number of large excursions took place, eight of them were initially not present in the system and, of those, four had total amounts less than .1 volume percent in all of the simulations that is they were essentially insignificant. The remaining four constituents, albite, amorphous silica, biotite, and K-clinoptilolite had maximum total concentrations of 1.16, 1.14, 1.34, and 2.39, respectively. The constituents that were present in major amounts that &d show large excursions were the two
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Figure 1. The number and amount of excursions for 70 different parameter sets
Figure 2. The sensitivity of the system to some selected parameter sets.
types of volcanic rock fragments and the two smectites. This is not surprising given that the rock fragments are highly unstable constituents and that the smectites are the immediate result of the reaction of the rock fragments. This information points to closer examination of those constituents. Figure 2 shows a graph of the number of times changing a particular parameter resulted in any mineral in any simulation having an excursion in the top three excursions for that mineral in all simulations. This is a gauge of the entire system’s sensitivity to any particular parameter set. As an example, when the rhyolitic volcanic rock fragment’s equilibrium constant was given a value a factor of 1/100 of the base value, two minerals had excursions that ranked in the largest increases in concentration for that mineral compared to all the simulations. Nine minerals had decreases that ranked in the largest three decreases for all simdations, and, in total, 11 of the 21 minerals showed their largest excursions for t h ~ sparameter change. Note that, in this scheme, it was not
considered an excursion if the change was in the constituent whose corresponding parameters were changed. That is, if a rate or equilibrium constant for a mineral was changed, changes in that mineral’s concentration was not counted as a ‘significant effect’. The seven parameter changes shown counted for 63 percent of the system sensitivity when measured in this way. Not only total mineral amounts, but maximum and minimum local amounts can also be treated in a similar fashion. When this is done, it is clear that the rate constants, equilibrium constants, and grain size of the volcanic rock fragments are far and away the parameters to which the system, as a whole, is most sensitive. When the simulations that are being considered include information regardmg the spatial andor temporal distribution of system constituents, these data need also be examined in light of the parameter changes. Figure 3 shows a graph of the predicted spatial distribution of the mineral biotite in this flow-through system after 40 days of fluid flow. In 198
this graph, and those following, the lines are labeled as follows: RG = rhyolitic volcanic fra&ments;BG = basaltic volcank fragments; KC = K-clinoptilolite, AS = amnpbous silica; Q = equilirimn constant; K = rate constant; D = grain diameter; V = velocity. So. a line labeled B G & R G Q x 10. D x 5 was the result of a simulation in which the equilibrium canstant for both rhyolitic and basaltic volcanic rock fragments was multiplied by 10 and the grain diameter was mulbplied by 0.5. Basc refers to the shulation to which t he others were compared in the prenous discussion. Figure 3 shows that for the mineral hiotite, the spatial distributions and amounts of the mineral were similar for all simulations plotted, except for one. For biotite, line RG Q x .01 represents the simulation for which the lowest total amount of bictite was predicted to occur. Line BG & RG Q x 10, D x .5, represents the simulation for which the greatest local concentration of biotite occurred although this is not the simulation in which the largest total amount of hiotite m d . The BG & RG Q X 10, D x .5 line represents the largest local concentration predicted. Figure 4 is a similar plot for the mineral dolomite. Note, however, that the amount of dolomite present is very small. Once again, the RG K x .01 Line represents the smallest total amount predicted to form and the BG & RG Q x 10, D x .5 line represents thdargest local concernation predcted. The same situation exists for the mineral dioctahedral smectite as shown in A w e 5 Generally, the same situation exists for most of the minerals consid& in these simulations and it is tempting to say that, for this system the BG & RG Q x 10, D x .5 simulation and the RG K x .01 simulation bracket the expected system behavior in a general way if you have confidence that your estimates for equilibrium rate constants are within two orders of magnitude, plus or minus, of being correct.
The question then becomes one of wbether it is important if system behavior varies within this envelope. The lithology chosen for the above discussion was one of three examined in this study and was the one that showed the most effect of varying these parameters. Onc of the others contained fewer volcanic rock fragments and much less carbonate and the last contained fewer volcanic rock fragments, much less carbofiate, and more clay. In both these cases, excursions and variations were fewer and much analler in range than those shown. For the purposes of the particular project for which this work was being done, the general trends found in the simulations and the identification of important parameters was sdcient justification for the work performed. This is likely to also be !me in many other "realworld" projects.
3 CONCLUSION The data presented here, along with the other resdts of the simulations, show, I believe, that even in systems that can be expected to have a high sensitivity to changes in parameters, it is quite possible that knowing those parameter values even to only with a few orders of magnitude can still yield useful and defendable results. One should not assume, a priori, that the lack of wellconstrained and verified data would automatidly result in useless or questionable results. In many situations where we are dealing with real systems such uncertainty can be tolerated.
Figure 4. The predicted distribution of dolomite for a selected set of parameters.
Figure S. The predicted distribution of dioctahedral smectite for a selected set of parameters.
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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Evolution and self-organization of the w ater-rock system S .L.Shvartsev Institute of Oil and Gas geology SB RAS, Tomsk, Russia
ABSTRACT: It's shown that a water-rock geological system satisfies all essential attributes of selforganizing matter as considered in terms of modem synergy. The main premise of the self-organization processes is its equilibrium-nonequilibrium character determining the geologically long and strictly directional evolution with the formation of numerous new mineral phases and new geochemical water types throughout the Earth's crust. The main code of information directing the evolutionary development of this system is laid down in the structure of water solution.
origin of synergy self-organization phenomena in the area far from equilibrium. Let us consider this problem in more detail.
1 INTRODUCTION During many years we have studied water-rock equilibrium problems. As a first stage our research concerned fresh and salt waters in weathered zones (Shvartsev 1998a). Then we started to investigate these problems with respect to salt waters of sedimentary petroleum bearing basins of the West Siberian platform (Shvartsev et al. 1982). Recently we have also studied the equilibrium of strong brines of the chloride-calcium type of the Siberian platform with water-containing rocks (Shvartsev 1998b). Based on the large amount of factual material about equilibrium of different ground water types, we put forward the idea of an equilibriumnonequilibrium state of this system (Shvartsev 1995). According to this statement, under natural conditions, independent of depth and solution movement velocity, water is always out of equilibrium with respect to the minerals it dissolves. At the same time, it is in equilibrium with the other minerals it forms. It is important that the secondary mineral complex of this or that composition, equilibrium to water, develops during the entire period of rock-water system interactions at different depths rather than when local equilibrium is set in. An equilibrium-nonequilibrium state of the-rockwater system determines its ambiguity, promoting a continuous interaction due to the tendency to equilibrium, but nowhere and never fully achieved for a number of reasons under the geological conditions. Nonequilibrium of the system drives it to a continuous inner self-development, the formation of new structural-spatial formations (secondary minerals and geochemical water types), and the
2 THE ORIGIN AND COMPLICATION OF SPATIAL-TEMPORAL STRUCTURES OF WATER-MINERAL SYSTEMS For synergy, in contrast to cybernetics, nonequilibrium is not a source of death or the state of destruction. Vice versa, it is the basis for establishing an order, the cause of structural genesis, and the system's evolution as a whole. Nonequilibrium is a motive force for the evolution that leads finally to changes in irreversible energy (and matter) f l u e s arising when the closed megasystems tend to equilibrium. In these terms, the equilibrium-nonequilibriumnature of the rock-water system is the main factor of the evolution of mineral substance in diverse processes of its selforganization and self-development (Prigogine & Stengers 1984). One of the most important indicators of the system's self-organization is a purposeful origin progress, and disintegration of new spatial-temporal structures. It is the interaction of water with rocks, primarily aluminosilicates that fits best this condition. Thus, hydrolysis of aluminosilicates leads to the destruction of an initial solid phase, the chemical destruction of water with the simultaneous formation of a new mineral phase of different composition: MSiA10n+H20=M++OH"+[ Si(OH),-*]n3'1-[A10(OH)6]2201
Therefore, while interacting, the two phases (water and rock) give rise to a new (secondary) formation. It can be formed either in place of a mineral destroyed by water with repetition of the structure of this mineral (pseudomorphous substitution) or at a considerable distance from it. In any case the third component arises from the two system’s components (water and rock), which is an explicit indication of the system’s complication and development. The originating secondary mineral phase forms new structural elements thus giving to the developing system new properties not intrinsic to it before. At the same time, the composition of the water solution itself changes forming a new geochemical medium based on the chemical components obtained from the rock and the water. This new geochemical medium influences the nature of the secondary mineral phase that finally modifies its own composition. With weathering instead of kaolinite, for instance, montmorillonite starts to form. Therefore, the system is getting more complicated because a new structural element arises, better adjusted to certain conditions of the medium and reflecting a more profitable state in terms of energy. The formation and growth of ‘secondary mineral phases in integrity with the geochemical medium is a self-developing process. This is so because it proceeds in a direction from the separated parts toward the bonded ones, which, according to Ashbey (1962) is the most important indication of the selforganizing system. The formation of the secondary mineral phase represents in essence the origin of a qualitatively new system consisting of new (additional) structural elements. They are more regulated and adapted to the environmental conditions. Therefore, they gradually expand from the local area (the point of origin) covering increasingly new geological space. Examples are ultra-fresh siliceous waters of the tropical regions associating with laterites, soda water of forest-steppes associating with the secondary carbonates, strong brines of chloride-calcium type related to gypsiferous-saliferous formations.
Correspondingly, one geochemical type of water converts into another, one mineral phase transforms into another.
Figure 1. Principle scheme of the relationship of dispersed and concentrated elements during water-rock system evolution. A I - points, reflecting the generation initiation of the new secondary phase. Notes: 1 - the curve of the dispersed element concentration; 2 - elements concentrated by the secondary phases; 3 - zones of probable water flow to the contact with rocks; 4 - places where the water interaction with rocks probably stops.
In Figure 1, an example is given of the succession of the secondary mineral fomation observed in nature. Water moving in rocks is successively saturated relative to hydroxides of Fe and Al, then toward clay minerals, -calcite, gypsum, etc. Passing to equilibrium to a new mineral, water always retains equilibrium to the preceding phase. For instance, at point F at the same time water is equilibrium to gibbsite, goethite, kaolinite, hydromica, montmorillonite, gypsum, and a number of minerals not presented in the Figure. During the evoluttonary development, the layer of a new mineral phase appears. It causes the breaking of the available links, and mechanical isolation of initial components of the system, i.e. water, and primary rock, reacting with each other from the very beginning. To continue interaction within the initial system, new-water must arrive at the reaction front through the secondary mineral that was already developed and naturally react with it. As this new water is undersaturated both with respect to the primary and secondary rocks, at first they also dissolve this secondary mineral (for example calcite) and come to a balance with it. Then in accordance with the physical chemistry laws water solution
3 MECHANISMS OF GEOLOGICAL SELFREGULATION OF THE WATER-ROCK SYSTEM The formation of the secondary mineral phase leads the system to a principally different more complicated state called in synergy as “stationary”, being in the condition of a movable equilibrium. In the developing system, there can be several such stationary states successively alternating each other or existing simultaneously. Their number is unlimited and, in principle, can be infinite. Elements accumulate continuously in water solution.
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reaching the surface of primary rock forms the same secondary phase, the quantity of which continuously increases. So, system stabilization takes place at the level of secondary phases reproduction that gradually widens the area of its distribution in geological space. Due to this stabilizing mechanism laid down in the water-rock system, a huge mass of carbonates has developed on the Earth with the volume continuously growing up. The same is true with respect to all the other secondary mineral phases. The self-organization of the water-rock system is its unique property directed to preserve and multiply those new mineral phases that arise in the course of its evolution. In water-rock systems therefore, development directed to create new natural formations and their stabilization in geological space and time, some other important properties of the system manifest themselves. First, it is the system “memory” to reproduce the structural and self-like substance forms, which promotes preserving the primary organization of matter, its integrity and ability for self-development and spread in space. Second, this heredity hides the system mechanisms of regulating its relationships with the environment that, being transformed, promotes the growth of such mineral phase that on dissolution could influence the geochemical condition of the environment. Third, the system can attenuate or exclude random interactions, behave in a strictly prescribed mode, stabilize the evolution process at a certain stage of its development, and thus defend itself from external random events, which serves an indication of selforganization (Shvartsev 1998a). The above allows introduction of a notation of water-rock system geological self-regulation, i.e., its ability for evolving into new stationary states with typical spatial structures capable of selfdevelopment and dissemination with time and space. In this case the entire system retains its structuralfunctional parameters and can further develop in a complicated way, although it has already passed into an. energetically more profitable state (Ebeling 1985).
4 THE PROBLEM OF INFORMATION IN GEOLOGICAL SYSTEMS Information is the main chain in the teaching of selforganization of matter. Not a single system, if it is the self-organizing one, can function without information transmission. What is the mechanism of obtaining evaluating and using information? In this connection, it is necessary to remember that information in synergy comes out as a result of solving an alternative. The system evolutionary
development, its origin, and destruction are the information process fixed in the structure of formations arising in this case. Therefore, the structure of the state of the system is the basic and practically the only source of information in inanimate nature (Shcherbakov 1990). By modem theory, in liquid water, there is interaction among embryos of crystal transformations rather than disordered molecules intrinsic to any liquid phase of matter. These embryos represent conglomerates of “blinking clusters” consisting of molecules bonded via hydrogen links and floating in more or less free water (Frank and Wen 1957). By this hypothesis, though the clusters being identified with ice-similar carcasses, their structure is not identical to this of ice-tridymite. The authors believe that though the mechanism of uniting into clusters is unknown, it provides their diverse structure and the link with monomer water molecules. Most researchers believe that the nature of liquid water can be determined at my prescribed moment of time by the resemblance of ice-similar associates (clusters) to monomer water molecules (Eizenberg & Kauzmann 1969). This is the ratio of molecule groups to single molecules deprived of these links. This ratio depends on a number of factors among which the most important is the temperature and composition of the water solution. With changes in both of them, the structural features of water also change. It is important that in this case the evolution of the water-rock system leads to a selective concentration of chemical elements in the water solution, which in turn determines the directed change in water structure at the same temperature. Therefore, changes in water structure encode information serving as a signal for ions in the solution to change their behavior, to draw together, and settle. The conclusion follows from all the above indicates that the structure of water incorporates the mechanisms of using, evaluating, processing, and passing information. The multiplicity of structural elements and their variability under the influence of external factors promote the generation of new information fixed in the structural forms created later. Having analyzed the available factors, we propose that: the main code of information directing the evolutionary development of water-rock system is laid down into the structural features of water solution. This is due to capability of this water solution to multiple structural changes under a slight influence of chemical ions, temperature, and pressure, electric and magnetic fields, etc. The structural changes of water solution serve the bases for providing information, which directs the action
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Frank, H.S. & W.Y. Wen 1957. Structural aspects of ionsolvent interaction in aqueous solutions: a suggested picture of water structure. Discus. Faraday Soc. 24: 133-145. Eizenberg, D.S. & W. Kauzmann 1969. The structure and properties of water, New York: Oxford University Press. Prigogine, I. & I. Stengers 1984. Order out of chaos. Man‘s new dialog with nature, Toronto, New-York, London, Sydney: Banton books. Shcherbakov, A.S. 1990. Self-organization of matter in inorganic nature: philosophical aspects of synergy. (In Russian). Moscow: Moscow University Press. Shvartsev, S.L., Pinneker, E.V., Perelman, A.I., Kononov, V.I. & A.D. Nazarov 1982. Principles of hydrogeology. Hydrogeochemistry. (In Russian). Novosibirsk: Nauka. Shvartsev, S.L. 1995. Equilibrium-nonequilibrium state of the water-rock system. In Y.K.Kharaka & 0.V.Chudaev (eds), Proceedings of the 8‘” International Symposium on WaterRock Interaction: 75 1-754. Rotterdam: Balkema. Shvartsev, S.L. 1998a. Hydrogeochemistry of the zone of hypergenesis. (In Russian). Moscow: Nedra. Shvartsev, S.L. 1998b. Brines in Siberian Platform: Geochemical and isotopic evidence for water-rock interaction. In. G.B. Arehart & J.R. Hulson (eds), Proceedings of the 91hInternational Symposium on WaterRock Interaction: 357-360. Rotterdam: Balkema.
of the mechanism of selection of the ion interaction forms in the solution with the subsequent formation of particles and their “pushing O U ~ ” from the medium of origin (Shvartsev 1998a).
5 CONCLUSION 1. The water-rock system is in equilibriumnonequilibrium state in all points of the Earth crust that makes it internally contradictory, capable of spontaneous, continuous, geologically long evolution with the formation of principally new mineral phases and new geochemical types of water. 2. The water-rock interaction is of stage character: a strictly definite assemblage of secondary phases and a definite geochemical type of water correspond to each stage. The more prolonged and intensive is the water-rock interaction, the more secondary phases are formed and the greater is the difference between the composition of rock and water. 3 . In the course of the water-rock system evolution the amount of secondary mineral phase in close association with given geochemical type of water increases and involves more and more geological space. 4. The water-rock system possesses the ability to concentrate some chemical elements and to disperse other elements; this finally leads to their deep geochemical differentiation and generation of principally new mineral forms. 5. The water-rock system is self-organized because it is contradictory, both equilibrium and nonequilibrium in character, non-linear, has the capacity for spatial-temporary development, contains ability for assimilation, accumulation, evaluation and is able to pass on the information. The main code information directing the evolutionary development of the water-rock system is laid down into the structure of water solution. 6. The properties of self-organization and selfdevelopment determine the water-rock system as one of the fundamental and basic in the development of inorganic matter of the pre-biotic stage of the evolutionary setting of geological processes and structures. It was this system that gave birth to many others inheriting many of its features. REFERENCES Ashbey, W.R. 1962. Principles of self-organization. Principles of self-organization. : 3 14-343. Oxford: Pergamon Press. Ebeling, W. 1985. Thermodynamics of self-organization and evolution. Geochim. Cosmochim. Acta 44(6): 83 1-838.
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Isotopic and chemical characteristics of old “ice age” groundwater, North Iceland A.E.Sveinbjornsd6ttir & S .Arn6rsson Science Institute, University of Iceland. Dunhagi 3, I07 Reykjavik, Iceland
J .Heinemeier Institute of Physics and Astronomy, University of Aarhus, DK-8000, Denmark
ABSTRACT: It is suggested, on basis of chemistry and isotopic composition, that some of the natural groundwater in the Skagafjordur Valley, North Iceland, originates as precipitation under different climatic conditions from those prevailing at present. The 6D values of these samples are more negative than of those from any highland springs in the watershed and a negative correlation is observed between 6D and C1 of marine origin in the water. These waters range in temperature from about 13 to 90 “C and have a very distinct chemistry that reflects a maximum amount of water-rock interaction. They show the highest I4C age of all the waters in Skagafjordur which, after a boron based correction, ranges between 3000 and 70.000 BP, demonstrating mixing to a varying extent between “ice age water” and more recent precipitation. 1 INTRODUCTION 2 WATER CHEMISTRY AND ISOTOPES Chemical and isotopic data on natural waters in the Skagafjordur region, northern Iceland has been compiled over the last 3 years. This is the most comprehensive data set on groundwater from a particular area in Iceland with 253 water samples, where the water temperature ranges from ambient to about 90 “C. Major and trace elements have been analyzed together with oxygen, deuterium and carbon isotopes. Tri tium has also been measured in some of the samples. Earlier, deuterium isotopes of a few samples had been used to trace the origin of the thermal water of the Skagafjordur Valley into the interior of the country, where the isotopic composition of precipitation is simijar to that of the groundwater in the Valley (Arnason 1976). Interpretation of the new isotopic and chemical data contradicts the earlier interpretation and it is clear that most of the isotopically lightest water cannot be traced into the highlands, It has been suggested that this water originates from the last glaciation (AndrCsdcittir et al. 1998). At that time the hydrological conditions of the Valley allowed infiltration of seawater (which is not possible today) and the precipitation was isotopically lighter than at present due to colder climate. Hence the light geothermal water contains components of seawater and precipitation from that time. In this contribution we will focus on the groundwater that definitely has “ice age water” fingerprint.
The natural waters in the Skagafjordur Valley and in the highlands of the watershed have been divided into five groups on basis of their chemistry and geological occurence. Figure 1 shows that 6l80and
Figure 1. Relation between 6’”O and 6D in the Skagafjordur natural waters. Legends: plus(rivers), open boxes (lakes), open crosses (peat waters), triangles (highlands springs), open dots (groundwater in lowland).
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of the waters follow the meteoric line that has been defined for Iceland (Sveinbjornsd6ttir et al. 1995), apart from some lake water, that shows substantial evaporation, the peat waters that plot just below the line and the isotopically lightest groundwater samples, which show slight oxygen shift up to 2.5%0. It can also be seen from the figure that some of the groundwater within the Skagafjordur Valley is more depleted in deuterium than the highland springs and, therefore, there origin cannot been traced directly to the highlands. This is even more clear in Figure 2 where the relationship between 6D and water temperature is shown. The most depleted highland spring has a 6D value around -100%0 whereas fourteen of the groundwater samples are more depleted, the difference being up to 20%0.Their water temperature varies from 13 to 90 “C. One of the water samples (98-3059, Hofsvellir well-1) is exceptional in its chemical and isotopic composition. It has about 14%0 lower 6D value than any mean annual value for precipitation in Iceland at present and its chemical composition suggests much more extensive water-rock interaction than for the rest of the samples.
Figure 3. Relation between 6D and the C1 of marine origin in the water (dots). The “ice age” water is shown as circles.
ing to today’s local precipitation and the other to water relatively high in Cl of marine origin but depleted in 6D. 3 14CAGE O F GROUNDWATER Figure 4 shows the relationship between 6D and the 14 C apparent age of the Skagafjordur waters. Surface waters and cold springs show modern 14C values, but vary from -70%0 to -96%0 in their 6D values. A
Figure 2. Relation between 6 D and water temperature of groundwater (circles) and highland springs (dots).
The relationship between 6D and the C1 of marine origin in the water as evaluated using the method described by Arn6rsson and Andrksd6ttir (1995) is shown on Figure 3. It can be seen from the figure that the data points fall into two populations. One shows decreasing C1 with decreasing 6D (more negative values) and can be accounted for by altitude effect. The other population consists of the fourteen depleted samples and shows a negative correlation between these two parameters. That can be explained by assuming the geothermal water to consist of at least two components, one correspond-
Figure 4. Relation between 6D and the I4C apparent age of the water. Symbols have the same signature as in Figure 3.
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clear relationship is detected between 6D and I4C apparent age of the groundwater where the isotopic values of the groundwater decrease with increasing 14 C apparent age. The figure shows also that the waters that are believed to display an “ice age” fingerprint are the most depleted samples and those that give the oldest I4C apparent ages. However, an overlap in the 14Cages is observed in the age range 12000 to 18000 BP for the two data groups. It is noticeable that the most depleted 6D sample (Hofsvellir well-1) does not give the highest I4C apparent age. A general trend is observed between the 14C apparent age of the Skagafjordur water and water temperature as demonstrated in Figure 5, where the age increase with increasing water temperature, apart from the “ice age” waters that show no correlation with temperature and cover the entire temperature range.
boron based model overestimates the 14C dilution and gives too high a 14C concentration. Using the 613C values the adsorbed carbon in these waters is suggested to be derived from decaying organic matter. The boron based dilution model in Figure 6 suggests that the corrected I4C concentration approches zero for the exceptional sample (Hofsvellir well; rock derived carbon = 14.7 mmoledkg).
Figure 6. Correlation between the measured I4C concentration and the estimated carbon derived from rock. Symbols have the same signature as in Figure 3 .
Figure 5. Correlation between water temperature and the I4C apparent age of the water. Symbols have the same signature as in Figure 3 .
3.1 The Boron based correction model The results of Sveinbjornsd6ttir et al. (1995, 1998) suggest that the groundwater boron concentration can be used to estimate the fraction of rock derived carbon (I4C dead CO,) in the water samples. The correlation between the boron concentration and the 14C concentration in pMC (percent of modern carbon) is exponential, as expected if the boron concentration is a measure of dilution with I4C-dead CO, from the rock. Figure 6 shows the correlation between the measured 14C concentration and the estimated carbon concentration in the water that is leached from the rock. The curve fitting was performed only with samples that are considered by their 613C values to be unaffected by young organic carbon. For the samples that do not fit the curve the
There is a direct relationship between the amount of rock-derived Cl component in the water and the amount of water-rock interaction. Figure 7 shows the relation between the I4C activity (uncorrected) and the rock-derived C1 in the “ice age waters” (apart from the Hofsvellir well), as evaluated with the method described by Arn6rsson and AndrCsddttir (1995). Comparison of Figures 7a and 7b demonstrates the relative success of the boron based correction model. Figure 7a shows no clear relation between the two parameters whereas Figure 7b, which shows the I4C activity after the boron based correction, demonstrates that with increasing waterrock interaction (i.e. with increasing rock derived C1 in the water) the 14Cactivity decreases and hence the age of the water increases. Figure 8 suggests that with increasing age C1 of marine origin increases, demonstrating that with increasing I4C age the ratio of the “ice age water” component in the groundwater also increases.
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Figure 7. Correlation between the rock derived carbon and the a) measured I4C concentration and b) the I4C concentration after the boron based correction.
4 CONCLUSIONS
This observed age range reflects mixing to a varying extent of “ice age water” and more recent precipitation. It is also clear from this study that not all the natural hot waters with relatively heavy isotopic values and low I4C age can be traced to their origin on the basis of their deuterium content alone, as they may have a small component of the old glacial water. Hence, this work demonstrates the danger of using deuterium data of recent precipitation only to trace the origin of natural waters.
It is suggested, on basis of chemistry and isotopic composition, that some of the natural groundwaters in the S kagafjordur Valley originate as precipitation at different climatic conditions from those prevailing at present. Although these waters range in temperature from about 13 to 90 “C they have a very distinct chemistry that reflects a maximum amount of water-rock interaction. They show the highest I4C age of all the waters in Skagafjordur, which after boron based correction for dead carbon due to interaction with the host rock, range between 3000 to 70.000 BP.
ACKNOWLEDGEMENTS This project is financed by the Icelandic Research Council and the Danish Natural Science Research Council. REFERENCES
Figure 8. Correlation between marine derived C1 in the water and the I4C concentration after the boron based correction.
AndrCsdbttir, A., Arnbrsson, S. & Sveinbjornsdbttir, A.E. 1998. Geochemistry of natural waters in Skagafjordur, NIceland. I. Chemistry. In Arehart & Hulston (eds). WaterRock Interaction: 605-608. Rotterdam: Balkema. Arnbrsson S. & A. AndrCsdbttir. 1995. Distribution of B and C1 in natural waters in Iceland. Geochimica et Cosmochimica Acta, 59: 4125-4146 Arnason, B. 1976. Groundwater systems in Iceland traced by deuterium. Reykjavik, Societas Scientiarum Islandica 423236~. Sveinbjornsdbttir, A.E., Arnbrsson, S., Heinemeier, J. and Boaretto, E. 1998: Geochemistry of natural waters in Skagafjordur, N-Iceland. I1 Isotopes. In Arehart & Hulston (eds). Water-Rock Interaction: 653-656. Rotterdam: Balkema Sveinbjornsdbttir,A.E., J. Heinemeier & S. Arnbrsson. 1995. Origin of 14C in Icelandic Groundwater. Proc. ISh Radiocarbon Conference. G.T. Cook, D.D. Harkness, B.F. Millerand & E.M. Scott (eds). Radiocarbon 37: 551-565.
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Supercritical Water-Rock Interaction for Development of Deep-seated Geothermal Reservoirs N.Tsuchiya, N.Hirano, G.Bignal1 & K.Nakatsuka Department of Geoscience & Technology, Tohoku University, Sendai, Japan
ABSTRACT: Dissolution reactions of granite, and quartz, with pure water have been investigated under sub- and supercritical conditions in order to evaluate geothermal reservoir dynamics. The supercritical region, up to the critical point for water (- 374”C, 22 MPa), has been inferred to be a homogeneous state, which neither conforms to a true liquid phase, nor a true vapor phase. In terms of dissolution behavior of granite and quartz, the supercritical region may be subdivided into two apparent phases, comprising a ‘liquid-like’ region and a ‘vapor-like’ region. Solvent properties of supercritical fluid in the liquid-like region are more similar to those of subcritical water, whilst dissolution reactions of reservoir rocks in the vapor-like supercritical state are considered to be weak. The vapor-like supercritical region provides optimum conditions for development, and energy utilization, of an artificial HDR reservoir.
1 INTRODUCTION HDR/HWR (Hot Dry Rock / Hot Wet Rock) hydrothermal systems constitute a clean energy resource of great potential, since deep-seated geothermal reservoirs (DSGR) store huge amounts of renewable thermal energy. HDR projects have been undertaken in a number of areas in Japan, and in several other countries. Figure 1 shows representative temperature-depth conditions for some well-known HDR reservoirs, at subcritical thermodynamic conditions. Recently, the Kakkonda Granite was intersected by production drilling at the Kakkonda geothermal field, northeast Japan, with the rock temperature at approximately 3.5 km depth being up to 500°C (Doi, et al. 1998; Ikeuchi et al. 1998). Utilization of a subcritical geothermal reservoir (at >374”C and 22 MPa for pure water), underlying a conventionally developed geothermal reservoir, takes advantage of thermal energy transferred to supercritical water. High temperature rock masses, supporting a DSGR at supercritical conditions, are highly potential reservoir targets for development of HDR systems. Large volume neo-granitoids, compared to their associated volcanics, provide a suitable heat source to host a long-lived DSGR (Bando et al. 2001). In this paper, we describe supercritical water-rock interaction processes at sub- to supercritical conditions, for pure water, in hypothetical DSGR system in terms of the dissolution characteristics of granite and individual quartz crystals.
Figure 1. Temperature-depth conditions for Deep -Seated Geothermal Reservoir (DSGR) and HDR test sites. The broken line indicates typical temperature profile in the Kakkonda geothermal area modified and simplified after Dai et al. 1998.
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2 SCHEMATIC MODEL OF DSGR Our schematic representation of a magmatichydrothermal system is shown in Figure 2, with a convective hydrothermal system, overlying a deeper conductive region that may potentially host a supercritical DSGR. ‘Conventional’ geothermal reservoirs (A) may be tapped by wells, of <2.5-3km depth, sited to intersect faults and other permeable features (that supply fluids to surface features, B).
Multiple magmatic intrusions (C) are the 'heat source' for most 'high-temperature' geothermal systems, with the brittle-plastic transition in felsic intrusive rocks occurring at -370-400°C (at D). A few wells have penetrated the brittle-plastic transition, but these have little permeability (and/or produce gas-rich hypersaline brines), since interconnected pore spaces are typically closed with the onset of plastic conditions. The core (E) of the intrusion(s) is partly crystallized, and exsolves magmatic fluids, which accumulate near the top of the intrusion. Here, high fluid pressures may produce hydrothermal fracturing, expulsion of fluid across the brittle-plastic transition and brecciation of overlying rocks (Fournier 1999). Countering the process that increases permeability is deposition of quartz and other minerals in veins and breccia matrix, coincident with development of Cu-Au-MOporphyry systems (e.g. at F). Clearly, the downward circulation of a dilute fluid at hydrostatic fluid pressures, and corresponding
3 EXPERIMENTAL Hydrothermal dissolution experiments, using a batch type autoclave (300cm3), were conducted on synthesized tabular quartz ( 3 0 x 1 5 ~ 3 mm) and medium-grained cubic granite (1 cm'). The maximum temperature and pressure in our experiments was 600 "C and 60 MPa respectively. Reaction rate was determined during a long-term experiment, with reacted solution collected through a sampling tube and an equal volume of fresh solution injected into the system via a pre-heating system. Physical conditions in the autoclave, such as temperature and pressure, were maintained at a constant level throughout the experiment. Quantitative analysis for cation composition of the solution was carried out by ICP, and sample weight change and morphology were also examined.
4 RESULTS Figure 3 shows the variation in Si concentration for quartz dissolution in pure water at 30 MPa and 60 MPa under three temperature conditions (400"C, 500°C and 600°C). In the case of our 30 MP experiments (Fig. 3a), dissolution of quartz is inferred to have reached an equilibrium state within 24 hours. The greatest Si concentration was obtained in the 400°C experiment, and secondly for the 600°C experiment. In contrast, dissolution reactions for 400°C and 500°C experiments at 60 MPa (Fig. 3b) were slower than for the 600°C experiments, with the latter having the greatest Si concentration and the 400°C experiment having the lowest Si concentration.
Figure 2. Schematic model of a magmatic- hydrothermal system (described in the text), with a 'conventional' convective system at intermediate depths, and deep-seated, supercritical geothermal reservoir at greater (>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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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
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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
A
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individual mineral arains reaional litholoov short&, sedimentary mineral-water interface diffusion transport basins hydrated pore &fracture Openings faults and joints sedimentan,, deDositional environments . mineral inclusions intrusive rock bodies
.-
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fracture dilation extent of annual ground-water flow aquifer heterogeneity ecosystems to ecoregions
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mineralization & alteration faults stress regimes basin development gravity driven ground-water flow palhs
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thermally driven ground-water flow paths surface-water flow paths Fracture-controlled flow 4 multiple
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.
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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
<--(Tpl2ABp)
dT
RT
d2
dT2
(Tpl2AB)
(6)
Figure 1. Natural logarithm of Henry's constant (Ink,) of COz at saturated water pressure (P sat.) and 500 bar, plotted against reciprocal temperature. Solid curves are model predictions.
where "g)' at pedice stands for ideal gaseous component at T of interest and standard pressure, index "1" denotes pure H20, and "2" - the solvent. 2.2 Fitting procedure and the results To calculate any thermodynamic property of the dissolved species (Equations 3-6) one needs to know the corresponding property of the solvent. In this study the thermodynamic properties of pure water were calculated adopting the unified equation of Hill (1990). Data for ideal gaseous H20 are from Glushko (1 98 1). The Of V2 at state conditions (250c, bar) Of aqueous species together with all available sets of experimental data on kH were used to estimate the three empirical parameters 228
Figure 2. Partial molar volumes of aqueous COz at subcritical (left) and supercritical (right) conditions. Symbols correspond to experimental data. Solid curves are model predictions at infinite dilution.
3 DISCUSSION AND CONCLUSIONS The proposed equation of state (EoS) provides good prediction for the whole set of the investigated thermodynamic properties (chemical potential, entropy, molar volume and heat capacity) of aqueous species at infinite dilution in a wide T-P range encompassing the H20 critical region. In terms of Henry constant k~ the accuracy is better than 0.07 log10 units for nonpolar species, and 0.1 units for polar ones (H2S, as an example). This last discrepancy may be due to the fact that proposed model does not take into account the electrostatic interaction between solute and solvent molecules.
Figure 3. Apparent molar heat capacity of COz at subcritical (left) and supercritical (right) temperatures and 280 bar. Symbols localize experimental data at a CO2 molality approximately 0.192. Solid curves are model predictions at infinite dilution and T, P of interest.
proach for predicting thermodynamic properties of "large" organic dissolved molecules. And finally, application of our pseudo-virial equations (3-6) also allows prediction of concentration dependencies of the thermodynamic properties of aqueous species. As an example the comparison between experimental and calculated excess molar volumes of H20-CO2 system is given in Figure 7.
Figure 4. Temperature dependence of Ink,, for Ar at saturated water Pressure. Symbols Co~esPondto experimental data; the Solid Curve is the model extrapolation Performed stemming from the stardard state Properties of aqueous Ar.
The adopted empirical parameters could easily be estimated even in the case of lack of experimental data. Specifically, they could be retrieved using known values of ,LQ, S 2 , and V2 at 298 K and 1 bar. As an example, standard state thermodynamic properties of dissolved Ar from Shock et al. (1 989) lead to the values: 4 = 0.0733, a = -7.6895, b = 11.4657. These values when used to predict temperature dependencies of kH, V2, and Cp,2of aqueous Ar result in fair accuracy of the model extrapolations (Fig. 4-6). The proposed EoS can be assumed as a development of the semiempirical approach derived by Japas & Levelt Sengers (1989). In the framework of our these authors adopted the linear between and the temperature T ' .Here we used an approximation (2) that gives better linearity to the T-derivative of Ink!,. It is to be noted that equation (2) is somewhat close to the Redlich & Kwong (1949) equation where intermolecular parameter is set proportional to T-'.~. Usage the adjustable parameter 4 as a scaling volume factor allows to expand the proposed ap-
Figure 5. partial molar volumes of aqueous Ar at subcritical (left) and supercritical (right) conditions. Symbols denote experimental data; solid curves are the model extrapolation performed stemming from the standard state properties of aqueous Ar.
229
Figure 6. Apparent molar heat capacity of Ar at subcritical (left) and supercritical (right) conditions. Symbols denote experimental data, and solid curves are the model extrapolation performed stemming from the standard state properties of aqueous Ar.
Glushko, V.P. I98 1. Thermal constants of substances: Directory in ten issues. Issue 10, Part 1-2. Moscow, Academy of Sciences of the USSR. Hill, P.G. 1990. A unified fundamental equation for the thermodynamic properties of HZO..J. Phys. Chem. Re$ Data 19: 1233-1274. Japas, M.L. & J.M.H. Levelt Sengers 1989. Gas solubilities and Henry's law near the solvent's critical point. AfChE J. 35: 705-7 1 3. Mason, E.A. & T.H. Spurling 1970. The virial equation of state. In J.C Rowlinson (ed.), The International Encyclopedia of Physical chemistry and Chemical Physics. Vol. 2, topic 10. Thefluidstate. New-York. Plyasunov, A.V. 1991. The approximate computation of Henry's constants of non polar gases at water supercritical temperatures. Dokl. Ak. Nairk SSSR 32 1 : 107 1 - 1074 (in Russian). Plyasunov, A.V., O'Connell, J.p. R.H. Wood 2000. Infinite dilution partial molar properties of aqueous solutions of nonelectrolytes. 1. Equations for partial molar volumes at infinite dilution and standard thermodynamic functions of hydration of volatile nonelectrolytes over wide range conditions. Geochim. Cosmochim. Acta 64: 495-5 12. Redlich, 0. & J.N.S. Kwong 1949. On the thermodynamics of solutions. V. An equation of state. Fugacities of gaseous solutions. Chem. Rev. 44: 233-244. Shock, E.L., Helgeson, H C. & D.A. Sverjensky 1989. Calculation of the thermodynamic properties of aqueous species at high pressures and temperatures: Standard partial molal properties of inorganic neutral species. Geochim. Cosmochim. Acra 5 3 : 2157-2183. Tanger, J.C. & H.C. Helgeson 1988. Calculation of the thermodynamic and transport properties of aqueous species at high pressures and temperatures: Revised equations of state for standard partial molal properties of ions and electrolytes. Amer. J. Sci. 288: 19-98.
Figure 7. The excess volumes of CO2 - HzO mixing at 400°C and two different mole fractions of COz.
ACNOWLEDGEMENTS This research was supported by the Russian Foundation for Basic Research, Grant 01-05-64368 and Russian Ministry of Education. Author is greatly thankful to R. Wood, A.V. Plyasunov, and A.V. Zotov for fruitful discussions. Author is also very grateful to the anonymous referee for his comments and correction of the manuscript. REFERENCES Akinfiev, N.N.1997. Thermodynamic description of H,O-gas binary systems by means of Redlich-Kwong equation over a wide range of parameters of state. Geochem. Intern. 35(2): 188-196.
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Water-Rock lnteracfion 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Soultz granite - saline water interactions at 175-200°C and 10-50 bar: experimental and therrno-kinetic modeling approaches M .Azaroual BRGM, Water Division, av. C . Guillemin, BP 6009, F-45060 Orlkans, France
V.Plagnes LSCE, CNRS, av. de la terrasse, F-91198 GifYvette Cedex, France
I .Matsunaga National Institute for Resources and Enviroizment, 16-3 Onogawa, Ibaraki 305-8569 Tsukuba, Japan
ABSTRACT: Laboratory and in-situ (at depths between 2000 and 5000 m) water - granite interactions phenom-
ena were studied in the framework of the development of the future European Hot Dry Rock (HDR) geothermal exchanger at Soultz-sous-Forets (Alsace, France). The target reservoir consists of a fractured granite with high saline solutions (- 100 g/l, - 200°C). Experiments of saline water (1, 10, 25 g/l) - granite interactions were achieved in dynamic systems. On the other hand, thermo-kinetic modeling results were compared to the laboratory experimental observations. A good compromise was obtained assuming local chemical equilibrium between the percolating fluid and zeolites, smectites, illite, calcite and goethite. This agreement was obtained when measured N2-BET specific surface was reduced by 0.001 to 35%. This result confirms the systematic discrepancy between operational kinetic constants and that determined in very constrained and simplified systems. Hence, the difficulty of precisely determining water - minerals exchange surface area and the misunderstood surface reaction mechanisms limits greatly the predictive capacity of thermo-kinetic modeling software.
I INTRODUCTION Exploration and exploitation of energy resources (petroleum, geothermal reservoirs, etc.) and also the radioactive waste disposal require sufficiently accurate reactive transfer models to understand the evolution of these natural and industrial hydrogeological systems. The Hot Dry Rock (HDR) principle is based on deep fluid circulation through a network of natural and artificial fractures to extract geothermal energy. The performance and the lifetime of a HDR geothermal reservoir will be controlled essentially by the magnitude of water-rock interactions and hydromechanical constraints. In this paper, results of laboratory and modeling works on Soultz granite - water interaction with varying temperature (175 - 2OO0C),salinity (0 - 50 g/l) and pressure ( 10 - 50 bar) are presented. 2 EXPERIMENTAL WORKS
Over the two last decades, experimental dynamic systems have been developed to simulate geochemical mass transfer induced by fluid flow through various geological environments. Two types of dynamic high temperature experimental devices (at the
BRGM and the NIRE) were used to study waterrock interaction mechanisms in the framework of European HDR project being developed at Soultzsous-Forets (Alsace, France). The advantage of that system is largely defended in the literature (PoseyDowty et al., 1986). 2.1 The BRGM device A schematic description of the fluid flow equipment was presented in Azaroual & Fouillac (1997). All components of the system are made of the inert materials PEEK (Poly-Ether-Ether-Ketone) and Titanium. The flow rate can reach 10 ml/min. The titanium reaction cell (I.D. 50 mm, L. 120 mm) containing the rock grains is placed in a furnace with a reservoir where the fluid is stored after flowing through the reaction cell and before sampling. The storage reservoir is closed by a piston maintained by N, gas under experimental pressure conditions, keeping the fluid under stable temperature and pressure. The fluid reaches experimental temperature before entering the cell and can be cooled before sampling. A device allowing pH measurement and samples for silica analysis before degassing and cooling is also installed on line. Three saline water
231
Soultz granite experimental series were achieved. Table 1 gives salinity, flow rate and duration of each series.
Table 2. Experiment conditions achieved on the NIRE device at 200°C. The flow rate is 0.1 ml/min for all experiments. *T = 175°C for this experiment. Exp. No
Table 1. Conditions of experiments achieved on the BRGM device at 180°C. Exp. No Salinity Flow rate W/R Duration PS-400 PS-500 PS-600
(g/l) 0.14 25 50
(ml/min) 0.09 0.09 0.09
ratio 0.37 0.36 0.37
(days) I6 22 16
Sltz02" Sltz07 Sltz08 Sltzl2
Type of granite Fresh Gr. Fresh Gr. Altered Gr. Fresh Gr.
Salinity (g/l) 0 1 1
10
W/R ratio 0.43 0.43 0.48 0.42
Duration (days) 11 11 10 10
2.3 S o u k granite
Only fresh granite was used in the BRGM (BUreau de Recherche Gkologique et Minikre) experiments. Various proportions of Soultz brine (25 and 50 %) and diluted water from the distribution network of Soultz city (series PS-400) were used. The present day natural Soultz fluid is of Na-C1-CaHC03 type and highly saline (- 100 g/l) with temperature of 200 "C at 5000 m (in GPK-I). The in situ pH (5.2), pCO? (2.24 bar) and saturation indexes of calcite, anhydrite, barite and quartz were calculated using the SCALE2000 software (Azaroual et al., 2001) integrating the Pitzer (1995) approach. All these minerals are nearly in equilibrium with the Soultz brine. The detailed concentrations of major elements are (in molal): [Cl] = 1.65; [Na] = 1.15; [Ca] = 0.16; [K] = 0.08; [HCO?] = 0.01; [Li] = 0.02; [Mg] = 0.01.
The petrography of cored samples from GPKl and EPS I boreholes has been described in previous studies (Genter, 1989; LedGsert, 1993). The crushed fresh granite samples from GPK-1 (sample K-20, from a depth of 1995 m) with size between 1 and 2 mm was used in the BRGM and the NIRE experimental works. The fresh granite is a porphyritic Kfeldspar granite, rich in Na-perthite plagioclase and Fe-Mg biotite with minor components such amphiboles, sphene, apatite, zircon and allanite. The altered granite from EPS-l (sample K194, from a depth of 2165.05 - 2165.29 m) used in the NIRE experiments is quartz, K-feldspar and illite granite with traces of metal oxides, chlorite and calcite. 3 WATER-ROCK INTERACTION THERMOKINETIC MODELLING The thermo-kinetic software EQ3/6 (Wolery and Daveler, 1992) was used to simulate water - interactions under experimental conditions. The rate law integrated and modeling principles of geochemical reactions was detailed in Azaroual et al. (2001).
2.2 The NIRE device
In the NIRE (National Institute for Resources and Environment) device, the storage reservoir for injection fluid is closed by a piston maintained by Ar gas under experimental pressure conditions in order to keep the fluid under stable temperature and pressure and to prevent contacts with ambient atmosphere. The fluid circulates through the system using a HP pump (0.01 - 9.99 ml.min-'). The cell of percolation is a pure titanium cell (I.D. 11 mm, L. 100 mm). The high ratio of length-to-diameter (9.1) minimizes mechanical and diffusional dispersions. The fluid reaches the experimental temperature (<350"C) in the autoclave before percolating through the reactive cell. The temperature and pressure continuously recorded are constant inside the cell. Three type of fluids were used as initial solutions: distilled water (pH =5.8 at 25"C), low salinity fluid (pH = 5.84, TDS = 1 g/l, [Na] = 1.914, [Ca] = 3.129, [SO,] = 0.019, [Cl] = 8.135), high salinity fluids (pH = 5.9, TDS = 10 g/l, [Na] = 115.2, [K] = 7.9, [Ca] = 15.7, [Mg] = 0.58, [Cl] = 156.24). All concentrations are in mmol/l. 232
3.1 Modeling input constraints With thermodynamic data of aqueous and mineral species, kinetic parameters are the most important constraint on thermo-kinetic modeling. Table 3 presents the apparent kinetic constants of dissolution reactions of Soultz granite minerals. Table 3. Apparent kinetic constants for Soultz granite minerals (by Azaroual & Fouillac. 1997). Mineral Quartz K-Feld. Plagio. Biotite Illite Hematite Calcite
log k (25") -17.4a -1 5.5b - 16.5 -16.1 -17.1 -14.0 -15.0
Ea"
logk
logk
logk
(1 75°C) (180°C) (200°C)
21.3 13.8 8.4 13.0 16 10.0
-13.8 -16.0 -16 -12.9 -13.7 -11.5 -12.55
-13.1 -16.4 -15.5 -12.5 -13.6 -11.3 -12.5
-12.4 -13.6 -13.3 -12.3 -13.3 -11.0 -12.3
"Activation energy in kcal/mol; the units of kinetic constants are in mol/cm2/s.
Measured specific surface area using BET-N2 method for fresh and altered granite is 0.6 and 1.8 (in m2/g), respectively. The saturation index, pH, and concentrations of different elements in the outlet fluids were used as guidelines to accept results. Some observations of secondary minerals reveal precipitation of zeolites, aluminosilicates for experiments with pure water at 175, 180 and 200°C. For series with saline waters, calcite, illite, smectites and zeolites are observed at 180 and 200°C. Experiments with high salinity fluid at 200°C (Sltz12; 10 g/l) shows supersaturation vis-h-vis of illite and muscovite but the steady state was not achieved. Interestingly, experiments with portions of Soultz brine (Le., PS-400, PS-500, PS-600) show more “buffered systems” (figure 1A).
Figure 1. pH evolution of the outlet fluids for different experimental series. The pH is an important parameter to explore because it is very sensitive to both dissolution and precipitation reactions in the system. Figure 1A shows that a steady state was attained after 200 h of fresh granite water interactions at 180°C. Steady state was achieved (at 50 h) for low salinity experiments (Figure 1B) for both fresh granite (FGr) and altered granite (AGr) at 200°C. In contrast, for high salinity (10 g/l), the system is more reactive and unstable
(Figure 1B). Note that the modeling study using a quasi-stationary reaction path approach as in EQ3/6 is only possible in the hypothesis of steady state achievement of the dynamic system. 3.2 Modeling Results After several tests, the best agreement between measured, computed concentrations and pH was obtained with precipitation of calcite, hematite saponite, witherite and laumontite for experiments with water of Soultz city distribution network at 180°C (Fig. 2A). Anhydrite, huntite, saponite and magnetite were predicted to precipitate for high salinity fluid (25 g/l) at 180°C (Fig. 2B).
Figure 2. Simulation of the spatial distribution of precipitations along the path of the weathering fluid in the 12-cm percolation cell. Figure 3 shows the spatial repartition of thermodynamic affinities (TA) of primary mineral dissolution reactions along the path of the weathering fluid. Thermodynamic equilibrium (TA = 0) will be achieved more rapidly for systems with low salinities (Fig. 3A). Hence, the dissolution processes will be more important for low salinity experiments. For
233
the same time interval (- 24 h), the amounts of the silica and aluminum dissolved is more important for PS-400 (0.2 g/l) than for PS-500 (25 g/l) and PS-600 (50 gA). In contrast, ionic exchange reactions are certainly more important in the case of high salinity systems as demonstrated by high concentrations of Rb, Cs and Li. Figure 3B confirms the low reactivity of high salinity systems. TA is still far from equilibrium and demonstrates that the dissolution reaction kinetics are slower.
quasi-stationary state is very disturbed. It is preferable to use salted solutions whose chemical composition corresponds to fractions of the current brine of the system and not of major salts. 4 CONCLUSION Modeling thermo-kinetics of the mechanisms of water-rock interactions induced by the exploitation of geothermal energy in a context of Hot Dry Rock systems is not yet predictive because of the ignorance of the reactional mechanisms. Nevertheless, experimental work on the scale of the laboratory constitutes a significant support for better forcing (in an indirect way) the models. The water of the distribution network of Soultz is very reactive with respect to the granite. The use of certain specific salts to approach the salinity of the brine of Soultz is likely to amplify the phenomena of water-rock interaction and to prevent an operation in quasi-stationary state of the exchanger. It is preferable to use fractions of the brine mother of the tank even the brine itself if it is possible. From a thermodynamic and kinetic point of view, the order of the reactivity of minerals in the system is the following: plagioclase > K-feldspar > biotite > quartz > calcite. This order does not hold account for the fact that calcite is very reactive and can be treated using the local equilibrium approach.
REFERENCES
Figure 3. Simulation of the spatial distribution of thermodynamic affinities of primary minerals along the path of the weathering fluid in the 12-cm percolation cell. 3.3 Discussions These experiments of water - rock interaction show that modeling is not yet predictive. While using the measured kinetic and thermodynamic parameters, it ignorance of the reaction mechanisms makes it very difficult to use these parameters before confrontation with laboratory experiments. To obtain a satisfactory agreement, we were obliged to reduce the reactive surface of minerals from 0.001 to 35 %. Other share, the achievement of the quasi-stationary state by a dynamic system is very difficult to demonstrate. The use of some salts (Sltz12 series) shows that the
Azaroual, M., Ch. Hurtevent, Ch. Kervkvan, S. Brochot & M.-V. Durance 2001. Quantitative prediction of scale depositions induced by oil production: application of the thermo-kinetic software SCALE2000 (V2.01, Proceedings of SPE congress, N o 68303. Aberdeen 30-31 January, 200 1. Azaroual, M. & Ch. Fouillac 1997. Experimental study and modeling of granite-distilled water interactions at 180°C and 14 bars. App. Ceochenz. 12: 55-73. Genter A. 1989. GCothermie roches chaudes skches: le granite de Soultz-sous-For& (Bas-Rhin, France). Thise de doctorat, Universite' d'orle'ans, 201 pp. LedCsert B. 1993. Fracturation et palkocirculations hydrothermales. Application au granite de Soultz-sousFor& Thkse de 1'Universitk de Poitiers, 170 p. Pitzer K.S., 1995. Thermodynamics. 3'd edition, University of California, Berkley, pp. 626. Posey-Dowty, J., D. Crerar, R. Hellemann & D. Clarence 1986. Kinetics of mineral-water reactions: theory, design and application of circulating hydrothermal equipment. Am. Min. 7 1 : 85-94. Wolery, T.J. & S.A. Daveler 1992. EQ6, A computer program for geochemical aqueous speciation-solubility 246 pp. calculations (V7.0), UCRL-MA-110662-PT-I,
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Solubility and reaction rates of oxides and hydroxides to high temperatures with in situ pH measurement P.Bknkzeth, D.A.Palmer, D.J.Wesolowski & C.Xiao Chemical and Analytical Sciences division, Oak Ridge National Laboratory, P.O. Box 2008, Oak Ridge TN 37831 -6110
S .A.Wood Department of Geology, University of Idaho, Moscow ID 83844, U S .
ABSTRACT: Summaries of experimental studies involving the equilibrium solubility of the crystalline solid phases boehmite (AlOOH), zincite (ZnO), and neodymium hydroxide (Nd(OH)3), and the dissolutiodprecipitation rates of aluminum solid phases (boehmite and gibbsite, Al(OH)3) are presented. These experiments were performed over a wide range of pH (2-lO), temperature (30 to 290°C), and ionic strength (0.03-5.0), using a hydrogen-electrode concentration cell (HECC). This cell provides continuous, accurate in situ pH measurements of solid/solution mixtures to 295°C. Some of the potentiometric investigations were complemented by batch (AlOOH and ZnO), and flow-through column (ZnO) experiments. This research represents the first studies ever reported of mineral solubility profiles across the entire pH range of natural waters, and their precipitatioddissolution rates at temperatures above 1OO"C, involving direct pH monitoring. 1 INTRODUCTION Knowledge of the solubility of metal oxides and hydroxides in aqueous solutions under controlled conditions is of direct value in understanding natural and industrial processes, as well as yielding the thermodynamics of the metal-containing species in solution that are in equilibrium with these solid phases. Indeed, owing to the high degree of insolubility of most oxides in near neutral pH solutions, solubility measurements often provide the best means of identifying and quantifying these dissolved species. We recently measured the solubility of minerals such as boehmite (Benezeth et al. 1997, 2001; Palmer et al. 2001), zincite (Wesolowski et al. 1998, Benezeth et al. 1999), and neodymium hydroxide (Wood et al. 200 l), and the dissolutiodprecipitation rates of aluminum solid phases (boehmite and gibbsite) by taking advantage of recent advances in analytical techniques, the application of in situ pH monitoring, and experience with flow-through methods, or combinations of these approaches. 2 MATERIALS AND EXPERIMENTAL METHODS 2.1 Materials All solutions were prepared from reagent grade chemicals and distilled deionized water. Concenused to make up the desired experimental solutions. The method of synthesis and characterization of the
solid phases used has already been published (e.g., Benezeth et al. 1997, 1999, Wesolowski, 1992, Wood et al. 2001). In all cases, X-ray diffraction (XRD), Scanning Electron Microscopy (SEM), and the surface area (BET) were used to confirm the crystallinity and surface area of the solid phases taken before and after a number of solubility and kinetic experiments.
2.2 Experimentalprocedure The HEXC has been described in numerous publications (e.g., Benezeth et al. 1997, 1999, Palmer et al. 2001) and was used in this work for the solubility measurements and to study the kinetics of dissolutiodprecipitation of aluminum solid phases. The initial configuration of the cell in a typical experiment containing identical acidic solutions is as follows: H2 ,Pt
I HC1, NaCl I I HCl, NaCl, S I Pt, Hz Reference
Test
where S represents the solid phase of interest. In the case of ZnO and Nd(OH)3, HTr (Triflic acid) and NaTr (sodium trifluoromethanesulfonate) were used instead of HCVNaCl due to the strong complexation of zinc and neodymium with chloride. After equilibration at temperature, each electrode responds exactlv Nernstian to the half cell reaction. Since the cells share a common hydrogen hgacity, the potential between the electrodes is defined as: 235
AE = -(RT/F)ln(aH',,~/aH',f) + Elj
as were required in the previous batch investigations.
A solution of known pH, (-log[&] in molal units) is used in the reference cup and an identical large excess of inert salt is added to both cups, such that the activity coefficient differences and liquid junction potentials (Elj) between the solutions are minimized. Consequently, the pH, of the test solution can be very accurately determined (to 0.001 pH, units) and also monitored continuously over time frames of seconds in kinetic titrations, up to months in longterm solubility experiments. The convention used in our laboratory is that H" is not associated with the medium anions and any ion pairing is treated implicitly by the stoichiometric molal activity coefficient model employed. PH, can be readily converted to the activity scale by applying an appropriate activity coeEcient/ion pairing model. Solution samples can be withdrawn from this cell for analysis of metal content by ICP, AA (graphite hrnace and flame) and spectrophotometry (depending on the metal concentration). Also, acidic or basic titrant (at matching ionic strengths) can be metered independently into the test cell to change the pH, after attainment of equilibrium (within ca. 12 hours) so that the solubility profile can be mapped out.
Figure 1. Solubility profile for boehmite at 203.3"C in 0.03m ionic strength (NaC1).
The observed effect of increasing ionic strength (from 0.03 to 5 molal) is a shift in the solubility minimum to higher pH,, due to an increase in the stability field of A13' and minimization of the stability fields of the hydrolyzed species other than the alurninate anion. For example, in acidic solutions an increase in salinity from 0.03 to 1 molal increases boehmite solubility by 1.2 orders of magnitude. It appears likely that such conditions can be achieved either in soil and surface waters affected by acid rain and acid mine drainage, or by condensation of acidvolatile-rich steam in geothermal systems, or in hydrothermal ore deposits, etc.
3 METAL, HYDROXIDE/OXIDE SOLUBILITY AND RATES 3.1 Boehrnite solubility The solubility of pure synthetic boehmite has been measured in non-complexing solutions over a wide range of pH (2-lO), temperature (100 to 290°C,) and ionic strength (0.03-5 molal NaC1). As an example, Figure 1 shows results obtained at 203°C in 0.03 molal NaCl which allow a comparison to be made with two previous batch-studies of Bourcier et al. (1993) and Castet et al. (1993). The shape of the solubility curve is controlled by the relative stabilities of Al(OH)?-" species (x=O-4) as a hnction of pH, temperature, and ionic strength. As can be seen, at this low ionic strength we have obtained excellent agreement at high pH, where Al(0H)i is the dominant aluminum species in solution, with these authors. However, discrepancies between our fit (solid line) and the two previous works, as well as the predictions of Pokrovskii & Helgeson (1995), were found at the solubility minimum, due to the higher predicted stability of the Al(OH)3°(aq)species by these authors. In our experiments, equilibrium was approached from supersaturation at the solubility solubilities cannot be due to insufficient equilibration times. These differences likely stem from the use of direct pH measurements, which do not necessitate the use of pH buffers or mass and charge balance equations to determine the pH at high temperatures,
3.2 Zinc oxide solubility Hydrolysis constants of aqueous zinc species have been determined by solubility measurements of zincite performed at 0.03 molal ionic strength (NaTr) as a hnction of pH (5-1 1) and temperature (25-250°C), using the HECC. Separate experiments were also performed in alkaline solutions (0.03 molal NaOH) using batch and flow-through reactor techniques. Treatment of the experimental data obtained shows that only monomeric species are involved and that four species Zn2+,Zn(OH)', Zn(0H)z" and Zn(OH)3are needed to fit the data at each temperature investigated. The solubility quotient values obtained (Benezeth et al. 1999) are in good agreement with the previous work for Zn2' and Zn(OH)3-, (Khodakovagree markedly for the intermediate species. The discrepancies arise mainly because the previous studies were performed, 1) in alkaline to near-neutral media, preventing accurate evaluation of the thermodynamic properties of Zn(0H)' and Zn(OH)z", and 2) at temperatures ranging mostly from 100 to 350"C, and extrapolated to 25°C. However, we shown in this 236
study that precipitation of a metastable solid phase, Z~I(OH)~(,,, on ZnO surface, associated with precipitation from higher solubilities, occurs for temperatures higher than 200°C in near neutral medium. The formation of this solid phase was checked by various spectroscopic measurements (e.g., X P S , Raman). This discrepancy, at the minimum of the solubility, is still under investigation by measuring the solubility of ZnO from 150 to 300°C in NaOH and NH3/HTr solutions using a flow-through reactor technique.
3.4 Dissolution/precipitation rates of aluminum solid phases
3.3 Neoi@miumhydroxide solubility The solubility of Nd(OH)3 has been recently determined in NaTr solutions (0.03-1.0 m) from 30 to 290°C by in situ pH measurements W C C ) . This study permitted extraction of the solubility constants for Nd3' and Nd(OH)3" over a wide range of temperatures, and the intermediate species Nd(OH)2+ and Nd(0H); at 250 and 290°C. Nd3' ion (a good surrogate for 3+ actinides) is the predominant species over the majority of the pH range investigated, with solubility constants decreasing with increasing temperature (retrograde solubility). The extrapolated, infinite dilution Kd values proved to be in fair agreement (0.5 log units compared with the experimental uncertainty at 25°C of 0.3 log units) with some previous low temperature results, but much larger variations were also found. It is suggested that in some previous work the solid phase was contaminated with surface layers of less soluble Nd(OH)C03. Thus, the current study has provided an empirical equation that yields Qa values as a hnction of temperature to 300°C and ionic strength to 1 molal. Differentiation with respect to temperature provided a reliable enthalpy of dissolution of Nd(OH)3 because a wide temperature range was investigated. The AH? value for Nd(OH)3 cr derived therefrom (-1414 f 5 W-mol-') is in excellent agreement with the value reported by Merli et al. (1997) of -1415.5 k 1.4 kJ-mol-I.The solubility quotients of the intermediate hydrolyzed species, Nd(0H)g and Nd(OH)2', were determined only at 250 and 290°C where they become more important at low PHm, but still are dominant only over a narrow PHm range as was observed for the aluminum analogues. It was found that estimates of K,1 (Hass et al. 1995) were two orders of magnitude higher than measured in this study. Similarly, the solubility minimum defined by the Nd(OH)30specie is about three orders of magnitude lower than previously estimated (Hass et al. 1995). This represents the first such study of the solubility and speciation of a rare earth element at temperatures above 100°C with in situ pH measurement. The Nd(OH)3(,, solubility results are also in good agreement with the results reported in the literature for La and Gd (Deberdt et al. 1998).
The ability to perturb pH isothermally by addition of acidic or basic titrant opens the door for studies of the kinetics of dissolutiordprecipitation, even for relatively fast reactions. For pH-dependent dissolutiordprecipitation reactions, the pH becomes a completely sufficient indicator of reaction progress, which can be continuously and precisely monitored using this potentiometric method. The resolution of our instrumentation permits observation of pH changes of < 0.001 log units, or 0.25% in the concentration of H'. The degree of super- or undersaturation is solely a fUnction of how far the solution composition is perturbed by titrant addition from the equilibrium solubility curve. Preliminary results were obtained on the dissolutiodprecipitation rates of boehmite (from 100 to 290"C), and gibbsite (3050"C), in neutral to basic solutions at 0.1 molal ionic strength (NaCl media). The principle of "detailed balancing" (e.g., Rimstid & Barnes, 1980, Lasaga 1998) was applied under the assumption that the dissolution rate (Al(0H)i +- AlOO&,, + O H + H20, hiss) obeys a first-order dependence on hydroxide concentration: d[Al(OH)i]/dt = hi& moH-, where C stands for the ratio of total surface area of boehmite to the mass of solution (the activity of water is a secondary, but significant factor). Similarly, the precipitation rate (Al(0H)i -+ AlOO&,) + O H + H20, kprec)of boehrmte must obey a first-order rate law: d[Al(OH)i]/dt = kprecC M ~ ( O H The ) ~ - .rate of dissolution is first order in base, while the rate of precipitation is independent of pH. The results obtained thus far establish that the dissolutiordprecipitation rates of both boehmite and gibbsite can be studied accurately and expediently near equilibrium using the unique capabilities of the HECC. Boehmite proved to be an ideal phase in that its dissolutiordprecipitation kinetics permitted demonstration of reversible equilibrium solubilities at all temperatures studied and no other phases formed as a result of incongruent dissolution or precipitation from a supersaturated state. However, the dissolution and precipitation rates of boehmite were too fast to observe large changes (i.e., far-from equilibrium), so that the utility of this method is demonstrated better with gibbsite at lower temperatures as illustrated in Figure 2.
237
4 CONCLUSIONS The thermodynamic data of aqueous ions obtained in these studies, over a wide range of temperature and medium, can now be used to update geochemical reactiodtransport codes such as EQ3/6, SUPCRT, Chiller, etc., and allow better understanding and modeling of aqueous metal speciation, masstransport, and deposition in a number of industrial
Figure 2. Solubility profile for gibbsite at 50°C in 0.1 molal ionic strength, NaC1, from batch experiments (Wesolowski & Palmer, 1994) showing the paths of repetitive addltions of acidlc titrant in the HECC.
(fossil and nuclear powered steam generators) and natural systems (hydrothermaVgeotherma1 fluids). This approach has been applied to the study of heterogeneous reactions, including adsorption of H'/OH on oxide surfaces and is entirely amenable to studies the kinetics of aluminosilicate dissolutiodprecipitation, and the corresponding equilibrium solubilities, as well as to the rates and equilibrium constants of mineral transformation reactions. Acknowledgments T h s research was sponsored by EPRI, Inc. and by the Division of Chemical Sciences, Geosciences, and Biosciences, Ofiice of Basic Energy Sciences, U.S. Department of Energy, under contract DE-AC05-000R22725 with Oak Ridge National Laboratory, managed and operated by UT-Battelle, LLC.
REFERENCES Benezeth, P., Palmer, D.A. & D.J. Wesolowski 1997. The aqueous chemistry of aluminum. A new approach to high temperature solubility measurements. Geothermics 26 : 465-48 1. Benezeth, P., Palmer, D.A. & D.J. Wesolowsla 1999. The solubility of zinc oxide at 0.03m NaTr as a function of temperature with in situ pH measurement. Geochim. Cosmochim. Acta 63: 1571-1586. Benezeth. P., Palmer, D.A. & D.J. Wesolowsla,2001. Aqueous high temperature solubility studles. 11. The solubility of boehmite at 0.03 ionic strength as a function of temperature and pH as determined by "in situ" measurements. Geochim. Cosmochim. Acta (in press) Bourcier, W. L., Knauss, K. G. & K.J. Jackson 1993. Aluminum hydrolysis to 250°C from boehmite solubility measurements. Geochim. Cosmochim. Acta 57: 747-762. Castet, S., Dandurand, J. L., Schott, J. & R. Gout 1993. Boehrmte solubility and aqueous aluminum speciation in hydrothermal solutions (90-350°C). Experimental study and modeling. Geochim. Cosmochim. Acta 57: 4869-4884. Deberdt, S., Castet, S., Dandurand, J. L., Harrichoury, J.C. & I. Louiset 1998. Experimental study of La(OW3 and Gd(OH), solubilities (25 to l5O0C), and La-acetate complexing (25 to 80°C). Chem. Geol. 151: 349-372. Hass, J.R., Shock, E.L. & D.C. Sassani 1995. Rare earth ele-
238
ments in hydrothermal systems: Estimates of standard partial molal thermodynamic properties of aqueous complexes of the rare earth elements at high pressures and temperatures. Geochim. Cosmochim.Acta 59: 4329-4350 Khodakovskyi, I.L. & A. Yelkin 1975. Measurement of the solubility of zincite in aqueous NaOH at 100, 150, and 200°C. Geokhimiya 10: 1490-1498. Merli, L., Lambert, B. & J. Fuger 1997. Thermochemistry of lanthanum, neodynuum, samarium and americium trihydroxides and their relation to the correspondlng hydroxycarbonates. J. Nucl. Mater. 247: 172-176. Lasaga, A.C. 1998. Kinetic theory in the Earth Sciences. Holland H.D. (Ed.), Princeton New Jersey, 81 lpp. Palmer, D.A., Benezeth, P. & D.J. Wesolowsla 2001. Aqueous hgh temperature solubility studles. I. The solubility of boehmite as functions of ionic strength (to 5 molal, NaCI), temperature (100 - 290"C), and pH as determined by in situ measurements. Geochim. Cosmochim.Acta (in press) Pokrovslai, V. A. & H.C. Helgeson 1995. Thermodynamic properties of aqueous species and the solubilities of minerals at high pressures and temperatures: The system Al2O3H20-NaCI. Amer. J. Sci., 295: 1255-1342. Rimstid, J.D. & H.L. Barnes 1980. The lunetics of silica-water reactions. Geochim. Cosmochim.Acta 44: 1683-1699. Wesolowski, D.J. & D.A. Palmer 1994. Aluminum speciation and equilibria in aqueous solutions: V. Gibbsite solubility at 50°C and pH 3-9 in 0.1 Molal NaCl solutions (A general model for aluminum speciation; analytical methods). Geochim. Cosmochim.Acta 58: 2947-2969 Wesolowski, D.J., Benezeth, P. & D.A. Palmer 1998. ZnO solubility and Zn2+complexation by chloride and sulfate in acidlc solutions to 290°C with in-situ pH measurement. Geochim. Cosmochim.Acta 62 : 97 1-984. Wesolowsla, D.J. 1992. Aluminum speciation and equilibria in aqueous solution: Part I. The solubility of gibbsite in the system Na-K-C1-OH-Al(OH)4 from 0 to 100°C. Geochim. Cosmochim. Acta 56: 1065-1092. Wood, S.A, Palmer, D.A., Wesolowski, D.J. & P. Benezeth 2001. The aqueous geochemistry of the rare earth elements and yttrium: XI. The solubility of Nd(OW3 and hydrolysis of Nd3+from 30 to 290°C at saturated water vapor pressure with in situ pH measurement. Geochem. Soc. Spec. Pub. (in Press). Ziemniak, S.E., Jones M.E. & K.E.S. Combs 1992. Zinc(I1) oxide solubility and phase behavior in aqueous sodlum phosphate solutions at elevated temperatures. J. Solution Chem. 21: 1153-1176.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
New insight on the chemical control of aqueous aluminum. Application for modelling water-rock interactions G.Berger & J.P.Toutain Paul Sabatier University - CNRS, UMR 5563,38 rue des 36 Ponts, 31400 Toulouse, France
ABSTRACT: A1 contents 'versus pH trends in both experimental and natural hydrothermal solutions are compared in the range 50-120 "C to constrain the control mechanisms of aqueous Al. A single complex trend is observable for synthetic Al, whatever being the mineral phase (kaolinites, smectites, feldspars, zeolites) in equilibrium. This suggests a crystallochemical control of aqueous A1 by the coordination of A1 in the solids. A1-pH trends of natural solutions obviously differ from the previous trends but fit with current growth experiments of Fe-bearing chlorites. This observation suggests that numerical codes should take this behaviour into account in order to constrain aqueous AI contents for modelling diagenetic and hydrothermal reactions. 1 INTRODUCTION Aluminium is present in most natural solutions as a trace element. It is a fundamental element in aqueous geochemistry because it controls the reactivity of aluminosilicates and influences common dissolution-precipitation reactions in most hydrothermal and diagenetic environments. In natural solutions, however, A1 is poorly documented even if its knowledge should be necessary for modelling water-rock interactions. Geochemists commonly by-pass the problem by assuming the solubility of hydroxides and aluminosilicates as reliable data to constrain A1 concentrations in models. In this work, we examine the validity of such assumption by processing data of AI concentrations obtained from both experimental and natural solutions in a wide range of pH, focusing interest on A1 in organics-free fluids from hydrothermal environments. Experimental data from synthetic solutions selected for this work were obtained around 100 "C in order to be compared with those of natural solutions . 2 EXPERIMENTAL EXPLORATION In Figure 1 we report, as a function of pH, aluminium concentrations measured in experimental solutions equilibrated or at steady state with respect to the following different aluminosilicates : Kfeldspar (Berger et al. submitted; Beaufort et al.
2001), zeolites (Wilken & Barnes 1998), glasses (Leturcq et al. 1999), kaolinite (Devidal et al. 1996), clinochlore (Aja & Small 1999), montmorillonite (Beaufort et al. 2001). Data for feldspar and kaolinite are completed with unpublished results. All experiments were performed at temperatures ranging from 90 to 125 "C, most of them being at 100 "C. In some of these studies, secondary minerals were clearly indentified ;but in others they were not. For pH ranging from 1.5 to 6, a single Al-pH inverse correlation fit the data, despite the variety of the minerals studied. For highest pH values (pH>7), two well-identified and positive A1-pH correlation can be observed. Such a U-shaped A1-pH relationship is a classical feature of aluminosilicates. However, one must note that data corresponding to several different mineral phases fits a single (low pH) or double (high pH) A1-pH relationship. It may be suggested that data reflect the various chemical coordination of AI in the solids. The trend B on Figure 1 corresponds to minerals having tetracoordinated AI among total aluminium whereas trend C corresponds to almost exclusively hexacoordinated AI.
3 Al-pH WATERS
RELATIONSHIP
IN
NATURAL
To check the relationships between A1 contents and measured pH of natural waters, we selected data in the literature for groundwaters collected at temperatures higher than 50 "C. We focused 239
attention on the hydrothermal igneous environment and processed data from both hydrovolcanic and active tectonic systems. Water condensates from geothermal vapour were eliminated. In this study, we considered groundwaters from 8 active volcanoes or geothermal areas (Vulcano, Nevado del Ruiz, Nishiki, Valles Caldera, Taal, Campi Flegrei, Chinchon, Las Tres Virgines geothermal field) and 6 tectonic areas (S or CO,-bearing granitic waters from Pyrenees, Alps, French Massif Central, Portugal and Corsica, SO,+Cl waters from the Siping metamorphic areas in China) which displayed rising temperatures higher than 50 "C. Between pH 0.6 and pH 9.44, the data display a general decreasing trend from high Al contents (up to 813 ppm) to low A1 contents (about 5 ppb). Above pH 6, the range of variation for Al concentrations appears much higher than that for lower pH. A very few of the data seem unrealistic. Two from Sulfur Creek in Valles Caldera (Michard 1989) show pH values about one unit below the general trend, and a single value from Nevado del Ruiz (Sturchio et al. 1988) display Al contents two orders of magnitude higher than a mean value defined from the general trend. They might be analytical artifacts and will not be considered for discussion. Finally, two groups of samples display large ranges of Al contents for small ranges of pH values. The first group is the S-bearing waters from Pyrenees. They show large Al variations (from 16 to 116 ppb) at almost constant pH (8.56 9.16, Michard 1990). The Al level of these waters are non representative of initial contents because of
late equilibration with kaolinite at rising temperatures (Michard 1990). The other group consists of waters from Campi Flegrei (Valentino et al. 1999). Except for four values at low pH, Al contents display a wide range (0.06 - 1.51 ppm) for pH values between 6.6 and 7.3). Careful examination of the middle-pH data, however, shows that a general negative trend exists between Al and pH which is much more consistent than the general trend. 4 DISCUSSION Ideally, the pH values considered in this study should be in-situ values. This is the case for selected natural data. In the case of experiments, because values supplied are measured at 25 "C, one would have to calculate them at experimental temperatures. Such a calculation requires the knowledge of dissolved carbonates or other aqueous species having an acid-base behaviour, which are not documented in most of the publications used. Therefore, such calculations have not been processed. The trends displayed on Figure 1, however, seem to us sufficiently unambiguous. In a first attempt, because Al seems controlled mainly by pH rather than by the nature of the mineral phase, one would expect that natural data from igneous contexts fit this general trend. The observation of both Figures 1 and 2 demonstrates that it is not the case, especially under alkaline conditions.
Figure 1. Al contents in experimental solutions equilibrated or in steady-state with various aluminosilicates. Temperatures of experiments are between 90 and 120 "C.
240
Figure 2. A I content (ppm) versus field pH for natural waters having collection temperatures higher than 50 "C. Data corresponding to volcanic areas are from Vulcano, Italy (unpublished results), Nevado del Ruiz, Columbia (Sturchio et al. 1988, Giggenbach et al. 1990), Taal, Philippines (Delmelle et al. 1998), El Chinchon, Mexico (Taran et al. 1998), Campi Flegrei, Italy (Valentino et al. 1999), Valles Caldera, USA (Michard 1989, Goff et al. 1985). Data corresponding to tectonic areas are from Chavez geothermal area, Portugal (Aires-Barros et al. 1998), Siping metamorphic district, China (Wang & Shpeyzer 1993, French PyrtnCes (Michard 1990), French Massif central (Sanjuan et al. 1988).
Figure 3: A1 data for current Fe-bearing chlorites experiments compared with the field trend (shadow area) and above mentioned A, B, C curves.
by using the different hypothesis, presented here, about the Al control in natural fluids.
Recent unpublished results from growth experiments of Fe-bearing chlorites are displayed on Figure 3. Both hexa- and tetracoordinated sites were evoked to account for B and C trends displayed in Figure 1. Figure 3 suggests the involvement of a third possible mechanism being the uptake of aluminium in the hydroxide layer of chlorites. The most important consequence of this work concerns the water-rock interactions modelling. It seems obvious that kaolinite, feldspar or hydroxides are not adequate analogs for the Alcontrolling phase. We will compare applications
REFERENCES Aires-Barros, L., Marques J.L., Graca R.C., Matias M.J., Weijden C., Kreulen R. & H. Eggenkamp 1998. Hot and cold CO,-rich mineral waters in Chaves geothermal area (Northern Portugal). Geothermics 27 : 89-107. Aja, S.U. & J. Small 1999. The solubility of low-Fe clinochlore between 25 and 175°C and Pv= PH,O. Eur. J. Miner. 11 ; 829-842. Beaufort, D., Berger G., Lacharpagne J.C. & A. Meunier 2001. An experimental alteration of montmorillonite into a
241
di + trioctahedral smectite assemblage at 100 and 200°C. Clay Minerals, in press Berger, G., Beaufort, D. & J.C. Lacharpagne 2001. Experimental dissolution of sanidine under hydrothermal conditions: mechanism and rate. Amer J. Sci., submitted. Delmelle, P., Kusakabe, M., Bemard, A., Fischer, T., De Brouwer, S. & E. Del Mondo 1998. Geochemical and isotopic evidence for sea-water contamination of the hydrothermal system of Taal volcano, Luzon, the Philippines. Bull. Volcanol. 59 : 562-576. Devidal, J.L., Dandurand, J.L. & R. Gout 1996. Gibbs free energy of formation of kaolinite from solubility measurement in basic solution between 60 and 170°C. Geochim. Cosmochim.Acta 60 : 553-564. Giggenbach, W.F., Garcia, N., Londono, A., Rodriguez, V., Rojas, N. & M.L. Calvache 1990. The chemistry of fumarolic vapor and thermal-spring discharges from the Nevado del Ruiz volcanic-magmatic hydrothermal system, Colombia. J. Volcanol. Geoth. Res. 42 : 13-39. Goff, F., Gardner, J., Vidale, R. & R. Charles 1985. Geochemistry and isotopes of fluids from Sulphur Springs, Valles Caldera, New Mexico. J. Volcanol. Geoth. Res. 23 : 273-297. Leturcq, G., Berger, G., Advocat, T. & E. Vernaz 1999. Initial and long-term dissolution rates of aluminosilicate glasses enriched in Ti, Zr and Nd. Chem. Geol. 160 : 3962. Michard, A. 1989. Rare earth elements systematics in hydrothermal fluids. Geochim. Cosmochim. Acta 53 : 745750. Michard, G. 1990. Behaviour of major elements and some trace elements (Li, Rb, Cs, Sr, Fe, Mn, W, F) in deep hot waters from granitic areas. Chem. Geol. 89 : 117-134. Sanjuan, B., Michard, A. & G. Michard 1988. Influence of the temperature of CO,-rich springs on their aluminium and rare-earth element contents. Chem. Geol. 68 : 57-67. Sturchio, N.C., Williams, S.N., Garcia, N.P., & A.C. Londono 1988. The hydrothermal system of Nevado del Ruiz volcano, Colombia. Bull. Volcanol. 50 : 399-412. Taran, Y., Fischjer, T.P., Pokrovsky, B., Sano, Y., AuroraArmienta, M. & J.L. Macias 1998. Geochemistry of the volcano-hydrothermal system of El Chinchon volcano, Chiapas, Mexico. Bull. Volcanol. 59 : 436-449. Valentino, G.M., Cortecci, G., Franco, E., & D. Stanzione 1999. Chemical and isotopic compositions of minerals and waters from the Campi Flegrei volcanic system, Naples, Italy. J. Volcanol. Geoth. Res. 91 : 329-344. Wang, Y. & Shpeyzer 1997. Genesis of thermal groundwaters from Siping’san district, China. Geochim. Cosmochim. Acta. 12 : 437-445. Wilken, R.T. & H.L. Barnes 1998. Solubility and stability of zeolites in aqueous solutions: I. Analcime, Na-, and Kclinoptilolite. Amer. Mineral. 83 : 746-761.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Testing a clay/porewater interaction model through a laboratory experiment Ph.Blanc, E.Gaucher, B .Sanjuan, C.Crouzet & A.Seron BRGM, OrlLans, France L.Griffault ANDRA, Chdtenay-Malabry,France
ABSTRACT: This paper presents an experiment performed to test a geochemical model of the porewater composition of the Opalinus Clay formation. Batch experiments were conducted on crushed argillite in contact with a synthetic solution, using a mixed CO2 + N2 gas phase to fix the PCO2. Small amounts of the suspension were periodically taken from the experimental cell. Analyses of the solution and of the solid fraction reveal that (i) most of the equilibrium between the argillite and the solution is achieved after about three weeks, and (ii) the solution composition, pH and Eh are all in close agreement with the model calculations, with the exception of adsorbed cation distribution on the clay fraction. However, the experiment suggests ways to improve the modeling, by including other minerals such as illite among the controlling phases. 1 INTRODUCTION
2 MATERIALS AND METHODS
This study takes place within the framework of the International Mont Terri Project which focuses on the scientific investigation of argillaceous rocks of low-permeability, as potential host rocks for nuclear waste. The formation under investigation, the Opalinus Clay, is well-characterized by studies related to various mineralogical, geochemical or geotechnical aspects (Thury & Bossart 1999). Recently, a geochemical model has been proposed for interpreting the mechanisms regulating porewater chemistry in the Opalinus Clay environment (Sanjuan & Gaucher 2001). Models for predicting the solution compositions in argillite are generally based on analytical data measured in the rock and the porewater (Bradbury & Baeyens 1998). The scope of the present work is to validate such geochemical modeling by experimentally equilibrating the rock with the solution predicted by the model. The present experimental work is based on the results of the modeling. To assess the accuracy of the model, a batch experiment is performed based on the equilibration of crushed argillite in contact with a synthetic solution. Solution composition corresponds to that calculated by the model. Only the solid to solution ratio is decreased, allowing to collect many samples of the solution in sufficient amounts to perform analyses on both the solid and the solution. Little or no changes in the composition of the solution should be observed if the model is correct.
2.1 Opalinus Clay samples Opalinus Clay is a compact clayey rock whose mineralogical composition is given in Table 1. The samples are taken from drilling cores stored in aluminium coated plastic bags. The bags were themselves placed into a glove box in an N2 atmosphere and all grinding operations took place inside the glove box in order to avoid any oxidation. The mineralogical relative proportions are obtained by combining different analytical techniques: XRD, TEM, DTA, Xray fluorescency and Cation Exchange Capacity (CEC) determination. The illite/smectite phase is found to be of the short-ordered type R1 with 30% smectite layers and the chlorite may incorporate about 5% smectite layers. The CEC determination gives a value of 16.3 k 1.5 meq/lOOg. Table 1. Mineralogical composition of the Opalinus Clay. YO Minerals YO Minerals Kaolinite 28 Calcite 14 Mite* 25 Dolomite 2 Illite/Smectite 13 Siderite 1 Chlorite 7 K-Feldspar 2 Quartz 6 Pyrite 2 * Weighted proportions of illite and muscovite could not be distinguished.
2.2 Starting solution The synthetic solution is prepared from the model solution whose composition is displayed in Table 2.
243
standard (quartz and calcite) or (ii) by comparing the areas of the 001 peak (kaolinite, illite, illite/smectite and chlorite). The determination of the exact proportions would require more investigations (Calvert et al. 1989). The accuracy of this simple method is about f 10 % against f 5 % for a complete analysis in the most favorable case (Reynolds 1989).
Minor or trace elements Fe, Al, Si are not added because equilibrium with the solution is quickly achieved by a weak dissolution of the argillite. This is taken into account by recalculating the speciation of the model solution, leading to a reference solution without Fe, Si or A1 (see Table 2). Table 2. Composition of the starting solution (in molkgw). Model Reference Analyzed Salts (gkgw) solution solution experimental solution Na 2.03 10-I 2.03 10-’ 2.05 10-I NaCl 10.36 K 1.68 10-3 1.68 10’3 1.71 10-3 KCl 0.13 CaCl, 1.82 Ca 1.64 10-’ 1.64 10’* 1.51 10, MgCI2.6H20 2.66 Mg 1.31 10-2 1.31 10-* 1.28 10-, C1 2.38 10-I 2.38 10-I 2.44 10-1 1.75 so4 1.23 10-’ 1.23 10-, 1.28 10-* Na2S04 -Si 9.37 io-’ -Fe 1.84 104 -AI 1.32 10-* Alk 1.65 10” 1.28 10‘3 1.27 10-3 NaHC03 0.11 7.10 7.11 pH 7.23
Table 3. Analyses performed on the solution Method K, Na, Ca, Mg Capillary Ion Analyzer Ionic Chromatography C1, SO4 Alk Titration Si Spectrophotometry Fe Spectrophotometry confirmed by ICPMS
3 ABRIDGED DESCRIPTION OF THE MODEL A brief description of the model is given below. A more detailed discussion provided by Sanjuan & Gaucher (2001). The composition of the porewater in contact with the argillite is accurately described by a mixed model in a closed environment, incorporating a preliminary step of mixing seawater and freshwater, cation exchange reactions between clay phases and mineral dissolution-precipitation. A seawater is diluted with a slightly ionized water, which, as a first approximation, was selected as being a pure water. The dilution factor is inferred from the C1 concentration of the porewater; it has a value of 2.33. With the exception of the chloride ion, PC02, so4 and pH, each element concentration is controlled by a mineral/solution equilibrium, listed in Table 4. It must be noted that all controlling-phases have been observed in the argillite either by XRD or TEM analysis.
2.3 Experimental procedure The equilibration experiment is conducted in a 500 ml polyurethane vessel, containing both the argillite and the solution. The vessel is enclosed in a glass reactor and its atmosphere is maintained at 99.63 % N2 and 0.37 % CO2 using micro-gates. The gas mixture is pre-moisturized by flowing through a first reactor containing only deionized water. The whole apparatus is thermostated at 25°C and kept under pressure since the entering flow rate is fixed at 215 ml/min whereas the exiting flow rate is maintained at 5 ml/min. The goal is to efficiently fix the CO2 partial pressure at 0.37 % CO2 without 0 2 which corresponds to our estimation of in situ conditions. This device is also used in order to avoid evaporation of the solution, which occurred in a previous experiment. 12 ml of the mixture shale + solution are withdrawn from the vessel at different times : 1 hour, 1 day, 1 week, 3 weeks, 6 weeks, 10 weeks and 12 weeks. After centrifugation, the solid is separated from the solution, washed three times with deionized water and then dried at 40°C during for about 15 hours. The solution is then filtered at 0.1pm before being analyzed.
Table 4. Model depicting the porewater composition. Parameter Controlling reaction Na, K Argillite/porewater exchange Ca Equilibrium with calcite, CaC03 Equilibrium with disordered dolomite, MgC03 Mg Mobile element C1, SO4 Si Equilibrium with quartz, Si02 Fe Equilibrium with siderite, FeC03 AI Equilibrium with kaolinite, Si2A1205(0H)4 Eh SO, - Pyrite FeS, couple
4 RESULTS
2.4 Analytical tools Table 3 lists the different types of analyses which are performed on the solution. The cation exchange capacity (CEC) is obtained after saturating the argillite surface with cobaltihexamine molecules, as determined by analyzing the extracted solution. This procedure is detailed by Gaucher et al. (2001). The XRD analyses are carried out on the bulk fraction and on the clay-size fraction (< 2 pm). The weighted percentages of the main phases are given (i) by comparing the areas of the main peak with an external
4.1 Solution compositions Solution compositions are displayed in Figures 1 and 2 with respect to equilibration time. It can be seen that after one week, C1, SO4, Mg and Na concentrations come very close to the calculated values. Calcium concentrations appear slightly lower than in the model, essentially because of a slight calcium discrepancy from the model in the starting solution (Table 2). 244
Figure 2. Eh and pH variations in the solution. Straight lines refer to the calculated values.
Figure 1. Composition of the equilibrated solution. Straight lines refer to the calculated values.
Significant departure from the calculations are observed for the total alkalinity and the potassium concentration. Eh and pH evolution indicates that predicted values are reached after 3 weeks. Silica concentration reaches 1.0 104 mol/l after one day and exhibits thereafter only weak fluctuations around this value. Iron concentration increases with the equilibration time with a final value of 2.5 1o - mol/l. ~
Figure 3. Mineralogical composition of the argillite.
Table 5. Cationic occupancy of the argillite (in meq/lOOg). Oh* 1 h 1 d** lw' 3 w 6 w l o w 12w NH4 2.1 1.7 1.3 1.0 1.3 1.7 1.5 1.7 K 2.0 1.2 1.4 1.4 1.5 1.9 1.7 1.5 Ca 4.5 8.1 7.9 10.1 10.2 7.9 7.2 7.0 Na 7.4 7.1 5.8 3.1 2.4 3.1 3.8 4.9 Mg 0.8 1.1 1.0 1.1 1.1 1.0 0.8 0.7 Total 16.7 19.2 17.3 16.7 16.6 15.6 14.9 15.8 * h stands for hour ; ** d stands for day ; w stands for week.
4.2 Solidfraction
'
Very few or no variation can be quantitatively detected in the argillite as shown in figure 3. The method followed for the semi-quantitative determination could be questioned, like many methods based only on the XRD analysis (Reynolds 1989). The variations of the cationic occupancy of the argillite are given in table 5 . The CEC is obtained by summing the contribution of all the cations. It can be observed that the CEC is almost constant with respect to the error bar on the initial value of 1.5 meq/100g. Most of all, the Ca/Na ratio is found to increase during the first three weeks. Thereafter, this ratio decreases toward the original value, which is not yet reached even after twelve weeks.
5 DISCUSSION For most of the parameters investigated, equilibration is achieved after three weeks. The values for Eh, pH and Na, Ca, Mg, Si, C1 and SO4 are in good agreement with the calculated data. This confirms the major influence of the exchange reactions on the Na composition and the control of calcium and magnesium by carbonates, calcite and disordered dolomite respectively. As expected, quartz is controlling the dissolved silica amount in the solution. Eh and pH values are in accordance with the calculations, 245
which implies for Eh that the couple SO4 - pyrite actually controls the redox system. It has been seen that total alkaliniy departs slightly from the calculated value (2.2 10- mol/l instead of 1.7 10” mol/l). In the model, the alkalinity is controlled by the equilibrium between atmospheric and dissolved CO2 and by limited dissolution of carbonate phases. Nevertheless, CoudrainRibstein & Gouze (1993) have suggested that the CO2 partial pressure could be buffered in some natural systems by mineral assemblages involving dolomite, kaolinite, chalcedony, chlorite and calcite. This buffering may not be strictly effective in our system since the PC02 is externally fixed at 0.0037 atm. whereas the mineral assemblage buffering gives a PC02 of about 0.01 atm. at 30°C (Coudrain-Ribstein & Gouze 1993). Unfortunately, chlorites exhibit large composition domains, complex compositiodstability relations and few relevant data are available so far concerning their stability in aqueous media (Walshe 1986). It is difficult therefore to take these phases into account in simulations. Similar questions arise fkom the potassium amount in the solution which starts at 1.7 10-3 mol/l, increases up to 5.5 10-3 mol/l after six weeks and then remains constant. Table 5 indicates that no potassium is released in the solution through exchange. Only illite or muscovite dissolution could explain such an increase. As for chlorite, relevant data describing the stability in solution are missing for 2:l clay minerals. For example, Vieillard (2000), in an extended review, indicates that no calorimetric data are available for the AG”f of any of these phases. Furthermore, the different methods for calculating the equilibrium constant yield an uncertainty of rf: 2 to k 3 log units, which strongly limits the use of such phases in modeling. A strong discrepancy with the model result is observed for the iron concentration (1.8 10-4mol/l versus 2.5 10-6 mol/l in the experiment). It appears clearly that siderite can no longer be claimed to control this element in solution. Illite and chlorite have been analyzed by TEM and these minerals exhibit strong amounts of structural iron. Other phases could be mentioned such as iron oxi-hydroxides. For example, replacing siderite by oethite would bring the iron concentration to 5.0 1015 mol/l. However it has not been found in the TEM observed argillite samples. Concerning the adsorbed species, the discrepancy between the initial and the final Ca/Na ratio seems to decrease with time. Analysis of the solid fraction equilibrated over a longer period of time will determinate if the ratio actually decreases toward the equilibrium with the solution.
6 CONCLUSION Evidence is provided that a geochemical model describing the composition of a solution in equilibrium with an argillaceous rock can be assessed by performing experimentally the equilibration between the rock and the predicted solution. Also, the results emphasize the ability of the experimental apparatus to control the PC02 and to maintain reduced conditions in both the atmosphere and the solution. The model developed by Sanjuan and Gaucher (2001) is then found to accurately describe the porewater composition of the Opalinus Clay formation. Improvement of the model would require relevant data concerning the stability of chlorites and 2:l clay minerals which, so far, are lacking. REFERENCES Bradbury, M.H. & B. Baeyens 1998. A physico-chemical characterization and geochemical modeling approach for determining porewater chemistries in argillaceous rocks. Geochim. et Cosmochim. Acta, 62: 783-795. Calvert, C.S., Palkowsky, D.A. & D.R. Pevear 1989. A combined X-ray powder diffraction and chemical method for the quantitative mineral analysis of geologic samples. In CMS Workshop Lectures, 1 , Quantitative Mineral Analysis of Clays Pevear D.R. & Mumpton F.A. (eds): 154-166. Evergreen, Colorado. Coudrain-Ribstein, A. & Ph. Gouze 1993. Quantitative study of processes in the Dogger aquifer, Paris Basin, France. Appl. Geochem. 8: 495-506. Gaucher, E.C., Cailleau, A. & L. Griffault 2001. Specific determination of ion exchange constants: the Opalinus Clay example. In G. B. Arehart (ed.) Proc. Intern. Symp. Water Rock Interaction-I 0 Italy. Rotterdam: Balkema (accepted). Reynolds, R.C. 1989. Principles and techniques of quantitative mineral analysis of clay minerals by X-ray powder diffraction. In CMS Workshop Lectures, 1, Quantitative Mineral Analysis of Clays Pevear D.R. & Mumpton F.A. (eds): 436. Evergreen, Colorado. Sanjuan, B. & E.C. Gaucher 2001. Analyses of argillite porewaters and geochemical modeling: examples of Mol (Belgium) and Mont Terri (Switzerland). Chem. Geol. (submitted). Thury, M. & P. Bossart 1999. The Mont Terri rock laboratory, a new international research project in a Mesozoic shale formation, in Switzerland. Eng. Geol. 52: 347-359. Vieillard, Ph. 2000. A new method for the prediction of Gibbs free energies of formation of hydrated clay minerals based on the electronegativity scale. Clays & Clay Minerals. 48: 459-473. Walshe, J.L. 1986. A Six-Component Chlorite Solid Solution Model and the Conditions of Chlorite Formation in Hydrothennal and Geothermal Systems. Econ. Geol. 81: 681703.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Dissolution of synthetic zeolites at low temperature - preliminary results J C a m a , X .Querol & C .Ayora Institute of Earth Sciences "Jaume Almera ", CSIC, Barcelona, Catalonia
E.Sanz & J.Ganor Ben-Gurion University of the Negev, Beer-Sheva, Israel
ABSTRACT: Dissolution experiments of synthetic NaP 1 zeolite were conducted using flow-through reactors at pH 3 and 25" and 50°C. The sample was synthesized from fly ash and contains significant amounts of other phases. Therefore, the release rates of the different elements depend on the dissolution rate of both the zeolite and the fast-dissolving accessory phases. After an initial period of non-stoichiometric release of aluminum and silicon, the release AvSi ratio approaches a constant value of 0.6, i.e., the stoichiometric Al/Si ratio of NaP1. Several hundreds of hours later, the AI / Si ratio begins to decrease with time, and the change in concentration is dominated by dissolution of accessory phases. 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. the present communication is to show that it is possible to extract reliable dissolution rates of zeolite from flow-through experiments of such material. Kinetics of zeolite dissolution at low temperature (
1 INTRODUCTION In the present communication we discuss preliminary dissolution experiments of synthetic NaP 1 zeolite conducted using flow-through experiments at pH of about 3 at 25" and 50°C. The zeolitic material was synthesized from fly ash produced in a coal-fired power station in NE Spain. These zeolites are potential candidates for use as sorbents in the treatment of acid mine waters due to their high retention capacity (20-30 mg of NH4' g-'). Knowledge of the stability of these zeolites under chemical conditions similar to those of the acid mine water is essential in order to model their effect on the chemistry of the mine water. The ultimate goal of the research is to investigate the dissolution kinetics and the solubility of these zeolites by carrying out flow-through and batch experiments. In particular, we study the rate dependencies on pH, the degree of saturation with respect to zeolite dissolution reaction, the Al inhibitory effect, and the catalytic effect of oxalic acid. The kinetic parameters and solubility data of the zeolitic material will be eventually used for computational modeling of several waste-involved scenarios using the reactive transport code RETRASO (Saaltink et al. 1998). Zeolitic material that is synthesized from fly ash is not pure and contains significant amounts of accessory phases. The release rates of the different elements during dissolution of such samples depend on the dissolution rate of both the zeolite and the fast dissolving accessory phases. The major objective of
2 MATERIALS AND METHODS 2.1 Raw material The zeolite material used in the present study is NaPl synthesized from fly ash by conventional and microwave-assisted hydrothermal alkaline activation experiments (Querol et al., 1997). The raw NaPl zeolitic material contains a significant amount of other phases. Based on X-ray diffraction patterns, the sample contains NaP 1 (Na6(Si10Al6032).12H20), quartz (SiOz), mullite (Al6Si2013), calcite (CaC03), tobermorite (Ca5Si6016(OH)2) and magnetite (Fe304). Presumably, along with these crystalline phases, amorphous Si and Al phases, as well as dried NaOH (used as a reagent in the production of the zeolite from the fly ash) are also present. The specific surface area was measured by the BET method, using 5-point N2 adsorption isotherms. Sample degassing lasted 24 h at approximate? 140°C. The specific surface area was 18.1 m2 g- (* 5%). The raw material was sieved to determine its size fraction distribution. Approximately 8, 20 and 32 wt.% 247
of the grains were coarser than 149, 53 and 25 pm, respectively. The remaining 40% corresponds to particles whose size is less than 25 pm. 2.2 Experimental setting and analysis Experiments were carried out using Lexan flowthrough reactors (ca. 35 ml in volume) hlly immersed in water-baths at 25 and 50 (* 0.2) "C. The flow rate is 0.05 or 0.1 mT, min-', yielding residence times within the cell of about 12 and 6 h, respectively. The steady state dissolution rate in a flow-through experiment can be calculated based on mass balance, using the equation:
where R, (mol m-2 s-') is the dissolution rate, vJ is the stoichiometry coefficient of component j in the dissolution reaction, 9 is the flow rate (L s-I), A is the total surface area (m2), and C,,,,, and Cl,Inpare the concentrations o f j (mol L-'), in the output and input solutions, respectively. The mass of zeolite in the cell was significantly decreased during the experiment as a result of its dissolution. The dissolved mass at each experiment was calculated daily based on the outflux of aluminum and silicon, assuming stoichiometric dissolution, and was subtracted from the initial mass. The daily total surface area was calculated by multiplying the mass by the initial (BET) specific surface area. Total input and output Si and Al concentrations were analyzed colorimetrically with a UV-visible spectrophotometer (uncertainty = rt 3%). The pH was measured at experimental temperature (uncertainty = f 0.02 pH units). 3 RESULTS
Figure 1. Dissolution of raw NaP 1 sample at 25 "C. Initial mass is 0.47 g; the input pH is 3. Variation of (a) Si and AI output concentrations (pIvQand output pH and (b) Aloutput / Sioutput ratio and accumulated loss of mass (%) versus time. The shaded area corresponds to a stirring stage, which is not discussed in the present communication.
From 400 to 1000 h (C), A I concentration continued to increase but in a slower rate, and Si concentration increased at a faster rate than the Al. Both Al and Si approached a maximum at 1000 h when the Al/Si ratio approached the stoichiometric ratio of 0.6. Thereafter (D), both the output Si and A I concentrations decreased, but the Al/Si ratio remained stoichiometric for another 400 h. During the remaining part of the non-stirred stage (E), Si concentration decreased significantly faster than Al, and as a result M/Si ratio decreased below the stoichiometric ratio. Throughout the experiment the output pH was significantly higher than the input pH. This difference in pH is explained by a combined effect of dissolution of the dried NaOH, and proton consumption due to the dissolution of the zeolite. The high output pH during stage (A) is probably solely controlled by the dissolution of the dried NaOH. The sharp decrease in pH during this stage and the consequent slow decrease of pH during stages (B) and (C) reflect the disappearance of the NaOH. During stages (B) and (C) the decrease in pH is halted due to in-
AND DISCUSSION
3.1 Dissolution of raw Nap1 at 25 "C The variation of output pH and Al and Si concentrations in a representative flow-through experiment as a hnction of time is shown in Figure 1. Although the input pH was 3 throughout the experiment, the initial output pH was 7.6 and it decreased with time towards the input value. Based on the Al and Si concentrations and the Al/Si ratio, we subdivided the non-stirred part of experiment to five stages (Fig. 1). During the first 100 h (A) the pH decreased sharply to about 4.5, Al output concentration was less than 10 pM, and Si concentration ranged between 200 and 300pM. During the next 300 h (B), Al concentration increased sharply whereas the pH continued to decrease and Si concentration remained constant. As a result, the Al/Si ratio increased to about 0.8. 248
crease in proton consumption as zeolite dissolved. Thereafter, the change in pH can be explained solely by this proton consumption. Due to the relatively high pH during stage (A), the solution is oversaturated with respect to several amorphous and crystalline aluminurn oxide, hydroxide and silicate phases (e.g., gibbsite and kaolinite). The source of silicon during this stage may be mainly from dissolution of amorphous silica. Although the solution may be undersaturated with respect to NaP1, its dissolution rate is expected to be very slow under pH>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
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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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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-
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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 <2-mm fraction. 2.2 Isotope dilution using enriched stable isotopes The soil sample (4g) was placed in a polyethylene (PE) test tube with 20mL 0.01 M Ca(N03)2 containing the spikes (enriched isotopes: 14Cd,53Cr, %U, 97Mo, 62Ni, 207Pb,203Tl,@Zn). The test tube was tumbled for 2 4 h on a rotary shaker. The suspended matter was removed from the suspension by centrifuging and filtration (0.45 pm, cellulose acetate). The liquid phase was acidified with HNO3. This solution was used for isotope ratio measurements (see 2.6).
the elements under investigation were added as neutral solutions prepared from nitrate salts. The test tube was tumbled for 24 h on a rotary shaker. The suspended matter was removed from suspension by centrifuging and filtration (0.45 pm, cellulose acetate). 4.9 mL filtrate was mixed with 0.1 mL Rh solution (1 mg/L) and 5 mL 0.3 M HN03. The resulting concentrations were analysed by HR-ICPMS (see 2.6). 2.6 Concentration and isotope ratio measurement Concentrations and isotope ratios were measured using an HR-ICP-MS instrument (ELEMENT, Finnigan MAT). MO, Cd, and T1, were analysed in low-resolution mode; Cr, Zn, Cu, Ni, and Pb in medium-resolution mode. Both external and internal calibration methods were used. Rh was used for the internal calibration. For ID analysis, the following isotope ratios were used to calculate the interchangeable amounts: 97Mo/98Mo,206Pb/207Pb, 111c%:l4cddO . 2 620 3 ~ ~ 2 0 5 5 ~ 21~, r / 5 3 ~ r66zn/68zn, , 63Cu/ Cu, NI/Ni. 2.7 Calculations The interchangeable fraction of an element in the soil sample is calculated by a common isotope dilution equation. Theory and equations are given in Gabler et al. (1999).
2.3 Recovery experiments 3 RESULTS
The soil sample (4g) was placed in a PE test tube with 20 mL 0.01 M Ca(N03)2 containing the elements under investigation. The elements were added as neutral solutions prepared from nitrate salts. The test tube was tumbled for 24 h on a rotary shaker. Spikes of enriched stable isotopes were added and the test tube was tumbled again for 24 h. The suspensions were then handled like the isotope dilution suspensions (see 2.2).
3.1 Recovery of added amounts by IDMS Recovery experiments are carried out to determine whether added and equilibrated amounts of the element of interest can be recovered by IDMS. For this purpose, these elements were added to each of five soil samples in batch experiments and tumbled until equilibrium was reached (24 h). After the end of this equilibration period, a spike of enriched stable isotopes was added, the resulting isotope ratio was analysed, and the interchangeable element fraction was calculated. If the added element remains interchangeable and is not bound irreversibly by the soil, it should be recoverable for IDMS analysis. Figure 1 shows the results of the recovery experiments for the elements Zn, Cu, TI, Pb, Cd, and Ni. The analysed concentrations represent the interchangeable element fractions determined by IDMS. These concentrations are shown in Figure 1 to be equal to the sum of SO (determined by IDMS) and the added concentrations. This clearly demonstrates that added amounts of the elements remain interchangeable and can be analysed by IDMS.
2.4 EDTA extractions The soil sample (4g) was placed in a PE test tube with 40mL 0.025 M EDTA. The test tube was tumbled for 24 h on a rotary shaker. The suspended matter was removed by centrifuging and filtration (0.45 pm, cellulose acetate). One mL of filtrate was mixed with 8.7 mL 0.15 M HN03, 0.1 mL Rh solution (1 mg/L) and 0.2 mL H202 (30 %). The mixture was placed in a UV digester for 30 min. to destroy dissolved organic matter and analysed by high-resolution inductively coupled plasma mass spectrometry (HR-ICP-MS; see 2.6). 2.5 Sorption experiments The soil sample (8 g) was placed in a PE test tube with 40 mL 0.01 M Ca(N03)2.Various amounts of 268
the element in the sorption equilibrium is in solution. Because IDMS analyses an isotope ratio in this solution, this method is very sensitive to contamination after the phase separation. Even small amounts of contamination at this stage of the analysis shift the isotope ratio towards the natural isotope ratio and result in strongly elevated SO values. Figure 2 shows that the completely different approaches of EDTA and IDMS give comparable results for the elements Cu, Cd, Ni, and Pb. Thus, both methods can be used for SO determination of these elements. For MO and T1, IDMS yields higher concentrations than EDTA. This can be explained by the fact that MO and T1 do not form stable EDTA complexes (Umland et al. 1971), which are essential for the EDTA technique. So for these elements, the IDMS approach should be preferred.
Figure 1. Recovery experiments. The dotted line indicates complete recovery.
3.2 IDMS and EDTA The results of the EDTA extractions and the IDMS determinations are given in Figure 2. The two techniques are comparable for the elements Cu, Cd, Ni, Pb and with some limitations for Zn. The two methods yield different results for T1, MO, and Cr. 3.3 Tl sorption isotherms
T1 sorption isotherms were determined for eight soil samples for which the EDTA and IDMS methods yielded considerably different SOvalues. The sorbed concentrations were calculated from the difference between the added Tl plus SO and the amount remaining in solution after the end of each equilibration experiment. SO was calculated from both the EDTA and IDMS results. Thus, two sets of sorption isotherms (Fig. 3) were obtained for Tl.
4 DISCUSSION 4.1 Comparison of IDMS and EDTA
IDMS and EDTA are two completely different basic approaches for determining SOin soil samples. The theory of the EDTA approach is that SOis transferred completely from the solid to the liquid phase by EDTA, a strong complexing agent. The element concentration (=SO) is then determined in the resulting EDTA solution. In the IDMS approach, the behavior of a spike in the sorption equilibrium is analyzed by determining the isotope ratio in the resulting extraction solution. It is assumed that SO has completely exchanged with the added spike so that at the end of the experiment the resulting isotope ratios at the sorption sites and in the corresponding solution are the same. The IDMS approach is independent of the position of the sorption equilibrium. The sorption equilibrium is normally shifted strongly to the side of the sorbed species. This means that only a small proportion of 269
Figure 2. Comparison of EDTA and IDMS. The dotted line indicates equivalence of EDTA and IDMS.
The difference between the results of the EDTA and IDMS analyses for Cr is difficult to explain, because a trend cannot be observed. Problems may occur with both methods. Cr(II1) forms kinetically stable complexes with humic substances, which remove the Cr from isotopic exchange reactions (Marx & Heumann 1999). Thus, the added Cr spike in the IDMS experiment may be irreversibly removed from the solution by these complexes and thus cannot exchange with the sorbed Cr. The formation of the
Cr(II1)IEDTA complex is very slow and usually carried out in boiling solutions (Umland et al. 1971). Hence, complex formation at ambient temperatures may be expected to be too slow for a application of the EDTA method. The EDTNIDMS plot for Zn (Fig. 2) shows that most of the samples plot near the equivalence line, but some samples plot significantly below this line, indicating the IDMS value is substantially higher than the EDTA value. The samples that do not plot near the equivalence line have high pH values and thus very low Zn concentrations in the solutions from which the isotope ratios are analysed. This indicates that after the phase separation step contamination with Zn, which is ubiquitous, shifted the isotope ratio towards the natural ratio, causing elevated IDMS results. Hence, for elements that are strongly sorbed and are susceptible to contamination, the IDMS method must be applied very carefully. The IDMS method is generally easier to use than the EDTA method, because the solutions obtained by the IDMS method can be analysed more easily by ICP-MS due to lower matrix concentrations. These solutions can be acidified directly for the ICP-MS determination and digestion of dissolved organic matter is not necessary.
4.2 Tl isotherms Sorption is described widely by the use of the nonlinear Freundlich isotherm (e.g., Filius et al. 1998, Boekhold et al. 1999): S=kCM
or
logS=logk+MlogC
(1)
where S is the sorbed element concentration, C is the
Figure 3. T1 sorption isotherms calculated on the basis of different approaches for determining So.
element concentration in solution and k and M are Freundlich parameters. As previously mentioned, the correct determination of SO is essential for determining isotherms, especially at low concentrations. Due to the low stability of the TVEDTA complex, the EDTA technique SO, leading to significant underestimates underestimation of the sorbed amounts S and thus
results in erroneous Freundlich isotherms at low concentrations (Fig. 3). If T1 sorption isotherms are calculated on the basis of IDMS results for SO (Fig. 3), the isotherms are linear in the log-log plot as postulated by the Freundlich equation (1). This demonstrates that the IDMS technique is a valuable alternative for determining of fractions of sorbed elements SOin soils.
5 CONCLUSIONS Enriched stable isotopes can be applied to determine fractions of sorbed elements in soils. For Zn, Cu, Cd, Ni and Pb, this technique yields results comparable to those for the conventionally used EDTA technique. For elements for which the EDTA method fails, the IDMS method provides a valuable alternative. The IDMS technique represents a completely different basic approach to conventionally applied extraction methods, is easy to handle and applicable for many elements, but requires a mass spectrometer. REFERENCES Boekhold, A.E., Temminghoff, E.J.M. & S.E.A.T.M. Van der Zee 1993. Influence of electrolyte composition and pH on cadmium sorption by an acid sandy soil. J. Soil Sci. 448596. Filius, A., Streck, T. & J. Richter 1998. Cadmium sorption and desorption in limed topsoils as influenced by pH: Isotherms and simulated leaching. J. Environ. Qual. 27: 12-18. Gabler, H.-E., Bahr, A. & B. Mieke 1999. Determination of the interchangeable heavy-metal fraction in soils by isotope dilution mass spectrometry. Fresenius J. Anal. Chem. 365:409-414. Heumann, K.G. 1988. Isotope dilution mass spectrometry. In F. Adams, R. Gijbels & R. Van Grieken (eds), Inorganic mass spectromefry: 301-376. New York: Wiley. Marx, G. & K.G. Heumann 1999. Mass spectrometric investigations of the kinetic stabiliy of chromium and copper complexes with humic substances by isotopelabelling experiments. Fresenius J. Anal. Chem 364:489494. Springob, G. & J. Bottcher 1998. Parameterization and regionalization of Cd sorption characteristics of sandy soils. I. Freundlich type parameters. Z. Pflanzenernahr. Bodenk. 161: 681-687. Umland, F., Janssen, A., Thierig, D. & G. Wunsch 1971. Theorie und prakiische Anwendung von Komplexbildnern. FrankfurVMain: Akademische Verlagsgesellschaft. Young, S.D., Tye, A., Carstensen, A., Resende, L. & N. Crout 2000. Methods for determining labile cadmium and zinc in soil. Eur. J. Soil Sci. 5 1:129-136.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
To stir or not to stir - implications for silicate dissolution experiments J .Ganor Ben-Gurion University ofthe Negev, Beer-Sheva, Israel
V.Metz Insrirure for Nuclear Wasre Management, Research Center Karlsruhe (FZK-INE),Karlsruhe, Germany
ABSTRACT: Experiments show that kaolinite dissolution rate increases with increasing stirring speed. The stirring effect is reversible, i.e., as the stirring slows down the dissolution rate decreases. We propose that under stirred conditions, ultra-fine kaolinite particles are formed as a result of spalling-off or abrasion of the kaolinite. These ultra-fine particles have a very high ratio of reactive surface area to total surface area and therefore, the dissolution rate increases with stirring. Balance between the formation and dissolution of the fine kaolinite particles controls the change in reactive surface area with time. Therefore, the stirring enhancement factor decreases as kaolinite dissolution rate increases. As a result, kinetic factors obtained under stirred conditions will be smaller than those obtained under non-stirred conditions. We recommend the standard use of non-stirred reaction vessels for kinetic experiments of mineral dissolution and precipitation that are slow enough to be surface-controlled. According to our calculations, experiments with fast-dissolving minerals, under condition in which the reaction is transport-controlled, will not be influenced by the stirring effect described in the present study. Such reactions will be better studied using stirred experiments.
1 INTRODUCTION
2 EFFECT OF STIRRING SPEED ON KAOLINITE DISSOLUTION RATE
Weathering rates of silicates observed in the laboratory are in general up to three orders of magnitude higher than those inferred from field studies. An important difference between mineral weathering in the field and dissolution in laboratory experiments is that natural systems are usually non-agitated whereas it is common practice in many experimental systems to stir or agitate the reacting mineral. It is widely accepted that under low-temperature conditions dissolution reactions of most silicates are limited by surface-controlled processes, rather than by the transport rate of products and reactants to and from the reacting surface sites (Berner 1978, Aagaard & Helgeson 1982, Murphy et al. 1989). Therefore, it is assumed that the method and rate of stirring do not influence the dissolution rate of most silicate minerals. However, several recent studies showed that agitation does affect dissolution rates of silicates (Amrhein & Suarez 1992, van Grinsven & van Riemsdijk 1992, Furrer et al. 1993, Komadel et al. 1998, Ferrow et al. 1999, Metz & Ganor submitted).
Flow-through experiments carried out at 25", 50" and 70°C and pH 2 to 4 show an enhancement of kaolinite dissolution rate with increasing stirring speed (e.g., Fig. 1). The effect of stirring speed on dissolution rate depends on temperature and pH (Metz & Ganor, submitted). The stirring effect is reversible, i.e., as the stirring speed decreases the dissolution rate slows down. Figure 2 illustrates the reversibility of the stirring effect. During the first stage of this experiment (stage A) the sample was not stirred; subsequently (stage B) stirring speed was set to 650 rpm, while at the end of the experiment (stage C) stirring was stopped again. The aluminum and silicon concentration increased as a result of stirring and returned to near the original concentrations when stirring was stopped. Accordingly, the dissolution rate increased from 1.0*10-'2mol m-2sP1(stage A) to 2.2.10-'2 mol m-2 s-' (stage B) and decreased back to 9.4-10-13mol mV2s-l (stage C), which is within error the same dissolution rate as that of stage A.
271
P
Metz & Ganor (submitted) proposed that under stirred conditions, ultra-fine kaolinite particles (grain
'f
(1)
=Ratef
where P is the production rate of the ultra-fine particles (mol s-'), sf(m2) is their surface area and Ratef (mol m-2 s-l) is their dissolution rate. The steadystate dissolution rate under stirred conditions is therefore Ratef
- sf
-I-Rateb S b +
Ratestked = Sf +Sb
where Rateb (mol m-2 s-') is the dissolution rate of the large kaolinite grains, which is equal to the dissolution rate under non-stirred conditions and Sb (m') is their surface area. In order to compare the stirring effect on reaction rate under different experimental conditions, the stirring enhancement factor, SEF, is defined as the relative difference between dissolution rate under stirred conditions and dissolution rate under non-stirred conditions:
Figure 1. Effect of stirring_speed on kaolinite dissolution rate at . pH=3 and 25°C.
SEF
E
Ratestirred - Ratenon-stirred
(31 Metz & Ganor (submitted) showed that the production rate of fine particles depends only on stirring intensity and is independent of temperature and pH. Therefore, at steady-state the amount of ultra-fine grains as well as their surface area is inversely proportional to the steady-state dissolution rate, and the stirring effect decreases with increasing dissolution rate. As will be shown below this dependence of the stirring effect on dissolution rate leads to errors in estimation of kinetic factors that are obtained under stirred conditions. -stirred
3 IMPLICATIONS FOR DETERMINATION OF KINETIC FACTORS BASED ON DISSOLUTION EXPERIMENTS
Figure 2. The change in Si concentration with time in a multistage flow-through experiment. The figure illustrates the reversibility of the stirring effect (see text). The solid line is a fitting of the model proposed by Metz & Ganor (submitted) to the experimental data of stage B (stirred) and C (non-stirred), respectively. The regression coefficient of the fitting, R2, is 0.94.
In order to demonstrate possible implications of the mechanism of the stirring effect that was proposed by Metz & Ganor (submitted) on determination of kinetic factors we model the results of simulated stirred and non-stirred dissolution experiments. The non-stirred dissolution rate is calculated assuming that the far-from-equilibrium rate law of the dissolution reaction can be written as:
size <
where ko, kl and k2 are rate coefficients, E, is the activation energy of the overall reaction, R is the gas constant, T is the temperature (K), a H + is the activity of H', nH+ is the order of the reaction with respect to H+ and CA/ is the aluminum concentration. To simu272
late the non-stirred dissolution rate (Rateb) under a specified set of environmental conditions (i.e., T, CA[, and pH) we took an arbitrary set of values of ) 1) and coefficients (i.e., ko, kl, k2, E, and n ~ +(Table substituted them into equation (4). In the following discussion we will refer to these simulated rates as the "realsim"rates and to the arbitrary values as the "reaIarbttcoefficients. We assume that the dissolution rate of the ultra-fine grains that were produced as a result of stirring (Ratef) is two orders of magnitude higher than that for the large grains (Helgeson et al. 1984),and calculate it using the equation
Ratej = 100 Rateb *
(5)
The surface area of the ultra-fine particles at steadystate was calculated by substituting Ratej and the production rate, P, from Table 1, into equation (1). The simulated stirred dissolution rate under the same environmental conditions was subsequently calculated by substituting sb (from Table 1) and the simulated values of Rateb, Rate4 sh into equation (2). We simulate different sets of stirred dissolution experiments. In each set, two environmental conditions were the same in all the simulations while the third one varied between the simulations. From each set we calculated the expected value of one of the coefficients (k2, E, and n ~ + as ) , is commonly done in interpreting dissolution experiments. Results of some of the calculations are shown in Table 1 and Figure 3. Figure 3a shows Arrhenius plots of simulated experiments at pH 3 and 1.5. Under stirred conditions the slope of the Arrhenius plot is shallower than the "real,i," slope and as a result the obtained apparent activation energy is less than the "realart," activation energy. The dissolution rate is faster at pH 1.5 than at pH 3 and therefore the slope and the obtained activation energy are closer to the "real" ones. Similar results are obtained when other kinetic effects such as pH dependence and A1 inhibition are examined (Figures 3b-c and Table 1). When the dissolution rate is slow the obtained kinetic effects under stirred conditions are significantly lower than the "realarbt' effects. At fast dissolution rate the difference between the stirred and non-stirred dissolution rates is less significant and the obtained values of the coefficient approach the "realarbf'values (Table 1). Metz & Ganor (submitted) measured apparent activation energies for non-stirred experiments and for experiments stirred at 650 rpm (Fig. 4). As was predicted above, the apparent activation energy for stirred experiments (8.5 k0.4 kcal mol-l) is significantly lower than that for non-stirred experiments (12.2kl.l kcal mol-I). The results of our study show that kinetic factors obtained under stirred conditions are smaller than those obtained under non-stirred conditions, and are better approximations for the real kinetic factors, at
Table 1. Comparison between kinetic parameters obtained in simulated stirred experiments and "realZ; values that were used as the input of the calculations. The ko and kl were not fitted in the modeling of the results. Their "realarb''values are ko=0.3 mol m-'s-' and kl=l. Set*
p H CA,
T
(pM) ("C)
P (m') (moVs)
sb
Ea k2 nH+ (kcal) (M-') 15 105 0.50 13.8 12.1 7.6 4.7 0.25 0.42 0.50 0.50
"realarb" 10-" T 0.5 40 8 T 1.5 40 8 T 3 40 8 T 4 40 8 8 PH 10 25 10 50 8 PH 10 100 8 PH 10 150 8 PH A1 2 25 8 1o4 * The sets are named according to the variable that varies between the different simulations in the set.
Figure 3. Results of simulated stirred and non-stirred dissolution experiments. (a) simulations examining temperature effect conducted at pH of 1.5 and 3 and constant A1 concentration of 40pM; (b) simulations examining pH effect conducted at constant temperature of 25°C and A1 concentration of 10pM; (c) simulations examining A1 inhibition conducted at constant temperature of 25°C and pH=2.
273
experiments with fast-dissolving minerals, under conditions in which the reaction is transportcontrolled, will not be influenced by the stirring effect described in the present study. Such reactions will be better studied using stirred experiments.
least for slow reactions. Fast reactions, on the other hand, may be transport-controlled and therefore should be studied using stirred experiments.
REFERENCES
Figure 4. Arrhenius plots of non-stirred and stirred kaolinite dissolution experiments. After Metz & Ganor (submitted).
Fortunately, if the reaction is fast enough the amount of fine particles in steady state is small and stirring does not influence the observed rate. The question is how fast should the reaction be in order to be "fast enough". Metz & Ganor (submitted) relate the stirring enhancement factor, SEF, to the production rate of the fine-particles using the equation:
It SEF is smaller than the analytical error, the stirring effect is negligible. For kaolinite, Metz & Ganor (submitted) estimate that the production rate of the ultra-fine particles for experiments stirred at 650 rpm range between 7.210''3 to 5.4.10-'2 mol s-'. Taking the maximum production rate observed by Metz & Ganor (submitted), an analytical uncertainty in the rate of 10% and surface area of the large grains of 1 m2, no stirring effect will be observed if the dissolution rate of the large grains is faster than 5.1 0-' mol m-2 s-'. This dissolution rate is 4 orders of magnitude slower than the surface-controlled nepheline dissolution rate at 25°C and pH 3 (5.10-7 mol m-2 s-') (Tole et al. 1986). Although the production rate of fine particles of other minerals in different experimental setting may be larger than that observed by Metz & Ganor (submitted), it seems that it can not be large enough to influence the rate of dissolution reaction of fast dissolving mineral under conditions in which the reaction may be transportcontrolled, and therefore should be studied using stirred experiments.
'
Aagaard, P. & H.C. Helgeson 1982. Thermodynamic and kinetic constraints on reaction rates among minerals and aqueous solutions. I. Theoretical considerations. American Journal of Science 282: 237-285. Amrhein, C. & D.L. Suarez 1992. Some factors affecting the dissolution kinetics of anorthite at 25 "C. Geochimica et Cosmochimica Acta 56: 1815-1826. Berner, R.A. 1978. Rate control of mineral dissolution under earth surface conditions. American Journal of Science 278: 1235-1252. Ferrow, E.A., Kalinowski, B.E., Veblen, D.R. & P. Schweda 1999. Alteration products of experimentally weathered biotite studied by high-resolution TEM and Mossbauer spectroscopy. European Journal of Mineralogy 11: 999-1010. Furrer, G., Zysset, M. & P.W. Schindler 1993. Weathering kinetics of montmorillonite: Investigations in batch and mixed-flow reactors. In Geochemistry of clay-pore fluid interaction, Vol. 4 (eds. D.A.C. Manning, P.L. Hall, & C.R. Hughes), pp. 243-262. Chapman & Hall. Helgeson, H.C., Murphy, W.M. & P. Aagaard 1984. Thermodynamic and kinetic constraints on reaction rates among minerals and aqueous solutions. 11. Rate constants, effective surface area, and the hydrolysis of feldspar. Geochimica et Cosmochimica Acta 48: 2405-2432. Komadel, P., Madejova, J. & J.W. Stucki 1998. Variables affecting the dissolution of smectites in inorganic acids. 35th Annual Meeting of the Clay Minerals Society, 1 1 1. Metz, V. & Ganor J. submitted. Stirring effect on kaolinite dissolution rate. Geochimica et Cosmochimica Acta. Murphy, W.M., Oelkers, E.H. & P.C. Lichtner 1989. Surface reaction versus diffusion control of mineral dissolution and growth rates in geochemical processes. Chemical Geology 78: 357-380. Tole, M.P., Lasaga, A.C., Pantano, C. & W.H. White W.B. 1986. The kinetics of dissolution of nepheline (NaAlSiO4). Geochimica et Cosmochimica Acta 50: 379-392. van Grinsven, H.J.M. & W.H. van Riemsdijk 1992. Evaluation of batch and column techniques to measure weathering rates in soils. Geoderma 52: 41-57.
4 CONCLUSIONS We recommend the standard use of non-stirred reaction vessels for kinetic experiments of mineral dissolution and precipitation that are slow enough to be surface-controlled. According to our calculations, 274
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The bentonite-water interface and its role in the adsorption processes of metals T.Gavri1oaiei University "A1.I.Cuza " Iasi, Faculty of Geography-Geology, Dept. of Mineralogy-Geochemistry, 20A B-dul Carol I , 66OO-Iasi,Romania
ABSTRACT: The use of montmorillonite (which is the main mineral in bentonites) like adsorbent is a wellknown process. The aim of the paper is to assess the adsorption capacity of a bentonite using the Langmuir and Freundlich non-competitive isotherm equations. The adsorption of aqueous solution of Fe"' and Cr"' (initial concentrations between 25 and 125 mg.1-') onto different amounts of clay mineral (0.lg to 0.3g) was determined. According to Langmuir isotherm, the maximum adsorptions for both ions were displayed for 0.1 g bentonite compared with the other two amounts. The maximum adsorptions and the equilibrium constants for Cr"' onto bentonite were smaller than those recorded for Fe"'. According to Freundlich isotherm for Fe'" adsorption, the highest Freundlich constant value was displayed for 0.3 g bentonite and the smallest value for 0.2 g bentonite, contrary to values recorded with Langmuir isotherm. 1 INTRODUCTION The removal of toxic heavy-metal contaminants from different effluents by sorption onto solids leads to high quality treated wastewaters. Many sorbent materials have been tested for various pollutants: cellulose matters, microorganisms, soils, different minerals, synthetic resins etc. These materials are capable of concentrating metal species from dilute aqueous solutions and accumulating them within the interface layer. The adsorption process by different materials arises from the coordination of the metallic ions to different functional groups on the adsorbent surface. These coordinating groups, generally include carboxyl, hydroxyl, phosphate or aminogroups. The kinetics of metal uptake through to physical adsorption at the adsorbent surface is very rapid and occurs in a short time after the two components come into contact. The removal of heavymetals contaminants is affected by several factors, which include the specific surface properties of the two phases, the physiochemical parameters of the solutions such as pH, temperature, initial metal ion concentration, the weight of adsorbent, the agitation and the aeration (Aksu et al. 1997). The bentonites are better adsorbents for heavy metal ions from their solutions. The actual uses are advanced through the clays modification for an amazing variety of applications in different fields. The main structural component of phyllosilicates is 2:l layer, which consists of octahedral sheets, sandwiched between two tetrahedral sheets. The 2:l lay-
ers are separated by an interlayer region, which contains hydrated alkaline and alkaline-earth cations. The main clay mineral is montmorillonite, which belongs to the smectites group (Lagaly 1994). In montmorillonite Al is the predominant cations in the octahedral sites. The adsorption depends greatly on the pH of the medium, surface charge state of the adsorbent and adsorption affinities (Campbell & Davies 1995). Chemical composition (oxide wt %) and the main properties of the sample used in experiments are given in Table 1. Table 1. Chemical composition (oxide wt %) and some mineralogical properties of the sample used in experiments. Oxides SiO, TiO, Fe203 CaO MgO Na20 K20
Fe0 MnO
so:-
Weight (%) 47.43 0.21 16.78 5.89 5.88 3.12 1.77 1.78 1.22 0.23 2.09
Loss on ignition,900"C 13.30
Mineralogical properties Montmorillonite content in bentonite (% min) Cation exchange capacity (bentonite) (meq/100g) Adsorbing index (%) Decolourisation power
(%I
Inactively substances
(%I
Bentonitic number Density ( g . ~ m - ~ ) Specific surface area (BET method) (m2.g-')
Values 7.5 10 13.10 1.72 5.0 18.22 0.11 2.84 67.2
The layer structure provides different types of surface and different sites for the interactions with ionic species from solutions. 275
A different model for the adsorption processes, the Freundlich isotherm, is also considered. This equation is expresses as:
2 MATERIALS AND METHODS A bentonite clay was selected for the adsorption experiments. The natural clay was crushed and sieved to recover a fraction used in the sorption experiments (0.09 mm). A. The main procedure required to construct adsorption isotherms is to add known amounts of sorbat to different amounts of adsorbent. After reaching the equilibrium, the amount of M"' adsorbed is determined from the difference between the concentrations in the initial and the equilibrium states. For this purpose, the bentonite samples (0.1 g, 0.2 g and 0.3 g, respectively) were weight directly into Erlenmeyer flasks. Different volumes (25 ml, 50 ml and 75 ml respectively) in five different concentrations of Fe'I and CrV' salts solutions (25, 50, 75, 100 and 125 mg.1-') were added to each adsorbent samples. The flasks were shaken vigorously and placed at room temperature for 2h. After equilibrium, the solutions were filtered and analyzed. B. Stock Fe"' solution was prepared from ferric chloride in distilled water (with 5 ml H2S04) and CrV' solution was obtained by dissolving the exact quantity of potassium dichromate in distilled water. C. The chemical equilibrium between adsorbent and ionic species from solutions can be represented by adsorption isotherm equations, which are mathematical functions describing quantitatively (at a constant temperature) the relationship between equilibrium metal concentration (ce) left in solution after binding and the initial metal concentration (CO).The non-competitive Langmuir and Freundlich isotherms are suitable for describing the adsorption of heavymetal ions by bentonite (Aksu et al. 1997). The expression of the non-competitive Langmuir model is given by the Equation (1) (Isac et al. 1995):
x = KF Ce""
(4)
where x is the ion amount adsorbed per unit of adsorbent weight (mg.g-'), ce is the metal concentration at equilibrium (mg.1-'), KF and n are constants relating to adsorption capacity and intensity, respectively. Equation (4) can be linearized in logarithmic form: In x = (l/n) In ce + In KF
(5) A plot of In x versus In ce should give a straight line. D. The concentration of unadsorbed CrV' and Fe"' ions were determined spectrophotometrically using diphenyl-carbazide and sulphosalicylic acid as complexing agents at 540 nm and 420 nm, respectively.
3 RESULTS AND DISCUSSION The case of sorption process considered here involves a solid phase (bentonite) and a liquid phase containing a dissolved species (Fe"' or CrV1)to be sorbed. Due to the "affinity" of the sorbent for the sorbate species, the latter is attracted into the solid and bound there by different mechanisms. The noncompetitive Langmuir and Freundlich models (at the optimum pH) were used to describe the monocomponent adsorption phenomena (Isac et al. 1995). It was shown in other place (Aksu et al. 1997) that the adsorption of CrV' - Fe"' from aqueous solution was more efficient with decreasing pH values. The equilibrium data for these two metal ions fit both the non-competitive Langmuir and Freundlich models. The calculation of the ion uptake is based on the material balance of the sorption system. The performance (its uptake) needs to be compared through the isotherms at the final equilibrium concentration. Equilibrium isotherms were constructed for each metal solution in contact with different amounts of adsorbent. Selected Langmuir and Freundlich isotherm plots are displayed in Figures 1-2, from experimental results. Not all of these plots will be done here, only those for 0.1 g adsorbent material. According to Langmuir isotherm (Table 2) the maximum adsorption for Fe"' was displayed for 0.1 g bentonite (rmax = 109.64 mg.g-') compared with the other two amounts (rmax = 70.57 rng.g-' for 0.2 g and rmax = 43.21 mg.g-' for 0.3 g bentonite, respectively). The equilibrium constants values varied from k = 0.341 (for 0.1 g bentonite) to k = 2.254 (for 0.2 g bentonite). The maximum adsorption for CrV' onto bentonite (rmax = 97.84 mg.g-' for 0.1 g, rmax = 36.62 mg.g" for 0.2 g and rmax = 30.65 mg.g-' for 0.3 g bentonite, respectively) and the equilibrium constants (k = 0.121 for 0.1 g, k = 0.212 for 0.2 g and k = 0.482 for 0.3 g bentonite, respectively) were smaller than those recorded for Fe"' ion.
r = rmax [kce / (l+kce)] (1) where r is the amount of metal bounds to unit of adsorbent weight (mg.g-'), rmax is the maximum amount of metal ion per unit of weight of adsorbent, Ce is the metal concentration left in solution after the equilibrium is attained (mg.1-') and k is a constant related to the affinity of the binding sites. This equation may be arranged in the linear form: From graphical plot of (ce/T) versus (Ce) it can be considered the terms rmaw and k. The shape of the Langmuir isotherm can indicate the "favourability" of the adsorption process (McKay et al. 1989). Equation 3 defines the equilibrium parameter "a": a = 1 / (l+kco) (31 The parameter indicates the shape of the isotherm as follows: a>l , unfavorable process: a=l , linear process; O
the adsorption's stabilities are of the same order of magnitude. The lowest yield of CrV' adsorption onto 0.3 g bentonite makes Freundlich empirical isotherm impossible because the ion does not obey the isotherm equation. Table 2. Values of Langmuir and Freundlich constants for Fe111 and CrVI ions adsorption onto different amounts of bentonite (0. Ig to 0.3g) and different volumes (25m1, 50m1, 75m1, respectively). Analyzed Langmuir isotherm Freundlich isotherm ions rmax k a KF n Fe 24.59 0.341 0,022 1.189 2.118 109.64 0.084 0.045 2.044 0.801 0.602 0.164 0.016 1.066 88.96 1.200 2.254 0.003 1.414 6.52 0.774 0.072 0.052 1.010 59.80 0.788 0.622 0.004 1.332 70.57 1.117 0.354 0.022 1.733 6.21 0.784 43.21 0.065 0.057 1.370 36.12 0.272 0.009 2.568 0.588 Cr "I 29.48 0.058 0.121 1.127 1.070 97.84 0.073 0.051 1.237 0.781 86.26 0,121 0.021 9.193 1.111 6.95 0.137 0.055 1.173 2.032 30.28 0.212 0.018 1.244 0.994 36.62 0.068 0.037 1.407 0.802 6.21 0.059 0.119 17.04 0.482 0.008 30.65 0.137 0.019 -
'"
Figure 1. Langmuir isotherm plots for the adsorption of Fe"' (a) and Crm (b) onto bentonite (0.1 g) for different volumes of solution (initial concentration between 25 and 125 mg.1-').
The examination of the adsorption maxima (r,,, Table 2) suggests that bentonite has lower capacities to adsorb both Fe"' and CrV' at 0.3 g compared with 0.1 g adsorbent material. The lowest values are displayed for the adsorption of Fe"' and CrV' from the same volume of solution (25 ml) onto 0.2 g and 0.3 g bentonite, respectively. An exception appears to Langmuir bonding energy (k) at the adsorption of ferric ion onto 0.2 g bentonite the sample being anomalous in its value. The Fe'*' bonding affinities were higher than the values for Crv' ions (Table 2). A large value of k also implicates strong bonding of Fe ions to the bentonite. The adsorption capacity of bentonite for CrV' ions was generally less than the adsorption capacity of Fe"' and an increase in the metal ion concentration decreases the adsorption yield for each metal ion. The shape of isotherms ("a" parameter) indicates the "favourability" of the adsorption process and a chemisorption's one (McKay et al. 1989). Referring to this parameter, both adsorptions are favorable processes. According to Freundlich isotherm for Fe"' adsorption (Table 2), the highest Freundlich constant value was displayed for 0.3 g bentonite (KF = 2.568) and the smallest one for 0.2 g bentonite (KF = 1-010), to recorded with Langmuir isotherm. Concerning the adsorption process for both ions the Freundlich constants (n) showed that
Figure 2. Freundlich isotherm plots for the adsorption of Fe"' (a) and Cr"' (b) onto bentonite (0.1 g) for different volumes of solution (initial concentration between 25 and 125 mg.1-I).
277
At the first stage of process a rapid equilibrium is establishes between adsorbent surface and the adsorbedhnadsorbed ionic species from solution. The magnitude of KF and n showed easy uptake of these two ionic species from solutions, with a high adsorption capacity of bentonite for CrV' ions for 0.1 g bentonite (KF 9.193) compared with Fe"' ions. Concentration of Fe"' and various Fe"'-hydroxocomplexes are significantly affected by pH values (Stone 1997). Fe(OH)3 amorphous forms first when Fe salts are added to water, but immediately slow conversion into hydro-complexes takes place. In the absence of the organic li ands, Fe"' speciation is 5 dominated by the Fe(OH)2 ions below pH 6.4 and by the Fe(OH)3(s) form at higher pH values. Hexavalent chromium solution species are thermodynamically stable as chromate (CrOt-) or dichromate (CrzO:-) anions over a large pH range (Stone 1997). Anions adsorption onto mineral surface is strongly pH-dependent: adsorption is generally greater at lower pH values and decreases with the increasing pH values. Adsorption of weakly bound anions, like chromium anions, is more dependent on the ionic strength, suggesting a greater dependence on electrostatic energy contributions to the free energy of adsorption. In the case of the analyzed species the adsorption process is best described as an outer-sphere complex, while the adsorption process of the strongly bound ionic species (phosphate, selenite) to mineral surface is described via an inner-sphere complex (Stone 1997). The most important difference between inner-sphere and outer-sphere complexation lies in the fact that inner-sphere complex affects the orbital system of the coordinated metal ions. The layer structure of clays minerals provides different types of surfaces and different sites for the interactions with anions or cations (Lagaly 1994). The two ionic species from their aqueous solutions can interact with the exchangeable interlayer cations (Na, Ca). The oxygen ions from Si-0-Si bonds are weak acceptors for hydrogen bonds. The isornophous substitutions in the tetrahedral layer intensi@ the surface interaction of the silicate layer with aqueous solution of ions (Low 1994). Many reactions are related to the interlayer cations: hydratation, complex formation, exchange reactions or solvation of interlayer cations.
level of metal concentrations. 1. The bentonite sample that was tested for Fe"' and CrV' sorption conformed both to the Langmuir and Freundlich models. If the Langmuir model gave a better fit to the data, the Freundlich model gave some irrelevant results. 2. The adsorption maximum (rmm) is accepted as a measure of sorption capacity. The highest values found for Fe"' ions indicated that ferric ion was better adsorbed than chromium ion. The adsorbent has a good capacity of sorption for the sorbate in the low residual concentrations range. According k values, the bentonite has a high affinity at 0.2 g for ferric ion and at 0.3 g for chromium ion, respectively. 3 . The Freundlich adsorption isotherm is less sensitive than the Langmuir one in describing the phenomena. For ferric solutions the highest Freundlich constant value was displayed for 0.3 g bentonite and the smallest one for 0.2 g bentonite, contrary to values recorded with Langmuir isotherm. 4. The values of equilibrium constants showed that the affinities of ionic species towards the sorbent decrease in the following order: Fe > Cr. 5. The ionic species uptake took place very fast during the initial stage of contact between the components. In the adsorption onto bentonite under acidic conditions, the main mechanism responsible for the metal uptake is the complex formation and less the ion exchange. The removal of ionic species from aqueous solutions through the adsorption onto minerals will be quite useful due to low costs of the processes and the wide availability.
=I
REFERENCES Aksu, Z., U. Acikel & T. Kutsal 1997. Application of multicomponent adsorption isotherms to simultaneous biosorption of iron (111) and chromium (VI) on C. vulgaris. J. Chem. Tech. Biotechnol., 70:368-378. Campbell, L.S. & B.E. Davies 1995. Soil sorption of caesium by Langmuir and Freundlich isotherm equations. Applied Geochem., 10:715-723. Isac, V., A. Onu, C. Tudoreanu & Gh. Nemtoi 1995. Chimie Jzic6, lucriiri practice. ChiTinFiu, Stiinp. Lagaly, G. 1994. Bentonites: adsorbents of toxic substances. Progr. and Polym. Sci., 95: 61-72. Low, P.F. 1994. The clay/water interface and its role in the environment. Progr. Colloid and Polym. Sci., 95: 98-107. McKay, G., H.S. Blair & A. Findon 1989.Equilibrium studies for the sorption of metal ions onto chitosan. Indian Jour. of Chemistry, 28A: 356-360. Stone, A.T. 1997. Reactions of extracellular organic ligands with dissolved metal ions and mineral surfaces. In J. Banfield & K.H. Nealson (eds), Reviews in Mineralogy, vol. 35, Geomicrobiology, Interactions between microbes and minerals, 309-344, Washington: Min. Soc. of America.
4 CONCLUSIONS As a result, it could be said that the mono-ion system could be defined with both the Langmuir and Freundlich adsorption and both of them could be used to model the adsorption of binary systems from aqueous solutions. The bentonite offered a practical approach to the multi-component removal of ionic metals from waste-waters due to the pH of them and 278
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Limiting mechanisms of borosilicate glass alteration kinetics: Effect of glass composition S .Gin & C.Jkgou Commissariat a 1 ’EnergieAtomique (CEA), Valrh6lMarcoule,France
ABSTRACT: A series of borosilicate glasses was studied to evaluate the effect of the glass composition on its leaching properties. Tests were conducted to determine the initial dissolution rate for each glass and the dissolution kinetics at advanced stages of reaction progress. The effect of the glass composition on the initial rate is interpreted on the basis of the network forming or modifying role of the glass elements and their free energy of hydration. The kinetic data at high reaction progress are used to discuss the protective role of the gels formed by recondensation of hydrolyzed species. The effect of the elements on the gel protective properties was found to be very different than their effect on the initial rate. Synergistic effects were identified and calcium was found to be a key element in the gel. Further characterization of the gel structure and texture will be necessary for a fuller understanding of this effect. drolyzed species (Gin etal., 2000a, b; Vernaz & Gin, 2000; JCgou et al., 2000). The Commissariat a 1’Energie Atomique (CEA) has been investigating simplified glass compositions for several years in order to better understand the elementary mechanisms of borosilicate glass alteration (JCgou, 1998; JCgou et al., 2000). This article presents the results obtained with a series of simplified glasses successively incorporating additional elements in order to assess their effects on the initial rate and on the protective properties of the gel. The element molar ratios in all the test compositions were the same as in the SON68 nuclear reference glass to permit valid comparisons with this complex formulation.
1 INTRODUCTION Borosilicate glasses are used in France for containment of the fission product solutions produced by reprocessing spent nuclear fuel. From the perspective of geological disposal of these long-lived wasteforms, predicting their long-term behavior under different alteration conditions is a necessary step in demonstrating the safety of a repository site. As water is the primary agent liable to result in the release of radionuclides, the mechanisms and kinetics of glass alteration must be studied in aqueous media. Work carried out over the last two decades has shown that nuclear glass alteration begins by an ion exchange process involving interdiffusion between protons and network-modifying alkali metals. The kinetics are subsequently limited by hydrolysis of the covalent bonds in the glass network, and as long as the solution remains sufficiently dilute the alteration rate remains at a maximum value (the “initial rate” ro),which depends only on the glass composition, the temperature and the pH. The rate then gradually diminish by several orders of magnitude compared with the initial rate as the solution becomes increasingly saturated-notably with respect to silicon. The diminishing rate under saturation conditions was first interpreted in terms of an equilibrium between the glass and solution (Grambow, 1985) or between the gel and solution (Bourcier, 1992), and more recently by the diffusion barrier effect of the gel formed by recondensation of a fraction of the hy-
2 MATERIALS AND METHODS 2.1 Glass chemical composition Table 1 shows the theoretical compositions of the ten test glasses. Analyses were performed to confirm that the compositions corresponded to the actual values. 2.2 Glassfabrication and specimen preparation The test glasses were melted between 1350 and 1500°C in a platinum crucible in an induction-heated melter according to the method described by Pacaud et al. (1990). The homogeneity of glass melts 1 to 6 was verified by Rayleigh-Brillouin scattering; the 279
reactor was placed in a 1-liter container with a few milliliters of water to limit evaporation. The containers were then placed on a stirring device mounted on rollers rotating at 3 rpm. Solution samples were taken at regular intervals from each reactor, ultrafiltered to 10 000 daltons, acidified, and analyzed by ICP-AES. 2.5 Expression of Results
The glass alteration rate is defined as follows:
Landau-Placzek ratios ranged from 50 to 200, indicating that these glasses contained no heterogeneities exceeding 15 A (Jkgou, 1998). In view of their compositions, the other simplified glasses were also assumed to be homogeneous. The glass samples were ground and sieved to recover the 40-100 pm and 100-125 pm fractions. The powders were cleaned several times ultrasonically in acetone and finally in water to eliminate any remaining fine particles. After drying, the specific surface area of the powder was determined by the BET method. The shape factor (i.e. the ratio between the BET specific surface area and the geometric surface area determined by considering the grains to be spherical) ranged between 2 and 3.5 for all the specimens; this range is characteristic of nonporous materials.
where r is expressed in g.m-2d-1, X, is the mass fi-action of element B in the glass, S/V is the ratio (m-I) between the glass surface area and the solution volume, and C, is the boron concentration (mg.1-I) in solution. Boron is a good alteration tracer for this type of glass, as it is primarily a network former and is not retained in the alteration products. In the case of a highly altered glass specimen, the grain size reduction is taken into account by means of the shrinking-core model described by Jkgou et al. (2000). For the test under saturation conditions, the altered glass percentage is calculated from the following relation:
2.3 Testing under "initial rate conditions "
where Vis the solution volume (1) and rn, is the initial glass powder mass (g).
The initial dissolution rates of the 10 glass specimens were determined by static leaching experiments with the 100-125 pm powder fraction at pH 9 (0.1N KOH) at 90°C (ltO.5"C) with a glass-surfacearea-to-solution-volume (S/V)ratio of 0.1 cm-I. The test reactors were 1-liter PTFE vessels in which the solution was continuously homogenized by a magnetic stirring rod. Solution samples were taken at regular intervals, acidified with 1N HNO, and analyzed by ICP-AES.
3 RESULTS AND DISCUSSION
The results are summarized in Table 2. The first line indicates the initial rates measured at 90°C and pH 9 for all the test glasses. Under these conditions, glass dissolution was congruent (rsi= r, = rAl...). The network-forming or modifying effect of the elements incorporated can be identified by analyzing these results. For example, adding aluminum to Glass 1 diminishes the initial rate (compare Glass 1 and Glass 2), while adding calcium to Glass 2 increases the initial rate. This simple reasoning shows the favorable effect of Al, Zr and Ce, and the unfavorable effect of Ca on the initial rate.
2.4 Testing under "saturation" conditions The glass alteration kinetics at high reaction progress were studied by static leaching experiments at 90°C with the 40-100 pm powder size fraction; the S/Vratio was adjusted to 80 cm-I. Each 500 ml PTFE Table 2. Summary of kinetic data. Glass
1
r, (g.rn.*d-') Csi* (rng.1-l)
4.5 1.1 2.4 1.7 1.6 2.2 1.1 9.5 8.5 320 200 175 164 164 194 85 320 345 9.2 8.9 9.1 9.0 9.0 9.3 8.7 9.2 9.1 55 43 7.3 6.4 6.7 3.5 25 12.5 7.5 1.6 x 10'' 3.2 x 10-3 3.7 x 10-4 4.9 x 10.' 5.4 x 10" 4.4 x 10-4 3.0 x 10'3 4.7 x 10-3 i5 x 10.' 280 350 6500 3500 3000 5000 370 2000 1170000
PH,,,, G,,,,yr) ("/.I lily,) (gm"d") rho (1 yr)
2
3
4
5
6
280
7
8
9
SON68 2.2 117 9.2 3.1 4.5 x 10.' 4900
Figure 1. Initial dissolution rate versus standard free enthalpy of hydration.
JCgou (1998) applied three structural models to simulate the effect of the composition on the initial dissolution rate of these glasses. The first model, based on calculating the nonbridging oxygen atoms (NBO), simply indicated a qualitative trend and was considered unsatisfactory mainly because the number of NE30 varies only slightly within this series (from 0% for Glass 7 to 3.5% for Glass 8), but also because this calculation does not take the nature of the chemical bonds into account. The model developed by Feng & Barkatt (1988), based on the bonding energy, was then tested with satisfactory results provided the AI-0 bond was assigned an energy three times higher than the value given by Sun
(1947), which is difficult to justify. Finally, we tested the thermodynamic model based on the free energy of hydration (Jantzen & Plodinec, 1984). Figure 1 shows the relatively satisfactory fit between the calculated and measured results. In this model the initial rate follows a relation of the type hr,, = a + bAG&d,where AGiYdis the free energy of hydration of the glass as calculated from Paul's model (Paul, 1977). This model allows for the nature of the hydrolyzed species, and correctly accounts for the results observed with this type of glass. It remains empirical, however, since the glass structure is taken into account in a highly simplified manner as a mixture of oxides and silicates. The kinetic data obtained during the test at 80 cm" are shown in Table 2 and Figure 2. In all the tests the pH rapidly stabilized at about 9; this can be explained by assuming that the buffering effect of boron offset the increase in the pH due to interdiffusion. After one year of alteration under these conditions, the altered glass fraction was highly dependent on the glass composition, but not in the same order as under the initial rate conditions. Glasses 2 and 7, for which initial rates were the lowest, exhibited the greatest alteration along with Glass 1. The glass compositions may be classified into three categories according to their alteration kinetics (Figure 2): those for which the dissolution rates diminished continuously (Glasses 1 to 7 and SON68); Glass 8, for which renewed alteration was observed; and Glass 9, for which alteration suddenly ceased after a few weeks.
Figure 2. Normalized mass loss versus time for the 10 glass specimens during static leaching at 90°C (S/V= 80 cm-').
281
Glass 1 was the subject of a detailed investigation (JCgou et al., 2000) showing that the diminishing alteration rate was not attributable to a decrease in the chemical affinity associated with saturation of the solution with respect to silicon, contrary to the results expected from the models proposed by Grambow (1985) or Bourcier (1 992). Thermochemical measurements providing an accurate assessment of the thermodynamic stability of simple glass compositions were performed by Linard (2000), confirming this results for Glass 1 as well as for the other simple glasses in the series. JCgou (2000) attributed the diminishing alteration rate of Glass 1 to the diffusion barrier effect of the gel formed by recondensation of hydrolyzed silica. The protective properties of the gel are related to the recondensation rate, which in turn depends on the dissolved silicon concentration near the glass surface and on the degree of network restructuring. The gel must be considered as a dynamic barrier that limits glass alteration only when a sufficiently dense layer has formed at the reaction interface and when the concentrations of dissolved species-notably silicon-are high enough to limit the gel dissolution (i.e. “saturation” conditions). From this perspective, the protective properties of the gel can be assessed directly by observing the sample alteration rates: the greater the drop in the rate, the more protective the gel. The following key observations can be noted on the basis of a preliminary analysis: - Under these alteration conditions, silicon alone forms a relatively nonprotective gel (Glass 1). - Adding A1 andor Zr to Glass 1 (Glasses 2 and 7) does not significantly enhance the protective nature of the gel. - Glasses containing Ca and one or more sparingly soluble elements (Glasses 3, 4, 5 , 6, 9 & SON68) produce more protective gels than Glasses 1, 2 and 7. - The association of Zr and Ca (Glass 9) is particularly effective, causing glass alteration to cease completely. - The properties of the gel containing Si and Ca (Glass 8) clearly evolved to a significant extent over time, resulting in renewed glass alteration. - Ce also enhances the protective quality of the gel. - Glass 6 constitutes a suitable analog for SON68 nuclear glass (same initial rate, same rate decrease at high reaction progress). An analysis of the kinetic data shows that the chemical elements do not have the same role under “saturation” conditions (in which the gel effect predominates) as under the initial rate conditions (where the glass structure is decisive). It should also be emphasized that there is no clear logical relation between the apparent silica solubility (C,,*) and the protective effect of the gel.
This study has revealed the essential role of the gel chemical composition on its protective properties. The next step will be to determine the key parameters of this effect by characterizing the structure (NMR, EXAFS, Raman spectroscopy) and texture (TEM, SAXS) of the gel, and by the use of atomistic models of the gel structural entities. FEFERENCES Bourcier, W.L., Pfeiffer, D.W., Knauss, K.G., McKeegan, K.D. & D.K. Smith 1990. A Kinetic Model for Borosilicate Glass Dissolution Affinity of a Surface Alteration Layer. In Scientific Basis for Nuclear Waste Management: Mat. Res. Soc. Symp. Proc. 176:209-216. Feng, X. & A. Barkatt 1988. Structural thermodynamic model for durability of nuclear glasses. In Scientific Basis for Nuclear Waste Management: Mat. Res. Soc. Symp. Proc. 112~543-554. Gin, S. & E. Vernaz 2000 Protective Effect of the Alteration Gel: A Key Mechanism in the Long-Term Behavior of Nuclear Waste Glass. In Scientific Basis for Nuclear Waste Management XXIV, Sydney, August 2000 (in press). Gin, S., Ribet, I. & M. Couillaud 2000. Role and Properties of the Gel Formed during Nuclear Glass Alteration: Importance of Gel Formation Conditions. International Workshop on Glass in its Disposal Environment, Bruges, April 2000. To be published in J. Nucl. Mat. Grambow, B. 1985. A General Rate Equation for Nuclear Waste Glass Corrosion. In Scientific Basis for Nuclear Waste Management., Mat. Res. Soc. Symp. Proc. 44: 15-27. Jantzen, C.M. & M.J. Plodinec 1984. Thermodynamic model of natural, medieval and nuclear waste glass durability. J. Non-Cyst. Solids. 67:207-223. Jtgou, C. 1998. Mise en kvidence expbimentale des mdcanismes limitant la cine‘tiqued’altdration du verre R7T7 en milieu aqueux: Critique et proposition d ’&volutiondu formalisme cinktique. Ph.D. thesis. UniversitC de Montpellier 11, France. Jtgou, C., Gin, S. & F. Larch6 2000. Alteration Kinetics of a Simplified Nuclear Glass in an Aqueous Medium: Effects of Solution Chemistry and of Protective Gel Properties on Diminishing the Alteration Rate. Journal of Nuclear Materials. 280:2 16-229. Linard, Y . 2000. Thermochimie des verres borosilicatb: Contribution a I’dtude de l’altkration des verres de confinement de dbchets radioactif. Ph.D. thesis. UniversitC Paris VII, France. Pacaud, F., Jacquet-Francillon, N., Terki, A. & C. Fillet 1989. R7T7 light water sensitivity to variations in chemical composition and operating parameters. In Scientific Basis for Nuclear Waste Management., Mat. Res. Soc. Symp. Proc. 127:105-112. Paul, A. 1977. Chemical durability of glasses: a thermodynamic approach. J. Materials Sci. 12:2246-2268. Sun K. 1947. Fundamental condition of glass formation. J. Amer. Soc., 30:277-28 1. Vernaz, E. & S. Gin 2000. The apparent solubility of nuclear glasses. In Scientific Basis for Nuclear Waste Management XXIV, Sydney, August 2000 (in press).
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Wafer-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeiflinger, Lisse, ISBN 90 2651 824 2
Heavy-metal binding mechanisms in cement minerals C .A.Johnson, I .Baur & FZiegler Swiss Federal Institute of Environmental Science and Technology (EAWAG), CH-8600 Diibendor- Switzerland
ABSTRACT: The increasing use of cement to stabilise wastes and the use of wastes as secondary building materials makes it ever more important to understand metal binding mechanisms in the cementitious matrix. Calcium silicate hydrate (CaOSi02-xH20 or CSH) and basic Ca sulphoaluminates such as monosulphate (3CaOA1203CaS04.12H20) or ettringite (3CaOA12033(CaS04).32H20)all have a potential for heavy metal uptake. The mechanisms are poorly understood. This paper presents sorption experiments of Zn to CSH(1) and selenate sorption to ettringite. The results indicate that Zn sorbs strongly to CSH(1). There is a rapid and a slow sorption step, the latter being ascribed to the difhsion of Zn into the CSH(1) particles. Selenate, on the other hand, sorbs only weakly to ettringite. This anion does not appear to exchange for SO;-, but is sorbed rapidly to the mineral surface.
1 INTRODUCTION The increased use of cement to stabilise wastes and the use of secondary materials in cement production has made it necessary to understand the potential risks associated with a possible rise in heavy metal concentrations in the hydrated cementitious matrix (Atkins and Glasser 1992, Spence 1993, Cougar et al. 1996). Studies show that in hydrated cementitious matrices, metal mobility is generally low due to physical factors, such as a low permeability, and by geochemical binding within the cement matrix. There appear to be three basic types of binding mechanism. A metal ion may (i) be bound in the alkaline cement matrix as an oxide, mixed oxide or other solid phase, (ii) sorbed onto surfaces or (iii) be incorporated into hydrated cement minerals. The precipitation of heavy metal containing solid phases does seem to be a limiting factor with regard to the second and third mechanisms. However, for heavy metals that are sufficiently soluble in basic media, such as Zn, Pb or Cr(VI), incorporation in hydrated cement minerals appears to be significant. Calcium silicate hydrate (CaOSiO2.xH20 or CSH) is the most abundant component of hydrated portland cement (-6O%), has a large capacity for ion uptake and is thus a prime candidate for heavy metal binding. Basic Ca sulphoaluminates such as monosulphate (3CaOA1203CaS04.12H20)or ettringite (3CaOA1203 3(CaS04).32H20) also have a potential for heavy metal uptake. Their structures are relatively tolerant of substitution without a change in structure. A number 283
of ions are reported to substitute at the different sites in ettringite namely divalent cations such as Pb2', Cd2+or Zn2' for Ca2', trivalent cations such as Cr3+, Ni3+ or Co3+ in lace of A13+ and diverse anions in exchange for SO42-.
Figure. 1 Schematic representation (adapted from Stumm 1992) of sorption isotherms representing the following mechanisms: a) adsorption, b) sorption and surface precipitation solution, c) adsorption and precipitation, d) adsorption and precipitation of metastable precursor and e) as d) but with the transformation of the metastable precursor into the stable phase
This paper illustrates the uptake of two ions into cement minerals. The first example is of Zn binding to CSH and the second of the anion selenate (Se04:-) to ettringite. The study explores the relationship between aqueous and solid-phase concentrations at
“quasi” equilibrium (see Figure 1) and uses supporting evidence to obtain an understanding of binding mechanisms. 2 METHODS
2. I Materials All chemicals were at least of p.a.-grade. Solutions were generally prepared using boiled ultrapure deionised water (Barnstead Nanopur, 17 MW). In order to prevent a CO2 contamination, all procedures involving alkaline solutions and solids were performed in a glove box in an argon atmosphere with pc02 < 1 ppm. The synthesis of CSH(1) (with a Ca:Si ratio of 1) was performed after Atkins et al. (1992). For the subsequent experiments, the size fraction < 63 pm was used Ettringite was prepared after the method of Atkins et al. (1991). The material was then sieved and the size fraction < 0.125 mm was used.
2.2 Analytical Method The pH value of the solutions was measured using a combined glass electrode (Metrohm 6.0202.100) calibrated by an acid-base titration. Dissolved Zn, Ca and Si concentrations were measured as described in Ziegler et al. (2000). Sulfate was analysed by ion chromatography. Selenium concentrations in the Se equilibration experiments were determined by atomic absorption spectroscopy (concentrations < 0.1 mmol L-’ by the graphite furnace technique, higher concentrations by flame technique). The resulting confidence intervals for the measured concentrations at a probability of 95% are generally at most A[Zn] = *lO%, A [Ca] = *5%, A [Si] = *5%, A [Se] = *10%, A [S] = *lO% and A [OH-] = *5%.
2.3 Sorption Experiments
experiments were conducted without addition of the CSH(1) suspension or only in an electrolyte solution at the same pH value (without Ca and Si addition. A stock suspension of CSH(1) was prepared by adding 1 g of CSH(1) to 0.5 L of the presaturated solution and equilibrated for 7 days on the rotary shaker at 150 rpm. To obtain the final CSH(1) suspension, the stock suspension was diluted by adding 1 mL 50 mL of a solution containing appropriate quantities of Ca and Si in a 50-mL HDPE bottle and equilibrating again for 7 days. Zn stock solutions containing Ca and Si at a neutral pH were prepared directly before the sorption experiments. For the sorption experiments, 1-mL aliquots of the Zn stock solutions were added to the CSH(1) suspensions and equilibrated. Afterwards, the solution pH was measured in a 20mL aliquot. The remaining sample was filtered (0.4 pm nylon filter), acidified (1% HNo3), and the dissolved Zn, Ca and Si concentrations were measured. All experiments were carried out in duplicate. Experiments with ettringite were carried out at one pH value (1 1.OO), at initial Se concentrations from 0.32 pmol L-’ to 63’500 pmol L-’ and equilibration times from 5 min up to 90 d. Certain experiments were conducted without addition of ettringite. An “equilibrated” stock solution containing concentrations equivalent to 84% saturation was prepared by the addition of ettringite (0.35 g) to 1L of deionised water. Ettringite (0.05 g) was added directly to 50mL HDPE bottles, to which 25 mL of the stock solution was added. These were equilibrated for at least 7 days at 25°C on the rotary shaker at 150 rpm. Selenate was then added and the mixture equilibrated. Afterwards, the solution pH was measured in a 10mL aliquot. The remaining sample was filtered (0.4 pm nylon filter), 5 mL was collected for S analysis and the rest was acidified (1% HNO3). The solutions were analysed for Se, S, Ca and Al. All experiments were carried out in duplicate. The solid phase samples analysed by X-ray powder diffraction (XRPD) to determine mineralogical composition (Scintag XDS 2000 diffractometer, Cu-Ka, 1.5406 nm, 2000 W, 45 kV, 40 mA).
Because of the high solubility of both CSH(1) and ettringite and in order to avoid changes in the stoichiometry of these minerals, the experiments 2.4 Thermodynamic Equilibrium Calculations were carried out in solutions that had been pre- Thermodynamic equilibrium calculations were performed as described in Ziegler and Johnson equilibrated with the appropriate solid phase. Experiments with CSH(1) were carried out at three (2001) and Baur et al. (2001). different pH values (1 1.7, 12.48, 12.78), at initial Zn concentrations from 4.8 pmol L-’ (undersaturated 3 RESULTS AND DISCUSSION with respect to zinc hydroxides) to 4800 p mol L-’ (supersaturated with respect to zinc hydroxides) and Zinc sorption to CSH(1) is illustrated in Figure. 2. equilibration times from 5 min up to 87 d. Certain There appears to be a linear relationship between 284
sorbed and dissolved Zn concentrations at concentrations below 10 pM after 4 days and after 87 days, corresponding to a Zn(II)sorbed:Siatomic ratio of approximately 15%, depending on pH. It can be seen that sorption increases during this time. It should be noted that at low Zn concentrations, it was found that
However, this cannot be incorporation via isomorphous substitution as a solid-solution in the classical sense. At solution concentrations greater than 10 pmol LThe isotherms deviate from linearity and indicate the precipitation of secondary phases. These were determined by XPRD to be P2-Zn(OH)2 below a pH value of 12 and Ca zincate (Zn2Ca(OH)6.2H20) above pH 12. The sorption isotherm for selenate on ettringite looks very similar to that of Zn on CSH I with a few exceptions. The solubility of the SeO)-)ion is high because there is no solubility-controlling phase involving Ca or Al. The relationship between dissolved and bound concentrations appears to be linear over a very wide range of concentrations, from 0.32 to 63’500 pmol L-’. It should also be pointed out that the affinity of the selenate ion for ettringite is a lot lower than the affinity of Zn for CSH(1). There are a number of possibilities with regard to binding mechanisms. Firstly Se0:- could substitute by ion exchange for SO:- as the latter is found in channels of the tubular ettringite structure as charge compensation. Should this be the case, one could expect typical ion exchange behaviour with the proportions similar Se:S ratios in solution and the
Figure 2. Sorption isotherms of Zn to CSH(1). The diamonds and circles denote experiments in which Zn was respectively undersaturated and supersaturated with respect to Zn oxide and hydroxide.
the sorption process was initially very rapid with over 50% sorbed within 30 minutes. This could be interpreted in terms of a fast and a slow sorption step. However Zn does not stay at the surface but instead Figure 3. Sorption isotherm of Se042-aqto ettringite. diffuses into the CSH(1) particles, as supporting evidence from electron microprobe analysis shows solid phase. This does no appear to be so, since the (Ziegler et al. 2001). However, Zn does not appear to highest Se concentrations used result in Se:S ratios of substitute for Ca or Si. Since the affinity of Zn 0.1 and 59 in the solid phase and in solution towards CSH(1) is very high with estimated surface respectively. For the lowest Se addition the ratios are 0~ Sorption, either inner coverages of almost above 10% and discrete surface 4-10-6 and 2 ~ 1 respectively. precipitates could not be detected, specific adsorption or outer sphere is a more likely explanation for the and surface precipitation seem not to be the dominant observed sorption. The poor binding capacity of ettringite for selenate sorption mechanisms although they cannot be excluded conclusively. The incorporation of Zn in the is surprising, since ettringites in which Se042- and a interlayer of CSH(1) appears to be the most probable number of other anions have been substituted for have been synthesised and characterised mechanism for the observed Zn sorption to CSH(1). SO-: 285
(Kumarathasan 1990). It appears that the ettringite containing S must be thermodynamically more stable than that containing Se. For anions such as MOO?-, the formation of CaMoO4 appears to greatly reduce anion mobility (Kindness et al. 1994; Baur et al. 2001). Part of the reason for this is the high Ca concentrations in hydrated cement pore waters. Thus anions of concern are likely to be those that do not form insoluble Ca salts. 4 CONCLUSIONS Zn binding in the cement paste appears to be significant. Zinc may bind to CSH(1) during hydration or after the CSH has formed. Sorption is thus likely to control pore water concentrations in most cases. Generalising to other heavy metal cations, one of the most important factors determining sorption behaviour appears to be the solubility of the metal hydroxides. It is not possible to extrapolate to the behaviour of heavy metal cations that form insoluble hydroxides. The binding of Se in the cement paste might be poor, if Se only sorbs to ettringite and because it does not form insoluble Ca salts. However, though Se might not bind in appreciable quantities to ettringite, there are other minerals, such monosulphate and the major component of hydrated cement pastes CSH, to consider. REFERENCES Atkins, M. et al., 1991. Solubility properties of ternary and quarternary compounds in the CaOA1203-S03-H20 system. Cem. Concr. Res. 2 1: 991-998. Atkins, M. et al. 1992. Cement hydrate phases: solubility at 25°C Cem. Concr. Res. 22: 241-246. Atkins M. & F.P. Glasser 1992. Application of portland cement-based materials to radioactive waste immobilization. Waste Manag. 12: 105-131. Baur I. et al. 2001. Leaching behaviour of cementstabilised incinerator ashes: A comparison of field and laboratory measurements. Accepted by Environ. Sci. Technol.. Cougar, M.D.L. et al.1996. Ettringite and C-S-H Portland Cement phases for waste ion immobilization: A review. Waste Manag., 16: 295-303. Kindness A. et al. 1994. Immobilisation and fixation of molybdenum(V1) by portland cement. Waste Management, 14: 97- 102. Kumarathasan, P. et al. 1990. Oxyanion substituted ettringites: Synthesis and characterization; and their potential role in immobilization of As, B, Cr, Se and V. Mat. Res. Soc. Symp. Proc. 178: 83-104. 286
Spence, R.D. (Ed.) 1993. Chemisw and microstructure of soIidiJied waste forms, Lewis Publishers, Boca Raton. Stumm W. 1992. Chemistry of the solid-water interface. John Wiley and Sons, New York. Ziegler F. et al. 2001. The sorption mechanisms of zinc to calcium silicate hydrate: sorption and microscopic investigations. Submitted to Environmental Science and Technology. Ziegler, F. & C.A. Johnson 2001. The solubility of calcium zincate (CaZn~(OH)6.2H20).Submitted to Cem. Concr. Res..
Wafer-Rock Inferaction 2001, Cidu (ed.), 0 2001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Thermodynamic elucidation of Eu anomalies in REE pattern in hydrothermal fluorite G .R Kolonin Institute of Mineralogy and Petrography SB RAS, Novosibirsk, Russia
G.P.Shironosova Institute of Mineralogy and Petrography SB RAS, Novosibirsk, Russia
ABSTRACT: Thermodynamic modeling of the origin of Eu maxima and minima in REE-fluorite has been carried out. The database was created using thermodynamic properties and parameters of the HelgesonKirkham-Flowers equation for REE complexes and other aqueous species. The developed model is based on the comparison of solubility curves for EuF3 and other REE fluorides, which have been formed aRer their interaction with fluid at temperatures 350 and 100°C. It has been obtained, that when Eu(II1) dominates, its fractionation to fluid can be two orders of concentration higher than La, Nd, Gd and Lu, and the appearance of Eu minimum in the solid phase is possible. On the other hand, in very reduced conditions, when Eu(I1) dominates, its fractionation to fluid should be much lower, as compared with Gd and other 3-valent lanthanides, resulting in the formation of Eu-maximum in REE-containing fluorite. 1 INTRODUCTION The anomalous behaviour of europium at REE partitioning in minerals and rocks is under discussion during last decades, but physical-chemical reasons of the appearance of its minima in some cases and its maxima in the others are not yet clear. Indicator properties of REE pattern in fluorites from some French deposits, including the Eu/Eu* ratio, have been used as an important marker values (Grappin et al. 1979). This ratio means a real Eu content, divided by its virtual (Eu") value, which was obtained in this work by interpolation of Sm and Tb contents as the nearest Eu neighbours. Later the data on partitioning of REE in fluorite, including the Eu behaviour, have been presented in a number of papers, in particular, in Marras et al.( 1999) for fluorite from 18 Sardinia deposits of various genetic types. The results of REE partitioning in fluorite from 25 tin-ore and fluorspar deposits of the Sikhote-Alin, Russian Far East (Kupriyanova et al. 2000) are mentioned This paper discusses Eu/Eu* ratios together with special data on Eu2+and Yb2' concentration in fluorite. On the other hand, calculations of conditions for Eu(ITI)+Eu(II) transition in hydrothermal fluid were made by (Sverjensky 1984) and then later by (Wood 1990). Besides, possible methods and approaches to use geochemical information on Eu distribution and fractionation including the influence of complex formation and waterhock ratio on initiation of Eu
anomalies of minerals and rocks have been considered (Bau 1991). The aim of this study is to apply the thermodynamic approach to the problem of the influence of complex formation at Eu fractionation during the interaction of hydrothermal fluid with fluoride solid phases. In the design and realization of this task the experience, obtained during the preparation of the previous authors paper (Kolonin & Shironosova 2001) was used as a specific preliminary basis. We believe that results of the following modeling calculations can be useful in discussing the reasons and conditions of specific chemical Eu behavior in mineral-forming processes. 2 INITIAL THERMODYNAMIC DATA Data base for the main aqueous species of Eu and other REE under discussion, including hydroxide, fluorine and chloride complexes was created using the needed information by Haas et al. (1995) and Shock et al. (1997). Gibbs free energies of solid REE fluorides were calculated using the data of Greis & Haschke (1982). Thermodynamic characteristics of five light REE fluorides from the Handbook (Barin 1989) were taken into account too. The computer code "Hch" (Shvarov 1999) was used for all calculations. Temperatures of 350 and 100°C have been selected to present the limits of this parameter.
287
The analysis of the dependence of REE free energies of both lanthanide solid fluoride phases and main aqueous species depending on their atomic number which was carried out before, revealed the presence of the distinct minimum for Eu(II1) compounds. The preliminary calculation of Gibbs free energies of some separate reactions of type: LnF3 solid + HzO + 2H‘+ LnOH2++ 3HFaq has been carried out to check the expected anomaly of Eu(II1) on the background of other REE. Figure 1 demonstrates, that there are negative difference in free energies for Eu(II1) reactions of all types for both tested temperatures. This simple way has supported a possibility of primary Eu(II1) fractionation to fluid in comparison with other REE (except Yb(II1) which is the second atypical REE showing the valence +2 in chemical reactions). It was a conclusive reason to carry out more extended modeling calculations of interaction of Eu and some other lanthanides with hydrothermal fluids. 3 RESULTS 3.1 Dependence concentration
of
EuF:
solubility
on HF
Figure 2 shows the concentrations of Eu(II1) and (11) as well as La, Nd, Gd and Lu at 350°C and 1 kbar in relation to the initial HF concentration. There are the curves for pH and total free F- concentration in fluid too. The distinct “mirror” character of F-curve relative to all solubility curves shows the depletion of solubility process when F- concentration increases. Nevertheless, the Eu(II1) concentration exceeds by two order of magnitude the concentrations of other REE within the whole HF interval. The dotted lines additionally reveal full domination of Eu(II1) over Eu(LI) or Gd by characteristic ratios. It is important to explain, that the EdGd ratio was especially calculated as a theoretical analogy of the Eu/Eu* ratio which is very popular in geochemical literature. The general conclusion from Figure 2 necessitates full Eu(II1) transformation to Eu(I1) in fluid with the subsequent decrease of its solubility, if the Eu maximum on the REE pattern in natural fluorite has been fixed. 3.2 The influence of pH on EuF3 solubility
All solubility curves at 350°C have clear V-shape form, when pH is changed within rather wide interval (Fig. 3). The total F- concentration,
Figure 1. The dependence of free energies of various solubility reactions for fluorides on REE atomic number at 25 and 350°C. The marks near curves mean formulas of separate aqueous complexes and temperatures. produced as a result of REE fluorides solubility, is close to 0.07+0.03 moVkg H20. Eu(II1) concentration predominates over La, Nd, Gd and Lu concentrations approximately two order of magnitude or more. The Eu(III)/Eu(II) ratio increases from 10’ in very acid fluid to about 105.5at pH = 4.5 and 10”’ or more in alkaline fluid. It is useful to notice that the real position of the alkaline part of the Eu(II) curve can be higher because Eu(I1) hydroxide complexes are absent in the used data base. The values log[Eu/Gd] are close to 3.3 in the most interesting pH interval. This figure support the main conclusion that only the reduction of Eu(II1) to E@) can cause the appearance of Eu maximum in natural REE-fluorite pattern. It is noteworthy that any extremum can not be found only in very reduced acid fluids because of closeness of EuF3 and GdF3 solubility in Eu(I1)-stable conditions. 3.3 Speciation of Eu in hydrothermalfluid
The general similarity of stability constant for all lanthanide group can be predicted from the data by Haas et al. (1995). It means that the main Eu(II1) aqueous species should be close to that discussed for La and other REE by Kolonin & Shironosova
288
Figure 2. Total REE concentrations in fluid, which is in equilibrium with their solid fluorides (left scale) and pH, log [Eu(III) / Eu(II)], log [Eu / Gd] (right scale) depending on the initial concentration of HF (the upper F-curve shows free fluoride concentrations).
Figure 3. The same as in Figure 2, but depending on pH of equilibrium solution.
4 CONCLUDING REMARKS
1. The following [Eu(III)] > [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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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.
293
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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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2
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.
295
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
loo.ooo
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Figure 5 represents the concentrations of Type (1) and Type (2) elements found in the two-week solution at 90°C and initial pH 3. The y-axis is concentration in ppm on a log scale and the x-axis is the material identification. Symbols have been offset, horizontally, where symbol overlap occurs.
Fe Mg Fe
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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.
298
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 << 0, AG2298 rapidly approaches c with moderate increases in x. If b M 0, then the exponential term in the above equation becomes unity. In applying Chen’s method Hemingway was unable to find combinations of reactant compounds beyond a second degree of reaction order and in some instances was restricted to only two reactions (e.g., kasolite, sklodowskite, cuprosklodowskite, and soddyite). It appears that Hemingway fit his data to an exponential function, AGY = aebx, but does not provide details of how this was accomplished for minerals for which he had only two estimated free energies (at x = 0 and x = 1) nor for any of his estimates does he provide values and sources for thermodynamic data used in his calculations other than that for schoepite. For reaction order 0, Hemingway summed oxide components. For reaction order 1, Hemingway used schoepite, chalcedony, metasilicate, and water as reactants. For reaction order 2, Hemingway used schoepite, chalcedony, uranate, and water as reactants to form the minerals. Hemingway noted that in the case of sklodowskite, a “reversal” in reaction order was observed both for the meta-silicate and uranate reactions suggesting errors in reported thermodynamic data. In this study Gibbs free energies of formation were recalculated by Chen’s method using thermodynamic data from Johnson et al. 1992; Wanner & Forest 1992; Pankratz 1982; Hemingway 1982; and Cordfunke & Loopstra 1971. The results of these calculations can be found in Table 1. As an example, AG; 298 was calculated for uranophane using the data available to Hemingway in 1983 from Robie et al. (1978). The phases for reaction order 0 are uranium (VI) trioxide, calcium oxide, chalcedony, and liquid water; phases for reac-
tion order I are wollastonite, schoepite, chalcedony, and liquid water; and, phases for reaction order 2 are calcium monouranate, schoepite, chalcedony, and liquid water. The Gibbs free energies of formation were regressed with Equation (1) resulting in:
Na2O as was the case in Chermak & Rimstidt (1989) as they found these oxides were outliers and H20 and Fe(OH)2 as was the case for regressions at higher temperatures by Chermak & Rimstidt (1990). The result of this regression at 298.15 K is: $298
= -45.206
+ 0.97986 AGi298
(2)
AG2
Equation 2 together with the value for 298 for U03 (Robie et al. 1978) were utilized to estimate
The A G i 2 9 8 estimated for uranophane is thus 6205.8 kJ mol-1 using data most likely employed by Hemingway.
or uranophane is about 8 kJ mol-1 more stable than that estimated using the older thermodynamic data.
ge298 for UO3. As pointed out by Chermak and Rimstidt the errors associated with these estimates are not known and could be significant. Nevertheless, this estimate together with values for the reduced free energies of oxides from Chermak and Rimstidt were used to estimate Gibbs free energies of formation for uranyl silicates (Table 1). This approach, with different phases chosen for the model is also the basis for estimates by Chen et al. (1999) (Table 1).
5 ESTIMATION OF GIBBS FREE ENERGY BY THE METHOD OF CHERMAK AND RIMSTIDT (1989,1990)
6 GIBBS FREE ENERGIES ESTIMATED FOR URANYL SILICATES FROM SOLUBILITY MEASUREMENTS (NGUYEN ET AL., 1992)
There have been observations that various properties such as volume, bulk modulus refractive index, and, stable-isotope fractionation can be modeled accurately by taking into account basic polyhedral units. Chermak & Rimstidt (1989) used this approach to determine the contributions at 298 K of the free energy component, gi,and the enthalpy component, hi, by multiple regression of the contributions by the polyhedral units [4lA1203, [6lA1203, [61Al(OH)3, 14ISiO2, [6lMg0, [6lMg(OH)2, [6lCaO, [s-ZlCaO, [681Na20, [s-121K20, H20, [6lFe0, [6]Fe(OH)2, and [6lFe203 for 34 minerals. Chermak & Rimstidt proposed that AG? and AH? of many silicate minerals for which there are no experimentally determined thermodynamic properties, or which are too complex and impure to be determined reliably by calorimetric measurements can be estimated at 298 K by summing the contributions of the polyhedral components for these minerals. It should also be noted that for silicate phases for which there are experimentally determined thermodynamic properties, but there are some doubts with respect to accuracy, estimates can be made for comparison.
Nguyen et al. (1992) synthesized four uranyl silicates: soddyite, uranophane, sodium boltwoodite, and sodium weeksite, and measured their solubilities at 303.15 K. Solubilities were measured under Ar gas and the pH controlled by adding either NaOH(aq) or HC104(aq) by pumps interfaced to a pH electrode. The pHs selected for the solubility measurements ranged from 3 to 4.5 in order to increase uranium solubility and to reduce the relative amount of uranyl hydroxide species formed. In this pH range it was assumed that Si02(aq), Na+(aq), and Ca2+(aq) were the predominant silicon, sodium, and calcium aqueous species respectively. The concentrations of the uranyl species were calculated from the solution composition using a speciation program. As an example of their estimates of free energies from solubilities, Nguyen et al. (1992) found that at pH 3 soddyite approached a steady state concentration of uranium of 1.93 x 10-2 molal after about 4 months. The calculated uranyl concentration from speciation considerations was 1.67 x 10-2 molal. The chemical analyzed aqueous silica concentration was found to be 5.12 x 10-3 molal. The ionic strength was listed as 8.43 x 10-2 molal. The Gibbs free energies of formation from the elements for the aqueous uranyl ion, aqueous silica, and liquid water were taken as -953.7, -833.79, and -237.53 kJ mol-l respectively (pers. commun. from W. F. McKenzie to S. N. Nguyen,
In contrast, when the most recent thermodynamic data is utilized, the resulting fit is: XAGI 298
a ebx+ C = 185.360 e-1.026- 6213.506
In this study we have reregressed
AG;298
vs
ge298 at 298 K for the oxides listed in Chermak & Rimstidt (1989) with the exception of K20 and
301
1990). The hydrolysis reaction for soddyite can be & Rimstidt (1989, 1990). Nevertheless there is good written as to fair agreement between free energies estimated by the two methods.[ A < 1 % for kasolite, haiweeite, soddyite, and sodium weeksite (synthetic); A < 2% for boltwoodite, cuprosklodowskite, weeksite, and sodium boltwoodite (synthetic), and A = 3.2, 3.2, It is important to note that a relatively small (= 1 and 4.8% for sodium boltwoodite, sklodowskite, and %) number is added to the sum of the Gibbs free enuranophane respectively]. ergies of the products of the hydrolysis reaction Previously calculated free energies by Heming(aqueous uranyl, aqueous silica, and water) to arrive way (1982) using the methodology of Chen (1975) at the Gibbs free energy for soddyite. One could thus for uranyl silicates for which only two reaction orestimate the free energy for soddyite within about ders were available are herein provisionally recom1% without any solubility measurements if the free mended. The difference between guessing the expoenergies for the hydrolysis products were known. nential fall off from the two reaction orders calcuThis, of course, assumes that soddyite is neither exlated in this study using updated thermodynamic tremely soluble in acid solutions nor extremely indata and that data used by Hemingway is thought to soluble in basic solutions. be small. Nguyen et al. assumed (1) idealized stoichiomeThe free energy calculated for uranophane in this tries for uranophane and sodium weeksite rather study using more recent thermodynamic data and the than their chemical analyzed stoichiometries in their data previously calculated by Hemingway is identicalculations, (2) they adopted approximations for cal. This is coincidental. Recalculating the free enDebye-Hiickel coefficients, (3) they assumed all ergy for uranophane using the data probably used by aqueous uranyl species were accounted for in their Hemingway results in a free energy 8 kJ more posispeciation calculations, (4) they assumed equilibtive. rium was attained, and (5) they assumed uranyl siliProvisionally recommended free energies for socates dissolved congruently. dium boltwoodite (synthetic) and sodium weeksite Assumptions 3, 4 and 5 are likely to contribute (synthetic) are those based on the methodology of the most significantly in departures between estiChermak & Rimstidt (1989, 1990). In the case of mated free energies and those based on solubilities. synthetic sodium boltwoodite there is fair agreement With respect to assumption 4, equilibrium is dif- with the expected exponential fall off (= 20 to 40 ficult to demonstrate owing to prohibitively slow rekcal/mole) from the value calculated by the method action rates, the impossibility of reversing experiof Chen (1975) compared to the value calculated by ments, and the persistence of metastable phases in the method of Chermak and Rimstidt. As for the low temperature experiments. case of synthetic sodium weeksite there is excellent We consider the most seriously erroneous asagreement between the free energy calculated by the sumption (5) to be that dissolution occured congrumethod of Chermak and Rimstidt and the value for ently rather than incongruently. Incongruent solubilthis thermodynamic property estimated from the ity occurs when a mineral upon dissolution reacts to solubility measurements of Nguyen et al. (1992). In form a new solid. This could be the case for the ex- this latter case there is also agreement with the experiments of Nguyen et al. where the amount of ura- pected exponential fall off of the value estimated by nium in solution is constrained by a secondary uraChen’s method and that estimated by the method of nyl mineral that precipitates. In fact this was the case Chermak and Rimstidt. for the sodium boltwoodite experiments as soddyite The reliability of these provisionally recomwas observed to form. As only the uranyl ion and mended Gibbs free energies for uranyl silicates is silica concentrations were measured in these ex- unknown. In the best of worlds, one would have periments, it is impossible to assess what is con- available for a particular thermodynamic property strained without measurements for sodium, potas- multiple estimates made by a variety of methods that sium, and calcium. are in exact agreement with one another as well as with that determined experimentally. This property would be in agreement with a number of phase 7 PROVISIONALLY RECOMMENDED equilibrium experiments and with field observations VALUES FOR GIBBS FREE ENERGIES OF from several localities. Unfortunately this is rarely W N Y L SILICATES the case and one must rely on other judgements to assess and provisionally recommend a particular esProvisionally recommended values for Gibbs free timated value. In many instances one must simply energies of formation for uranyl silicates are in boldrely on the reasonableness of a particular estimation faced type in Table 1. There are unknown and potechnique in that it is in accord with theoretical contentially large errors associated with estimating free siderations and it is generally successful in making energies by the extension of the method of Chermak accurate predictions.
Table 1. Gibbs free energies for uranyl silicates estimated in this study compared to other estimates.
Mineral Uranophane
Ca(H30>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) <-2796' Na(H30)(UO2)(SiO4).H20 2966e -2839d Sodium Weeksite (synthetic) <-90 10' -9093d Na2(U02)02.(Si205)3*4H20
-6113'
Mg(H30)2(uo2)2(sio4)2'4H20
-5896d -9143d -943 I f -9367g -3658e -3658' -3653g
-
-9O8ge
REFERENCES Baes, C.F., Jr. & Mesmer, R.E. 1976. Tlze Hydrolysis of Cations. New York Wiley. Chen, C-H. 1975. A method of estimation of standard free energies of formation of silicate minerals at 298.15 K. American Journal of Science 275: 801-817. Chen, F., Ewing, R.C. & Clark, S.B. 1999. The Gibbs free energies and enthalpies of formation of U6' phases: An empirical method of prediction. American Mineralogist 84: 650-66. Chermak, J.A. & Rimstidt, J.D. 1989. Estimating the thermodynamic properties (AG; and AH; ) of silicate minerals at 298 K from the sum of polyhedral contributions. American Mineralogist 74: 1023-103 1 . Chermak, J.A. & Rimstidt, D.J. 1990. Estimating the free energy of formation of silicate minerals at tugh temperatures from the sum of polyhedral contributions. American Mineralogist 75: 1376-1380.
CODATA Task Group on Key Values for Thermodynamics. 1976. Recommended key values for thermodynamics: 1975. Journal of Chemical Thermodynamics 8: 603-605. Cordfunke, E.H.P. 1964. Heats of formation of some hexavalent uranium compounds. Journal of Physical Cheinistry 68: 3353-3356. Eugster, H.P. & Chou, I-M. 1973. The depositional environments of Precambrian banded iron-formations. Economic Geology 68: 1144-1 168. Finch, R.J. 1997. Thermodynamic stabilities of U(V1) minerals: Estimated and observed relationships, Materials Research Society Proceedings 65: 1185-1192 Gayer, K.H. & Leider, H. 1955. The solubility of uranium trioxide, U03.H20, in solutions of sodium hydroxide and perchloric acid at 25°C. Journal of the American Chemical Society 77: 1448-1450. Hemingway, B.S. 1977. Written communication to Langmuir (1978) US Geol. Survey, Reston, VA. Hemingway, B.S. 1982. Thermodynamic properties of selected uranium compounds and aqueous species at 298.15 K and 1 bar and at higher temperatures-Preliminary models for the origin of coffnite deposits, US Geological Survey, OpenFile Report 82-619,90 p. 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 bars and 0 to 1000°C. Computers and Geosciences 18: 899-947. Langmuir, D. 1978. Uranium solution-mineral equilibria at low temperatures with applications to sedimentary ore deposits. Geochimica Comochimica Acta 42: 547-569. Nguyen, S.N., Silva, R.J., Weed, H.C. & Andrews, J.E., Jr. 1992. Standard Gibbs free energies of formation at the temperature 303.15 K of four uranyl silicates: soddyite, uranophane, sodium boltwoodite, and sodium weeksite. Journal of Chemical Thermodynamics 24: 359-376. Nikitin, A.A., Sergeeva, Z.I., Khodakovsky, I.L. & Naumov, G. B. 1972. Uranyl hydrolysis in a hydrothermal region. Geokhimiya 3: 297-307. Pankratz, L.B. 1982. Thermodynamic Properties of Elements and Oxides. US Bureau of Mines Bulletin 762, 509 p. Robie, R.A., Hemingway, B.S. & Fisher, J.R. 1978. Thermodynamic properties of minerals and related substances at 298.15 K and 1 bar (105 Pascals) pressure and at higher temperatures. US Geological Survey Bulletin 1452, Washington DC, 456 p. Robie, R.A. & Waldbaum, D.R. 1968. Thermodynamic Properties of Minerals and Related Substances at 298.15K and One Atmosphere Pressure and at Higher Temperatures. US Geological Survey Bulletin 1259, Washington DC, 256 p. Robins, R.G. 1966. Hydrolysis of uranyl nitrate solutions at elevated temperatures, Journal of Inorganic and Nuclear Chemistiy 28: 119-123. Rosenzweig, A. gL Ryan, R.R. 1975. Refinement of the crystal structure of cuprosklodowskite, C U [ ( U O ~ ) ~ ( S ~ O ~ O H ) ~ ] .6H20.American Mineralogist 60: 448-453. Stohl, F.V. & Smith, D.K. 1981. The crystal chemistry of the uranyl silicate minerals. American Mineralogist 66: 6 10625. Wanner, H. & Forest, I. (eds) 1992. Chemical Thermodynamics of Uranium. Amsterdam: North-Holland.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Negative pressure and water-mineral interaction in the unsaturated zone of soils L .Mercury UMR-CNRS 8616 "Orsayterre", Universite'de Paris-Sud, Lab. d'klydrologie, bat. 504, 91405 Orsay, France
P.Frey ssinet BRGM, Direction de la Recherche, 3, avenue C. Guillemin, BP 6009,45060 OrlLans cedex 2, France
Y.Tardy Institut National Polytechnique, ENSAT, B.P. 107,Auzeville, 31326 Castanet-Tolosancedex, France UMR-CNRS 6532 "Hydrasa", 40, av. du Recteur Pineau, 86022 Poitiers cedex, France
ABSTRACT: The comprehensive treatment of solution-mineral equilibria in the unsaturated zone of soils requires at first taking into account the specific state of water, generally under negative internal pressure. In a second step, we investigate how it is possible to consider solution-mineral equilibria in such systems. Roles of the pore size and the crystal shape are exemplified, studying the change of the equilibrium constant of the hydration reaction of boehmite into gibbsite. Roles of temperature and negative pressure on solubility are illustrated studying the solubility of gold into its hydroxides complexes. where AP is the pressure difference between water and air in bars, Yw-a the surface tension between water and vapour in humid air, in mJ/m2, rmeniscus The non-saturated domain of soils corresponds to the upper part above the water table (Pw,A,TER=PAIR~~(
1 INTRODUCTION
2.1 Theoretical aspects Along the planar piezometric interface from ground water to atmosphere: PWATER=PATM=I bar. In the saturated zone, below the water table (PWATER>PATM),hydrostatic pressure gains 1 bar about every 10 meters. In the unsaturated zone, above the water table (PWATER
305
meniscus towards air) and in which a grain of a mineral Y precipitates with spherical shape. Figure 2 details some relations existing within such system. Y grain precipitates as a concretion (Tardy 1997) in pores still filled by water, the size of which varies according to the relative humidity in such soil. Y appears to have a convex meniscus towards water (Fig. 2), so that it is at higher pressure than surrounding water.
2.2 Thermodynamic properties of water. Mercury & Tardy (1997a,b), Mercury (1998), and Mercury & Tardy (2001) detailed the expressions required to calculate thermodynamic (enthalpy, Gibbs free energy, heat capacity, entropy, volume) and physico-chemical (water-air surface tension, dielectric constant, pH at neutrality) properties of liquid water at negative pressures. They chose the standard state of water as "pure liquid at any pressure and temperature", proposed by Johnson et al. (1992). As a result, changes in water pressure account for the water state in soils. Thermodynamic features of soil water are calculated as function of water pressure, using an expression extrapolating the molar volume from positive pressure, where it is well documented, to negative one, where data are scarce. Other properties are derived using the classical relations of thermodynamics, enabling discussion of some geochemical features of soil water. Expression (1) implies that stretched water fills only suitable class of pores, so that changes in its properties can be expressed as function of either water pressure or filled pores size. Y ~ is- calculated ~ as proposed by Mercury & Tardy (1997b, 2001), permitting to deduce the radius of the meniscus at any water pressure. It has to be highlighted that the radius of the pore is equal to that of the meniscus if the contact angle with pore walls is nil, which is a present paper assumption. Following this theoretical frame, all the properties of minerals and aqueous species designed for positive pressures are extrapolated to negative pressures. Aqueous ions are assumed to be at the same pressure than the surrounding water. Their thermodynamical properties with pressure are calculated using the THIS model (Tanger & Helgeson 19SS), supposed suitable to the negative pressure domain. As we deal with diluted solutions there are neither osmotic effects nor change of the activity coefficients. Thus solution-mineral equilibria models are applicable to the unsaturated domain, where PWATER
Figure 2. Equilibrium conditions in soil water - minerals systems. Pressure differences and radius of interfaces implied in such representation are reported.
3 BOEHMITE-GIBBSITE IN SOILS: A CASE STUDY
Considering the gibbsite-boehmite equilibrium conditions in tropical soils, particularly in the case of metabauxites, Tardy (1997) has shown that almost amorphous boehmite develops in concretions at the expense of very well crystallized gibbsite in excretions, when temperature is sufficiently high and/or the relative humidity is sufficiently low (<72%). After thermodynamic properties of gibbsite and boehmite from Sposito (1996), close to those of Pokrovskii & Helgeson (1995), Gibbs free energy of hydration reaction (AlOOH+H20=Al(OH)3) is positive at 25OC, 1 bar: A~G"=+692 kJ.mol-', whereas the volume of reaction is negative: A~v"=-0.565 J.bar-'.mol-'. This means that thermodynamic equilibrium is reached for positive pressure, namely soil atmosphere at around 130%. In other words, gibbsite would never appear in tropical soils. In order to satisfy observed natural conditions, Trolard & Tardy (1987, 1989) proposed to change the Gibbs free energy of boehmite, of less crystallinity, without changing that of the wellcrystallized gibbsite, leading to A~G'=-800 kJ.mol-', and equilibrium conditions at R.H.=72%. This approach can be translated into the present pressure approach by the following: R.H.=72% means P~20=-437bars; dAfGoAIOo@-l 495 J.mol-' means P ~ 1 0 0 @ - 7 6bars; 6 dAfi0~l(o~)3'0J.mol-' means P A , ( O H ~1 bar.
2.3 Mineral thermodynamics in capillaries. Thermodynamic properties of minerals in capillary systems of the unsaturated zone of soils depend on the surface energy and the pressure difference between water and minerals, and also on the geometry of the fluid-solid interface. The geometry of the system is clearly determinant to a correct thermodynamic treatment of mineral equilibria in soil. The simplest scheme we could assume is to suppose a cylindrical capillary, made of a mineral X, where capillary water takes place (with concave
306
These data allow us to fit a stability domain for natural gibbsite in a R.H. range from 72% to 100%. Moreover, pore radius (delimited by gibbsite pore walls, see Fig. 2) can be calculated by iterative method from surface tension between water and vapour in humid air (Mercury & Tardy 2001). Supposing the radius of boehmite grain equal to that of the pore (Fig. 2), we can deduce surface tension between bad-crystallized boehmite (dA@"AloOH>O) and stretched water (PH20-d bar), using Laplace law (expression (l)), which here can be written:
considerations about the role of relative humidity and mineral crystallinity governs the equilibria of goethite-haematite in ferricretes and gibbsitekaolinite in laterites.
4 SOLUTION-MINERAL NEGATIVE PRESSURE
EQUILIBRIA
AT
After considering hydration reactions in presence of soil water (solid-solid equilibria), this paper deal with dissolution reactions (solid-solution equilibria), with the case study of the dissolution of gold in tropical soils. Native gold dissolves as monovalent dissolved gold, even in oxydizing conditions (~02'1 O-o.68 bar). Redox conditions in soils are here assumed to be in equilibrium with p O ~ = l 0 -bar. ~ ~ Dissolved gold species are cpsidered to be hydroxides complexes (Au+, AuOH , AuO-), the thermodynamic properties of which are taken from Shock et al. (1997). The solubility is here calculated as the sum of their concentrations, resulting from the dissolution of pure gold in soil conditions (Fig. 4). Figure 4 shows that increasing temperature has a dissolving effect at any pH, at odds with the usual dehydrating heating effect. Pressure effect depends
Figure 3. Equilibrium constant of boehmite hydration as function of pressure (bold figures, left y-axis) and temperature (normal figures, right y-axis).
-10
-1 1
.
spherules at 1 bar
-gold
-1 2
Figure 3 shows the change of the equilibrium constant of the hydration reaction of boehmite into gibbsite. Increasing temperature as well as decreasing pressure favours boehmite, which is the less hydrated pole of reaction. It has to be noted that this corresponds to the expected trend due to the "ice-like" nature of the hydration water (Tardy et al. 1999, Mercury et al. 2001). At lower pressure (arid zone), water subsists only in minor pores of soil, where boehmite should predominate as spherule minerals. At higher pressure (humid zone), larger pores of soils should display mainly gibbsite as large crystals. This corresponds clearly to natural observations, synthetized by Tardy (1997) in terms of concretionning (precipitation of fine solid in minor pore) and excretionning (large solids precipitating in larger pores) regimes. It appears that the grain size of the boehmite spherules, limited by that of the pore, the interfacial energy, the pressure difference among the system, satisfy our previous geometric model, leading us to take into consideration the petrographic shapes as input data for the thermodynamic treatment of solution-mineral interactions. This model may be tested in another cases since Trolard & Tardy (1987, 1989), for example, demonstrated that same kind of
w
-13
I
3"
-14
Y
-15
s I
-1 6 -1 7 T = 25OC
-1 8 -19 I -1 0
:
gold spherules at 50°C - - gold spherules at 25°C
-1 1 -12
,,.
/ 0 , /
:
:
...,.' ,/'
-13 -14
-1 5 -1 6
-17 -1 8
P = 1 bar
-19 2
3
4
5
7
6
8
9 1 0 1 1
PH Figure 4. Solubility of spherules of gold as function of water pressure (top) and temperature (bottom).
307
on the pH range. At low pH (pH<3.5), decreasing pressure favours gold solubility. Following the previous paragraph, this lead to predict that fine grains of gold dissolve in acidic capillary solution, suspended in minor pores. At higher pH, (pH>3.5), decreasing pressure (=drying air) favours precipitation of fine grains of gold in minor pores. Finally, comparison of the two parts of figure 4 illustrates that increasing temperature and decreasing pressure act in opposite manner for gold, except in restrained range of low pH.It has to be noted that increasing temperature from 25" to 50°C at 1 bar gives the same solubility shift (but with opposite sign) than decreasing water pressure from 1 bar (soil humidity=lOO%) to -2 kbar (soil humidity=22%) at 25°C. In other words, gold dissolving effect of heating (25"-+5OoC) would be exactly compensated by the precipitating effect of air dessication (100+22%) in the 3.5-1 1 pH range. For lower pH, the coupling of aridity and heat would be strongly dissolving. This kind of behavior change under temperature and pressure has to be taken into consideration where one accounts for unsaturated systems. Reverse behaviours are possible with respect to the expected one, where increasing temperature and decreasing pressure decrease solubility (Mercury & Tardy 2001).
5 CONCLUSION The "negative pressure approach" developed here allows us to take into account two new parameters in the calculations of equilibrium thermodynamics. 1. The hydric state of soil-air-water system of the unsaturated zone, where water undergoes a lower pressure than that of the atmosphere, which is simultaneously desiccated with respect to saturated air. 2. The crystalline morphology gives a pressure significance to the equilibrium shape, which may be a mean for minerals to persist in such system. Equilibrium solubilities calculated with classical methods, and including these new parameters, account correctly for the general trend of the studied mineral despite the present simplifications. Dissolving effects of stretched water have to be considered (despite the usual trend due to water itself) depending on the nature of studied minerals and ions (gibbsite, boehmite, goethite, haematite, kaolinite, ... and their different aqueous ions, simples and complexes). Reverse behaviours would exist at negative pressures, as those well known at positive ones (retrograde minerals). Thermodynamic treatment of the solution-mineral equilibria in soils has to consider the role of water pressure, pore sizes and crystals shape on the equilibrium constant of hydration and diszolution reactions.
REFERENCES Johnson, J.W., Oekers, E.H. & H.C. Helgeson 1992. SUPCRT92: a software package for calculating the standard molal thermodynamic properties of minerals, gases, and aqueous species from 1 to 5000 bars and 0" to 1000°C. Comput. Geosciences 18: 899-947. Mercury, L. 1998. Solution-mineral equilibria in the unsaturated zone of soils. Thermodynamics of capillary water and physico-chemical modelisation. Doc. BRGM 283, Orleans: BRGM ed. Mercury, L. & Y. Tardy 1997a. Negative pressure and thermodynamic properties of capillary water. A concise review paper. C. R. Acad. Sci. Paris 324: 863-873. Mercury, L. & Y. Tardy 1997b. Physico-chemical features of water in capillaries and fog water droplets. C. R. Acad. Sci. Paris 325: 947-954. Mercury, L., Vieillard, P. & Y. Tardy 200 1. Thermodynamics of ice polymorphs and "ice-like" water in hydrates and hydroxides. Appl. Geochem. 16:161-181. Mercury, L. & Y. Tardy 200 1. Negative pressure of water in soil. Geochemical consequences. Geochim. Cosmochim. Acta (in press). Pokrovskii, V.A. & H.C. Helgeson 1995. Thermodynamic properties of aqueous species and the solubilities of minerals at high pressures and temperatures: the system A1203-H20-NaC1.Am. J. Sc. 295: 1255-1342. Shock, E.L., Sassani, D.C., Willis, M. & D.A. Sverjensky 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. Sposito, G. (Ed.) 1996. The environmental chemistry of aluminum. 2nded., Lewis Pub. Inc. Tanger IV, J.C. & H.C. Helgeson 1988. Calculation of the thermodynamic and transport properties of aqueous species at high pressures and temperatures: revised equations of state for the standard partial molal properties of ions and electrolytes. Am. J. Sc. 288: 19-98. Tardy, Y., Mercury, L., Roquin, C. & P. Vieillard 1999. The concept of ice-like water: hydration-dehydration of salts, hydroxides, clay minerals, and living or inert organic matter. A concise review paper. C, R. Acad. Sci. Paris 329: 377-388. Tardy, Y. 1997. Petrology of laterites and tropical soils. Rotterdam: Balkema. Trolard, F. & Y. Tardy 1987. The stabilities of gibbsite, boehmite, aluminous goethites and aluminous hematites in bauxites, ferricretes and laterites as a hnction of water activity, temperature and particle size. Geochim. Cosmochim. Acta 5 1 : 945-957. Trolard, F. & Y. Tardy 1989. A model of Fe3'-kaolinite, AI3+goethite, A13'-hematite equilibria in laterites. Clay minerals 24: 1-21.
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Wafer-Rock lnferacfion 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Solubility of Na, Al, and Si in aqueous fluid at 0.8-2.0 GPa and 1000-1300°C B .O .Mysen Geophysical Laboratory and Centerfor High-pressure Research (CHiPR), Washington DC, USA
K.Wheeler Department of Geological Sciences, Brown University, Providence RI, USA
ABSTRACT: The silicate solubility in aqueous fluid in melt+fluid systems with 80 mol % SiO,, 20-14 mol % Na,O, and 0-6 mol % A1,0, at 0.8-2.0 GPa and 1000"-1300"C is in the 3-20 mol% range with a positive temperature- and pressure-dependence. The solution behavior in aqueous fluid is incongruent. The partition coefficient, Doxidefluidmelt, ranges between 0.02 and 0.5 for Na,O, 0.01 and 0.2 for A1,03, and 0.01 and 0.4 for SiO,, and increases linearly with increasing oxide content of the fluid (decreases with oxide content in the melt). Partial molar volume of H,O in the silicate-saturated fluids, VHfOfluid, is in the range 17-25 cm3/mol. Compared with the molar volume of pure H,O, VoH20, the V, OflUid values are 10-15 % lower at 0.8 GPa and as much as 15 % higher than VoH20 at 2.0 GPa. This volume difference diminishes as the system become more aluminous.
I INTRODUCTION
2 EXPERIMENTAL METHODS
Evidence for chemical interaction between rocks and aqueous fluids in the crust and the upper mantle of the Earth is extensive. High-grade metamorphism in deep continental crust can involve dehydration of hydrous phases with concomitant alteration of major and trace element patterns (e.g., Rollinson & Windley 1980, Fowler 1984, Whitehouse 1989). There is also evidence for transport of major, minor, and trace elements from the subducting slab to the overlying mantle wedge near convergent plate boundaries (e.g., Riter & Smith 1996, Brenan et al. 1995). Alteration of the bulk chemistry of peridotite in the mantle wedge above subducting plates via ingress of aqueous fluid can result in formation of biotite and K-richterite (Luth 1997, Konzett & Ulmer 1999). In order to describe these interaction processes, the solubility behavior of silicate components in aqueous fluids as a function of pressure, temperature, and bulk chemical composition is needed. Major important chemical components in upper mantle subduction zones and in the Earth's crust are alkalis, alumina and silica. An experimental study aimed at characterization of silicate-saturated aqueous fluids was conducted, therefore, in the system Na,O-Al,O,Si0,-H,O in the 0.8-2.0 GPa and 1000"-1300°C pressure and temperature range, respectively.
Starting compositions were synthetic glasses that contained 80 mol % SiO, with 20, 17, and 14 mol % Na,O and 0, 3, and 6 mol % A1,0, (denoted NS4, NS4A3, and NS4A6, respectively). The glasses were crushed to L. 20 ym grain size and stored at 110°C when not in use. High-pressure and high-temperature experiments were conducted in the solid-media, high-pressure apparatus in 0.75"-diameter furnace assemblies. Temperatures were measured with Pt-Pt,,Rh!, thermocouples with no correction for pressure on their emf. Pressure was calibrated against the melting point ofNaCl and the calcite-aragonite transformation. The uncertainties were 210°C and +O. 1 GPa, respectively. Two sample container configurations were employed. To determine the bulk silicate solubility in aqueous fluid the starting glasses were loaded together with double-distilled, deionized H,O into the Pt containers and welded shut while submerged in liquid N,. The bulk silicate solubility in aqueous fluids was determined as described by Mysen and Acton ( 1999). In order to determine NayAl, and Si distribution between melt and fluid, a double capsule technique was used. The starting material was loaded into perforated (-50 200 ym diameter holes) 3 mm OD Pt
309
capsule and welded shut. This capsule was loaded into an outer 5 mm OD Pt capsule to which the desired amount of H,O was added. This capsule was then welded shut while submerged in liquid N,. After an experiment, the outer capsule was opened and all silicate material that precipitated from the fluid during quenching was washed off the outside surface of the inner capsule. The inner capsule was then opened, and the hydrous melt extracted. The glass compositions were determined with the electron microprobe (JEOL 8900 operating at 15 kV and 10 nAmp). Pristine aqueous fluid cannot be extracted because of extensive precipitation of aluminosilicate materials from the fluid during quenching (Mysen and Acton 1999), and fluid composition was obtained by mass-balance from the composition of the hydrous, quenched glass and the bulk chemical composition. Attainment of equilibrium was assessed by time studies. Fluidmelt partition coefficients, determined at 1100°C,reached constant values after about I 100 min. Experiments were, therefore, conducted with durations of 1440 min or more. 3 RESULTS AND DISCUSSION
Figure 2. Bulk silicate solubility in aqueous fluids with temperature for compositions with 0 mol% (NS4) and 6 mol % (NS4A6) A1,0,.
The silicate solubility in aqueous fluid, Xsilicate, increases linearly with increasing temperature, T, for all compositions (Fig. 2), consistent with observations from other systems (Pascal & Anderson 1989, Manning 1994, Davis 1972, Shen & Keppler 1997). The (dXsillicate/dT)P for equivalent compositions in the system K,0-A1,03-Si0,-H,0 in the same pressure and temperature range is also positive (Mysen &Acton 1999). The (dXsilicate/dT)P in the K,O-Al,O,-Si0,-H,O system is, however, smaller than for compositions in
Figure 1. Bulk silicate solubility in aqueous fluids with pressure for compositions with 0 mol% (NS4) and 6 mol % (NS4A6) A1,0,.
The bulk silicate solubility in aqueous fluid is a positive and non-linear function of pressure (Fig. l), in qualitative accord with data for other systems such as SiO (Manning 1994)and NaAlSi,O, (Davis 1972). The pbsitive dependence of solubility on pressure becomes more pronounced as the pressure increases. Compared with the results for equivalent compositions in the system K,O-AI,O,-Si0,-H,O (Mysen & Acton 1999), the silicate solubility in the Na2O-A1,0,-SiO,-H,O system is generally slightly smaller for the most AI-rich composition (KS4A6 versus NS4A6), whereas for 0 and 3 mol % A1,0,, the silicate solubility in the aqueous fluid in the Na,OA1,0,-SiO2-H,O system is slightly higher than in the K2O-A1,0,-SiO,-H,O system at the same pressure and temperature.
Figure 3. Partial molar volume of H,O in silicate-saturated aqueous fluid, VHZOflUid, as a hnction ofpressure and temperature indicated. Also shown is the molar volume ofpure H 2 0(dashed lines; Haar et al. 1984).
the Na20-A1,0,-Si0,-H20 system. From stepwise regression, of the solubility data can be expressed with the equation: Xsilicate= 1.9k3.5-1 .3+0.9*X,1203+0.008t-0.002eT + 13t-4.P + 7.3k1.5P2. R2 = 0.89. (1)
310
Figure 4. A. Composition of coexisting silicate-saturated aqueous fluid (open symbols) and H,O-saturated silicate melt (solid symbols) projected from H,O. B. Enrichment of Na, Al, and Si in aqueous fluid coexisting with melt relative to their concentration in starting composition (enrichment factor) (NS4A3 - 3 mol % A1,0, as example) as a function of mol fraction of aqueous fluid [fluid/(fluid+melt)] in experimental charges.
Figure 5. Distribution of oxides between coexisting, silicatesaturated aqueous fluid and H,O-saturated silicate melt as a function of oxide concentration in fluids and melts as indicated. 0% A1,0, etc. denotes mol % A1,0, in starting composition.
increase in /3 [= - lN0(dV fluid/%') ] from 0.14 1GPaat 1000°Cto 0.172 G$;-? at 130n"C. For NS4A6saturated aqueous fluid p increases from 0.193 GPaI to 0.213 GPa-' in the same temperature range. The solubility of Na, AI, and Si in the aqueous fluids coexisting with aluminosilicate melts is incongruent so that calculated on a water-free basis, the fluids are enriched in Na and depleted in Si relative to the starting compositions (Fig. 4). Consequently, the abundance of Na, Al, and Si in coexisting fluids and melts depends on the fluidlmelt ratio for each of the compositions and at each pressure and temperature. Such a relationship, in turn, implies that unless these elements in melts and fluids show Henrian solution behavior, the fluid/melt partition coefficients, Doxidcflu'd/melt, depends on melt and fluid composition. This behavior is exactly what is observed (Fig. 5). For the most part, Doxid~uid/me'r is a nearly linear and positive hnction of the oxide is a negative concentration in fluid. The Doxldefluld/melt
In equation (l), XA1203 and Xsilicate are in mol %, T is "C and P is GPa. The solubility data for silicate in aqueous fluid can be used to derive partial molar volumes of H 0 in the fluid. The free energy of solution of silicate 1s; AGT(P)= 0 = AGT(1 bar) + RTln(a,20fluid/fH2,)+ V,20""'(P- 1) (2) is activity of H,O in the fluid,f,200 is where aH20flUld the fugacity of pure H,O (Haar et al. 1984), and VH20fluid is the partial molar volume of H,O in the fluid. The relationships between V, and pressure is linear and negative (Fig. 3). For Al-bearing is compositions the pressure-dependence of V also a function of temperature so that (dV,,:%dP) increases with increasing temperature. For the NS4A3 composition, this trend translates to an 311
function of the oxide concentration in the melt (Fig.
the Earth's crust and upper mantle. N. Nb. Mineral. 172: 227-244. Mysen, B. 0. & M. Acton 1999. Water in H20-saturated magma-Fluid systems: solubility behavior in K,OAl,O,-SiO,-H,O to 2.0 GPa and 1300°C. Geochim. Cosmochim. Acta 63: 3799-3816. Pascal, M. L. & G. M. Anderson 1989. Speciation of Al, Si, and K in supercritical solutions: experimental study and interpretation. Geochim. Cosmochim. Acta 53: 18431856. Riter, J. C. A. & D. Smith 1996. Xenolith constraints on the thermal history of the of the mantle beneath the Colorado Plateau. Geology 24: 267-279. Rollinson, H.R. & B. E Windley 1980. Selective elemental depletion during metamorphism of archaean granulites, Scourie, NW Scotland. Contrib. Min. Pet. 72: 257-263. Shen, A. H. & H. Keppler 1997. Direct observation of complete miscibility in the albite-H,O system. Nature 385: 710-712. Whitehouse, M. J. 1989. Pb-isotopic evidence for U-Th-Pb behaviour in a prograde amphibolite to granulite facies transition from the Lewisian Complex of north-west Scotland; implications for Pb-Pb dating. Geochimica et Cosmochimica Acta 53: 7 17-724.
5).
The incongruent solubility behavior of Na, Al, and Si in aqueous fluids probably reflects differences in the structure of coexisting H,Osaturated silicate melts and silicate-saturated aqueous fluids. From in-situ structural data (Mysen 1998), the main difference between the two phases is that in aqueous fluids, there probably are H,O solvation spheres around alkali-OH complexes, whereas there is no evidence for solvated H,O in the melts. With increasing pressure, the partition coefficients do, however, approach 1. This effect may be due to a shrinkage of the H,O solvation spheres in the fluid. This trend, when extrapolated to higher pressure, requires that there exists a pressure above which the system becomes supercritical. This suggestion is in agreement with in-situ observations of a transformation from coexisting aqueous fluids and silicate melts to a single, supercritical phase at pressures somewhere above 1 GPa (Shen & Keppler 1997). The pressuretemperature coordinates of the melt-fluid solvus in the latter experiments is, however, uncertain. On the basis of the present experiments solvus closure must occur above 2 GPa. REFERENCES Brenan, J. M., H. F. Shaw, R. J. Ryerson & D. L. Phinney 1995. Mineral-aqueous fluid partitioning of trace elements at 900" and 2.0GPa: Constraints on the trace element geochemistry of mantle and deep crustal fluids. Geochim. Cosmochim. Acta 59: 333 1-3350. Davis, N. F. 1972. Experimental studies in the system sodium alumina trisilicate-water:Part I. The apparent solubility of albite in supercritical water. Ph. D. thesis The Pennsylvania State University, State College, PA. Fowler, M. B. 1984. Large-ion lithophile element mobility in the lower continental crust; mineralogy and geochemistry of the hornblende-granulite subfacies at Gruinard Bay and its relationships with amphibolitefacies and granulite-facies end members. In M. Brown, (ed.), Metamorphic Studies; Research in Progress. 141. Geological Society of London, London, United Kingdom. Haar, L., J. S. Gallagher & G. S. Kell, 1984. Steam Tables. Th er m o dy n a m i c a n d Tra n sp o r t Properties and Computer Programs f o r Vapor and Liquid States of Water in SI Units. 320 pp. Hemisphere Publishing Corporation, New York. Konzett, J. & P. Ulmer 1999. The stability of hydrous potassic phases in lherzolitic mantle - an experimental study to 9.5 GPa in simplified and natural bulk compositions. J. Petrol. 40: 629-652. Luth, R. W. 1997. Experimental study of the system phologopite-diopside from 3.5 to 17 GPa. Amer. Mineral. 62: 1198-1209. Manning, C.E. 1994. The solubility of quartz in H,O in the lower crust and upper mantle. Geochim. Cosmochim. Acta 58: 483 1-4840. Mysen, B. 0. 1998. Interaction between aqueous fluid and silicate melt in the pressure and temperature regime of
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The surface chemistry of a gram-negative bacteria and its role in metal uptake B .T.Ngwenya Department of Geology & Geophysics, University of Edinburgh, West Mains Road, Edinburgh EH9 3JW
I .W.Sutherland Institute of Cell & Molecular Biology, University of Edinburgh, Mayfield Road, Edinburgh EH9 3JH
ABSTRACT: This paper presents results of acid-base titrations of a gram-negative bacterial surface, which are then used to model metal adsorption edges to the surface. Modelling of the titration data using FITEQL 4 reveals three distinct surface sites with intrinsic pK values of 4.3k0.1, 6.910.5 and 8.9k0.5. Metal adsorption experiments involving Pb, Cu and Zn show that at least two of these are involved in metal uptake. The concentrations of the different surface sites, as well as metal-stability constants differ slightly from those of other gram-negative and from gram-positive bacteria. These differences may originate from inherent variability in surface characteristics of individual species andor cell wall composition between gramnegatives and gram-positives. However, the small differences suggest that it may be possible to use a suite of microbial consortia in modelling water-rock interaction.
1 INTRODUCTION Bacteria are cornmon in many geologic environments and are involved in a variety of fluidrock reactions, including mine water acidification (Walsh & Mitchell 1972), mineral dissolution and growth (Thorseth et al. 1995), and immobilisation of heavy metals (Mackaskie & Dean 1987). It is therefore important that bacterial effects are incorporated in models of water-rock interaction and contaminant transport. However, our understanding of the mechanisms underlying these processes is still poor. For example, the mechanism by which bacteria promote mineral dissolution rates remains debatable (e.g. Ullman et al. 1996). Moreover, although bacterial immobilisation of heavy metals is well known, the determination of intrinsic thermodynamic parameters for modelling metal adsorption is relatively new (Plette et al. 1995, Fein et al. 1997). Recent studies suggest that metal-bacteria complexes for gram-positive soil bacteria have thermodynamic stabilities comparable to those of common organic ligands (Plette et al. 1996; Fein et al. 1997, Daughney et al. 1998). In contrast, relatively few studies (e.g. Seki et al. 1998, Texier et al., 1999) have examined the metal adsorption characteristics of gram-negative bacteria. The latter not only have less peptidoglycan in their cell wall but the peptidoglycan is covered in an outer 313
membrane with different surface polymers (Beveridge 1989). Since gram-positive strains exhibit different surface chemistry (Daughney et al. 1998), differences in cell wall composition are likely to lead to larger variability in surface reactivity. We test this hypothesis by conducting experiments on a gram-negative soil bacterium and comparing the results to previous studies. The results suggest common functional groups with similar deprotonation constants but site concentrations and metalstability constants appear to be species-dependent. 2 METHODS A copper-resistant strain of gram-negative bacteria was isolated from a diesel contaminated soil by incubating the soil for three weeks in a 2mM copper medium containing 30g/L tryptone soya broth of which 0.5% was yeast extract. A series of biochemical tests showed the strain to belong to the family Enterobacteriacea, and full identification to species level is currently undenvay. Experimental batches were grown from the primary culture by inoculating 1L of medium containing lg/L each of yeast extract and casein hydrolysate, plus the following salts: 1Og Na2HP04, 3g KH2P04, l g K2S04, l g NaClYO.2gMgS04.7H20, 0.001g CaC12.2H20 and 0.0001g FeS04.7H20. The cultures were incubated at 30°C for 48 hours and
harvested by centrifugation at 15,OOOg for 20 minutes at 4°C. The cells were washed 3 times in distilled water, washed in 0.1M HN03 for one hour to strip them of metals adsorbed from the growth medium, then rinsed four more times in distilled water. Finally, the cells were freeze-dried to yield a dry powder that was used in the experiments. Tests showed that the use of wet and dry cells did not affect the results. For the actual experiments, the cells were further soaked for one hour to re-hydrate them, and then washed three times in the electrolyte. Deprotonation constants and surface site concentrations were determined from acid-base titrations in a background electrolyte of 0.01M NaNO3. Replicate 50 ml suspensions of varying bacterial concentration were titrated with 0.5M NaOH using an automated Mettler Toledo DL53 burette assembly. The titrations were carried out in polythene beakers sealed to atmospheric gases and a positive Nz pressure was maintained throughout the titration. Sorption experiments were conducted as a function of pH using 25ml suspensions. The metal was added from 1000 mgL-' stock solutions to yield initial concentrations of 10k0.5 ppm. The metalbacterial suspension was equilibrated for 30 minutes (time established from kinetic experiments). Sampling involved pelleting the cells and pipetting lOml of the supernatant into an acid-cleaned bottle. These solutions were acidified to 1% v/v HN03 and stored at 4°C before metal analysis by Flame AAS using matrix-matched standards.
npnii where Ycalc is the calculated value, Y e p is the experimental value, Sexpr is the error associated with the experimental data, np is the number of data points, nu is the number of adjustable parameters in l the number of so called Group I1 the model and n ~is components, for which both the total and free (dissolved) concentration are known. The variance is thus a quantitative estimate of the success with which the specified model describes the data such that 0.1 IV(Y) 5 20 is considered a good fit (Herbelin & Westall 1999). It follows from (1) that as the number of adjustable parameters increases (more sites) the variance should increase unless the model with more parameters provides a better description of the data.
3 EESULTS AND DISCUSSION
3.I Surface Titrations Figure 1 shows a typical surface titration for a 21.3gL-' suspension of the bacteria, together with the titration for the electrolyte alone. The presence of bacteria imparts a significant buffering capacity to the suspension, due to the presence of active functional groups on the bacterial surface. It is also clear, however, that the bacterial suspension shows relatively weak (if any) inflection points, making determination of deprotonation points by conventional methods difficult. Instead, FITEQL 4 (Herbelin & Westall, 1999), an equilibrium speciation program was used to fit each titration in order to determine the number of sites, their deprotonation constants and their concentrations. The program fits data to a specified model by calculating a misfit function or variance, V(Y) between the experimental data and the model:
Figure 1. Titration of blank and bacterial suspension shows that hnctional groups on the surface of the bacteria buffer the electrolyte pH.
In the model, deprotonation represented generically as
reactions were
R-OH%R-O-+H+ (2) where R - 0 2 represents a protonated surface site. A constant capacitance model (Stumm & Morgan, 1996) was used to correct for changes in surface potential (w) due to development of surface charge (0) via:
314
(3)
where C is the capacitance. With this approach, intrinsic deprotonation constants (Kjnt) can be calculated from apparent equilibrium constants ( K ) for reaction (2) above via:
where F is the Faraday constant, A2 is the change in the surface charge of the surface species while R and Tare the gas constant and temperature respectively. Data from each titration was fitted successively with one, two, three, four and five sites and the variance used to judge the best fitting model. Reactions for the dissociation of water and the acidbase behaviour of the electrolyte were included in the equilibrium model, while activity coefficients were calculated using the Davies equation (Stumm & Morgan 1996). A surface area of 140 m'g-' of dry bacteria was assumed based on computations of Fein et al. (1997) and a surface capacitance of 8 Fm-2was found to provide good fits to the data. Adjustable parameters were intrinsic deprotonation constants and surface site concentrations. Whilst only models with one, two and three sites converged, the best fit required a three-site model, as found for other strains of bacteria (Plette et al. 1996, Fein et al. 1997, Seki et al. 1998). MeanpKint values (n = 6) for these sites were 4.3k0.1, 6.9f0.5 and 8.9t-0.5, while the site concentrations were 5.0?0.7x1O4, 2.2f0.6 x104 and 5.5f2.2 ~ 1 0mol/g ' ~ dry bacteria. The errors are one times the standard deviation of the mean. By comparing our pKint values with those from previous studies and deprotonation constants for well known functional groups, we assign sites with PKint = 4.3 and 6.9 to carboxyl and phosphoryl groups respectively. The third value is on the lower end of those reported by other workers and may be due to hydroxyl or amine groups (Plette et a1.1996; Fein et al., 1997 and Cox et al., 1999). Although of similar order of magnitude, the site concentrations for these bacteria differ slightly from those for grampositive bacteria reported by Fein et a1 (1997) and from those reported by Seki et a1 (1998) on gramnegative bacteria. It is not clear whether these differences reflect species differences or growth conditions.
metal adsorption reaction, assuming that the deprotonation reaction (2) occurs prior to metal adsorption thus:
R - 0- + M"+
R -OM("-')
(5)
where Wf was Pb2', Cu2+ and Zn2'. The site deprotonation constants and concentrations determined from titration data were used as input, as well as electrolyte and water acid-based equilibria. In Figure 2, the lines represent FITEQL generated curves assuming that adsorption occurs to one (Fig. 2a) andor two (Fig. 2b) sites. As with the titration modelling, the variance was used to select the bestfitting model, in conjunction with visual adherence of the model curves to the experimental data (Fig. 2).
3.2 Metal adsorption experiments
Figure 2 . Representative metal adsorption data (symbols) and FITEQL modelling (curves), showing that one site model fits the data for high bacterial concentrations, but two sites are required at low bacterial concentrations.
As expected, the amount of metal adsorbed increases with increasing pH (Fig. 2). These experimental adsorption edges were also analysed using FITEQL 4 to calculate site-specific stability constants for the
For all three metals, the one site model was adequate for high bacterial suspension concentrations, but low bacterial suspensions required two sites to fully fit 315
the data. These observations suggest that site concentration, which is dependent on weight of bacteria, is a primary control on adsorption. Fits to Pb data were less satisfactory, and only one-site models converged. In contrast, Fein et al. (1997) found that Cu2+ adsorption could be described by adsorption to one site regardless of bacterial concentrations, while Pb2+required two sites at low bacterial concentrations. This discrepancy requires further investigation. FITEQL calculations are consistent with the following stability constants for the different metals; Log &&,oxyl: Zn2+ = 3.3f0.1, Pb2+= 4.320.2, and c 8 = 4.4k0.2, Log Kphosphoryi: Zn2+ = 5.1k0.1 and Cu2' = 6.050.5. Because the two-site model did not converge for Pb data, the Pbphosphoryl stability constant could not be evaluated. Note that carboxyl stability constants for Cu2+and Pb2+are close to those for the gram-positive Bacillus subtilis from Fein et a1 (1997), but the Pb2+constant differs markedly from those reported by Seki et al. (1 998) for two gram-negative strains. Thus published studies so far have yet to resolve whether the differences observed are due to the inherent differences between species (Daughney et al. 1998), cell wall composition or experimental artefacts associated with differences in growth media.
4 CONCLUSIONS We have studied the acid-base behaviour of the surface of a gram-negative soil bacterium and used the calculated deprotonation constants to explain metal adsorption edges to the bacterial surface. The titration data is consistent with the presence of three distinct surface groups with pK values of 4.3k0.1, 6.9f0.5 and 8.920.5. At least two of these are involved in metal uptake at the pH conditions expected in nature. More importantly, we found small differences in the concentrations of the different surface sites and metal-stability constants from those of other gram-negative and from grampositive bacteria. The small differences between strains suggest that it may be possible to use a suite of microbial consortia in modelling, although further studies are required to confirm this inference. REFERENCES Beveridge, T.J. 1989. Metal ions and Bacteria. In Beveridge, T.J & Doyle, R.J. (eds.). Metal Ions & and Bacteria, 1-29. Cox, J.S., Smith, D.S., Warren, L.A. & F.G. Ferris 1999. Characterising heterogeneous bacterial surface fbctional groups using discrete affinity spectra for proton binding. Environ. Sci. Technol. 33 : 45 14-4521.
31 6
Daughney, C.J., Fein, J.B. & N. Yee 1998. A comparison of the thermodynamics of metal adsorption onto two common bacteria. Chem. Geol.. 144: 161-176 Fein, J.B., Daughney, C.J., Yee, N. & T.A. Davis 1997. A chemical equilibrium model of metal adsorption onto bacterial surfaces. Geochim. Cosmichim. Acta. 61 : 33 193328. Herbelin, A.L. & J.C. Westall 1999. FITEQL 4.0: a computer program for determination of chemical equilibrium constants from experimental data; Report 99-0 1, Department of Chemistry Oregon State University, Corvallis. Mackaskie, L.E. & A.C.R. Dean 1987. Use of immobilised biofilm of Citrobacter sp. For the removal of uranium and lead from aqueous flows. Enzyme Micobiol. Technol. 9: 213. Plette, A.C.C., Van Reimsdijk, W.H., Benedetti, M.F. & A. Van der Wal 1995. p H dependent charging behaviour of isolated cell walls of a gram-positive soil bacterium. J. Colloid Interface. Sci. 173: 354-363. Plette, A.C.C., Benedetti, M.F. & W.H. Van Reimsdijk 1996. Competitive binding of protons, calcium, cadmium and zinc to isolated cell walls of a gram-positive soil bacterium. Environ. Sci. Technol. 30: 1902-1909. Seki, H., Suzuki, A. & S-I. Mitsueda 1998. Biosorption of heavy metals on Rhodobacter sphaeroides and Alcaligenes eutrophus H16. J. Colloid Interface. Sci, 197: 185-190. Stumm, W. & J.J. Morgan 1996. Aquatic Chemistry, Third Edition, Wiley, Chichester, 1022pp. Texier, A.C., Andres, Y & P. Le Cloirec 1999. Selective biosorption of lanthanide (La, Eu, Lu) ions by Pesudomonas aeruginosa. Environ. Sci. Technol. 33 : 489495. Thorseth, I.H., Fumes, H. & 0. Tumyr 1995. Textural and chemical effects of bacterial activity on basaltic glass: an experimental approach. Chem. Geol. I 19: 139-160. Ullman, W.J., Kirchman, D.L., Welch, S.A. & P. Vandevivere 1996. Laboratory evidence for microbially mediated silicate mineral dissolution in nature. Chem. Geol. 132: 11-17. Walsh, F. & R. Mitchell 1972. A pH dependent succession of iron bacteria. Environ. Sci. Technol. 6: 809-8 12.
Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
A test of aqueous speciation: Measured vs. calculated free fluoride ion activity D .K .Nordstrorn U S . Geological Survey, Boulder, Colorado USA
ABSTRACT: More than 35 water samples from Yellowstone National Park have been analyzed for free fluoride ion activity by ion-selective electrode (ISE) potentiometry. Sample pH ranged from 2.3 to 9.0 and free fluoride concentrations ranged from <1% to >99% of total dissolved fluoride. Agreement between computed activities based on the WATEQ4F speciation code and the ISE measured activities was within 1-30% for all samples at or above 10-6molal, providing excellent corroboration for chemical speciation models.
1 INTRODUCTION Aqueous speciation models are used extensively to interpret water-rock interactions, reactive-transport processes, and solute-biota interactions (Nordstrom & Munoz 1994; Langmuir 1997). Evaluation of their reliability, however, is rarely done (Schecher & Driscoll 1987; Nordstrom & Ball 1989; Nordstrom 1996) and often only for relatively simple systems (Harvie & Weare 1980; Ptacek & Blowes 2000). Different techniques for measuring inorganic or organic speciation often give varying results and analytical speciation usually is not compared with computed speciation (Nordstrom 1996). Such comparisons are needed to evaluate the accuracy of aqueous speciation. Ion-selective electrodes (ISEs) provide direct measurements of free ion activities that can be compared with free ion activities computed from water analyses and a speciation model. This paper reports on a comparison between free fluoride ion activities measured with the fluoride ISE and those computed with the WATEQ4F aqueous speciation model (Ball & Nordstrom 1991) for geothermal waters from Yellowstone National Park.
2 YELLOWSTONE WATER CHEMISTRY Figure 1. F- concentrations vs. pH
Geothermal water samples from Yellowstone National Park have a wide range of pH (1-lO), temperature (ambient to boiling), and fluoride concentrations (0.1-48 mg/L). The wide variation in water compositions offers the possibility of a wide range of fluoride speciation that can be determined. Figure 1 is a plot of total dissolved fluoride concentration against pH for most of the major thermal areas in Yellowstone (data from Thompson & DeMonge 317
1996). Note the occurrence of high fluoride concentrations is restricted to waters with field pH values > 6. Waters with pH < 6 have fluoride concentrations < 10 mg/L with the exception of 2 samples from Crater Hills (that are also unusually high in boron, arsenic, and silica). The proportion of free F- to total F- at high pH should be greater because of competition from hydrolysis and the
higher concentration of OH- ions. At low pH, some fluoride might be partially lost by volatilization of HF(,, and it might be diluted by cool, shallow ground waters of lower TDS.
Jenne(1977), Nordstrom et al. (1990), and Nordstrom & May (1996).
4 RESULTS AND DISCUSSION
3 ANALYTICAL METHODS 3.1 Major ions Major cations and most trace elements were determined by inductively-coupled plasma spectrometry, major anions by ion chromatography except for alkalinity which was determined by titration, and fluoride which was determined by ISE. Most determinations had errors of about 5% and most detection limits were in the range of 1-100 ppb. Charge balances were commonly less than 5% and nearly all were less than 10%. The analyses are reported in Ball et al. (1998a,b).
3.2 Fluoride ISE method The fluoride ISE is highly sensitive and precise for the measurement of fluoride ions and is calibrated for total dissolved fluoride by mixing 1:l of sample or standard with “total ionic strength adjustment buffer” or TISAB to eliminate aqueous complexes of fluoride and ionic strength effects. The measurement of free F- activity must be obtained without any added reagents. The electrode is calibrated against a series of NaF standard solutions covering the linear working range of the electrode (about 100 ppb to at least 50 ppm). The activities of the standard solutions are calculated with the same model used for the speciation calculations, WATEQ4F (Ball & Nordstrom 1991). The electrode is then used to measure the electromotive force of the sample solutions. Kinetic problems and non-linearity of the electrode response begin to be noticeable below 100 ppb fluoride (or the equivalent activity, about 10-6molal, for a mixed electrolyte solution) and the uncertainty increases substantially. For the speciation calculation, the pH and temperature of the sample must be known under the same conditions as the fluoride ion activity measurement. Hence, the samples must be air-equilibrated before the activity measurements can be done (CO;! degassing). Sample volumes of 25 ml are used with a miniature teflon stirring bar on an electric stirrer. Insulation between the beaker and the stirring platform minimizes heating effects.
Comparison of measured vs. calculated free fluoride ion activity is shown in Figure 2. The values cover a range of about 4 orders of magnitude. The comparison is considered excellent, especially when the wide range in pH and solute composition is considered. The two methods agree to within 1-30% of each other. This comparison is the only one known for a speciation test on natural waters with the fluoride ISE. A deviation plot is a more sensitive measure of the variance. The deviation in percent difference relative to the measured value as a function of free F- activity is shown in Figure 3. This deviation shows a clear trend wherein the calculated value becomes systematically greater than the measured value as the activity decreases below 10-6 molal. This systematic bias could be caused by at least two possible phenomena. First is the inability of the ISE to reach an equilibrium potential rapidly at low fluoride activities. The other possibility is an ionic strength effect because no constant ionic medium (such as NaCl) was added to samples or standard.
3.3 Speciation computations
Complete water analyses with ambient lab temperature and lab pH values were input to the WATEQ4F speciation code. The thermodynamic data relevant to fluoride speciation can be found in Nordstrom & 318
Figure 2. Calculated vs. measured aF-
ity to the computed aqueous speciation for fluoride. aF. This bias may be caused by slow rates of equilibration for the fluoride ISE at low activities or the problems with changing ionic strength effects between the sample and the standards. 3. Free fluoride ion concentrations can range from < 1% of the total dissolved fluoride at low pH to > 99% at high pH. This wide range in concentration is explained by strong complexing with H' and A13+under acidic conditions. 4. Comparative evaluations of this type can demonstrate the reliability of aqueous speciation computations and increase our confidence in geochemical models.
ACKNOWLEDGMENTS
-100 -50 0 50 100 150 Figure 3. % difference [(meas.-calc.) x 100/meas.]
-150
The proportion of free to total dissolved fluoride varies as a function of pH. The range is from <1% free fluoride at low pH values to >99% free fluoride at high pH. The main complexes are HF', HFz-, AlF", A1F2+, and AlF3". Some examples of the dominant complexes are shown in Table 1 as a function of pH. Table 1. Fluoride speciation relative to pH. Ojo Caliente Cinder Pool field pH 7.72 4.22 3.95 8.61 lab pH field T 93°C 95°C 22°C lab T 22°C 6.24 F T (lng/L) 31.6 57.2 % F-fm 99.3 % HF" 7.9 0.08 %A~F~+ % A1Fg 8.7 23.3 % AlF3" 2.3 % AlF4-
I am very grateful for the assistance of Davison Vivit who helped in the initial phases of this study and to all those who helped in collecting and analyzing samples, but most especially Jim Ball, Blaine McCleskey, Greg Druschel, Jenny DeMonge, and Jo Burchard. I am thankful to Everett Jenne who encouraged me to pursue this line of research. Jim Ball and Howard Taylor graciously reviewed this manuscript and provided helpful comments. REFERENCES
Black Pool 2.61 2.54 89°C 22°C 5.28 0.8 2.7 43 .O 51.8 1.57
5 CONCLUSIONS 1. For 38 water samples of varying composition and pH values from 2.3 to 9, the agreement between analytical and computed aqueous speciation for the free fluoride ion is 1-30%, indicating a high reliabil319
Ball, J.W. & Nordstrom, D.K. 1991. User's manual for WATEQ4F, with revised thermodynamic data base and test cases for calculating speciation of major, trace, and redox elements in natural waters. U.S. Geol. Survey Open-File Report 9 1- 183. Ball, J.W., Nordstrom, D.K., Jenne, E.A., & Davison, V.V. 199%. Chemical analyses of hot springs, pools, geysers, and surface waters from Yellowstone National Park, Wyoming, and vicinity, 1974-1975. U.S. Geol. Survey Open-File Report 98-182. Ball, J.W., Nordstrom, D.K., Cunningham, K.M., Schoonen, M.A.A., Xu, Y., & DeMonge, J.M. 1998b. Water-chemistry and on-site sulfur-speciation data for selected springs in Yellowstone National Park, Wyoming, 1994-1995. U.S. Geol. Survey Open-File Report 98-574. Harvie, C.E. & Weare, J.H. 1980. The prediction of mineral solubilities in natural waters: the Na-K-MgCa-C1-S04-H20 system from zero to high concentration at 25°C. Geochim. Cosmochim. Acta 44:981-997. Langmuir, D. 1997. Aqueous environmental geochemistry. Englewood Cliffs, NJ:Prentice-Hall. Nordstrom, D.K. 1996. Trace metal speciation in natural waters: Computational vs. analytical. Water, Air, Soil Pollution 90:257-267. Nordstrom, D.K. & Ball, E.A. 1989. Mineral saturation
states in natural waters and their sensitivity to thermodynamic and analytic errors. Sci. Geol., Bull. 42:269-280. Nordstrom, D.K. & Jenne, E.A. 1977. Fluorite solubility equilibria in selected geothermal waters. Geochim. Cosmochim. Acta 41:175-188. Nordstrom, D.K. & May, H.M. 1996. Aqueous equilibrium data for mononuclear aluminum species. In G. Sposito (ed), The environmental chemistry of aluminum:39-80, 2nd edition. Boca Raton: CRC Press/Lewis Publishers. Nordstrom, D.K. & Munoz, J.L. 1994. Geochemical thermodynamics. Boston: Blackwell Scientific. Nordstrom, D.K., Plummer, L.N., Langmuir, D., Busenberg, E., May, H.M., Jones, B.F., & Parkhurst, D.L. 1990. Revised chemical equilibrium data for major water-mineral reactions and their limitations. In D.C. Melchior & R.L. Bassett, (eds), Chemical modeling in aqueous systems II:398-413. Washington, D.C.:Am. Chem. Soc. Symp. 416. Ptacek, C.E. & Blowes, D.W. 2000. Predicting sulfate mineral solubilities in concentrated waters. Rev. Mineral. 40:(in press). Schecher, W.D. & Driscoll, C.T. 1987. An evaluation of uncertainty associated with aluminum equilibrium calculations. Water Resour. Res. 23525-534. Thompson, J.M. & DeMonge, J.M. 1996. Chemical analyses of hot springs, pools, and geysers from Yellowstone National Park, Wyoming and Vicinity, 1980-1993. U.S. Geol. Survey Open-File Report 9668.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Fading of luminescense in feldspars - an autoradiographic method E.Oila, S .Pinnioja, M.Siitari-Kauppi & V.Aaltonen Department of Radiochemistry, University of Helsinki, Finland
A .Lindberg Geological Survey of Finland, Espoo, Finland
ABSTRACT: Irradiation-induced luminescence is known to be characteristic of feldspar and quartz minerals. In the present study, sensitivity of an autoradiographic method was tested to see whether or not, it would be sufficient to study differences in fading rates of luminescence. Results show that for these minerals differences were easily detected. An autoradiography with a digital image processing technique may offer a great help for investigate mineralogy of centimetric scale rock surface. However, the different parameters of the different samples can cause a lot of variations in the intensity of luminescence and the results can not be generalized. Every sample should be considered separately. 1 INTRODUCTION
been emitted from the mineral of a sample. Today, eqiupment for digitizeing the images are very sensitive and accurate. A digitized image is comprised of pixels and the size of the pixels depends on the capability of the equipment used. Digital images can be easily superimposed on one another and the decay rates can be calculated for each pixel. As a result, it is possible to get an image of decay rates of a whole surfaces, as displayed in Figure 1.
Ionizing radiation induces a luminescence in minerals which is characteristic in both intensity and spectra (McKeever 1985). Autoradiography has shown to be a useful technique for measuring the luminescense emissions of rock matrices (SiitariKauppi et al. 1998). An autoradiographic method for analysing the intensity of luminescence has been compared with the results given for crushed minerals by a TL reader (Pinnioja et al. 1999a, b). In this paper we studied if autoradiography combined with a digital image processing technique can also be an effective method for investigating the fading of luminescence from centimetric scale luminesent rock surface. The fading of luminescense at a constant temperature is a simple first-order decay, and it follows the equation 1. (McKeever 1985).
I(t) = I, exp(-t/z ) where t is time, I(t) is the intensity at the time t, I. is the intensity at time = 0 and zis mineral specific decay rate constant of luminescence. The different fading rates of luminescence from feldspar and quartz minerals are easily measured by exposing autoradiographic (Ag-film) with different delay times from end of the irradiation. The resolution of Ag-films is usually a few micrometres. A darkening by one dot on a film represents the exact amount of luminescence which has
Figure 1. Model of the measuring intensities from digitized autoradiographs which are exposed from same sample with differnet delay times. The intensity values of every pixel can be measured as a finction of delay time and place of sample (x,y) then fit the data in equation 1.
321
2 EXPERIMENTAL, 2.1 Samples Four thin slices were made of Rapakivi-granite (Y161 and Y-174). The Y-161 samples were made of fresh granite and Y-174 of altered granite. The samples were cored from Hastholmen, Loviisa, Finland. Thin slices were made of 2.5 x 4.5 cm2 in size and variations in thickness are listed in Table 1.
where OD is the optical density, I, the intensity of the background and I the intensiy of the sample. The measured areas were 0.3 mm . The areas were selected to ensure the uniformity of the mineralsamples. Quantification of the autoradiographs by digital image analysis has been described in detail earlier by Hellmuth et al. (1993, 1994). 3 RESULTS
Table 1. Data of the granite samples. Sample code
Drilling depth
Thickness of slice (mm)
(4
I I
I
Y-174-1 Y-174-2 Y-161-1 Y-161-2
I I
I
174 174 161 161
I I
I
1 0.1 1 0.1
I I
2.2 Irradiation Irradiation was carried out using a 500 Ci G°Coirradiation source twice with the same irradiating time (24 minutes) and during that time the dose was 470 Gy in one irradiation. Before each irradiation the samples were heated for two hours at 120 "C to eliminate previous luminescence of the samples. Irradiations were performed in normal atmosphare at ambient temperature.
2.3 Autoradiography After the first irradiation and a 15 minutes delay, the samples were then exposed on an X-ray film (Kodak X - OMAT MA) in the darkroom. The films were changed every 15 minutes and this was done 7 times. Thus, the first film was exposed 15-30 minutes from the end of the irradiation and the last one was exposed 105-120 minutes from the end of the irradiation. After the second irradiation the delay time was 120 minutes and 15 minutes exposures were made 5 times. Quantitative interpolation of the results was based on digital image analysis, which began with the digitizing of the autoradiograph into pixels (300 dpi). All the intensities of the subdomains were then converted into corresponding optical densities, which were dependent on the luminescence intensity of the different minerals. Because the image source (table scanner Ricoh FS 2) and the amplifier of the digital image analyser are linear, the digitized gray levels of the film could be expressed as intensities. Intensity here means the light intensity coming through the autoradiographic film. Optical densities were derived from intensities as follows:
Photographs of the samples Y-174-1 and Y-174-2 and the corresponding autoradiographs of the first exposure (delay time 15 minutes) are shown in Figure 2. Digital image analyses were done on feldspar grains (points A and B) and from quartz grains (points C). Figures 3-6 show fading of the optical density as a function of time at the selected points. Background value for intensities were measured individually for every film. Optical density decrease is in accordance to equation 1. At the very intensitve pixels some error can be caused by a saturation of the film; at some point the film did not darken lineary when luminescence increased. One irradiation to produce one fading curve is optimal. However, in this study two irradiations were carried out which caused some uncertainty about the curves. This can be seen as a small increase in OD values after second irradiation. Accurate mineral composition of the measured points was not carried out but the same kind of mineral areas were chosen for the analyses. This is a reason why the fading curves could not be compared with each other. However, microscopy was used to ensure that the selected grains from the same depth were uniform and similar. Furthermore, this ensured that luininescenc as a function ofihin slice thickness was comparable. Comparing the samples Y-174- 1 with Y- 174-2 and Y-161-1 with Y-161-2, it was apparent that the thicker samples give off more intense luminescence than thiner samples. The optical densities of fresh granite were higher than the optical densities of the altered granite.
4 CONCLUSIONS Autoradiography and optical densitometry using an application of digital a image processing technique was found to be a suitable method for determining the mineral specific differences of luminescence fading. However, the exact conditions of irradiation are difficult to repeat and all measurements should be done after only one irradiation episode. In addition, the different parameters of the different sam-
322
Figure 4. Measured optical density at the points as a function of delay time, determined by quantitative autoradiographic method. Sample Y- 174-2.
Figure 2. Rapakivi granite thin slices Y-174-1 and Y-174-2 (top) and corresponding autoradiographs of the surfaces after irradiation at a dose of 470 Gy and 15 minutes delay time. Autoradiographic analyses are from points A,B and C.
Figure 5. Measured optical density at the points as a function of delay time, determined by quantitative autoradiographic method. Sample Y- 16 1- I .
Figure 3. Measured optical density at the points as a function of delay time, determined by quantitative autoradiographic method. Sample Y- 174-1.
ples can cause a lot of variations in the intensity of luminescence and the results can not be generalized. Every sample should be treated separately. The benefit of the method is that whole rock samples can be analysed at once, allowing the fedious work of crushing and separating grains to be avoided.
323
Figure 6. Measured optical density at the points as a fimction of delay time, determined by quantitative autoradiographic method. Sample Y-161-2.
REFERENCES Hellmuth, K.H., Siitari-Kauppi, M. & A. Lindberg 1993. Study of Porosity and Mi ration Pathways in Crystaline Rock by Impregnation with ''C-polymethylmethacrylate. Journal of ContaminantHydrology 13: 403-418. Hellmuth, K.H., Lukkarinen, S. & M. Siitari-Kauppi 1994. Rock Matrix Studies with Carbon-14- Polymethylmethacrylate (PMMA); Method Development and Applications. IsotopenpraxisEiiviron. Health Stud. 30: 47-60. McKeever, S.W.S. 1985. Thermoluminescence of Solids. Cambridge University Press. Siitari-Kauppi, M., Pinnioja, S & A. Lindberg, 1998. An autoradiographic method for studying irradiation-induced luminescence in feldspar. Proceedings of the 9" international symposium on water-rock interaction. Arehart, G.B & J.R. Hulston (eds) Balkema, Rotterdam, (1998) pp. 859-862. Pinnioja, S., Siitari-Kauppi, M. & A. Lindberg 1999. Effect of feldspar composition on thermoluminescence in minerals separated fiom food. Radiat. Phys. Chem. 54505-5 16. Pinnioja, S., Siitari-Kauppi, M., Jernstrom, J. & A. Lindberg 1999. Detection of irradiation foods by luminescence of minerals - effect of mineral composition on luminescence intensity. Radiat. Phys. Chem. 5 5 : 743-747.
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Sampling techniques and pH measurement methods for geochemical analysis of deep groundwaters H.Pitsch* , C.Beaucaire* , P.Meier & S .Grappin Commissariat Ci 1’knergie atomique, DCCIDESDISESD, 91191 Cif sur Yvette, France *Present address: IPSNIDPREISERGD, CEA-FAR, BP N06, 92265 Fontenay aux Roses, France
ABSTRACT : Two deep aquifers enclosed between tertiary clay layers in the underground of the northern Belgian province Kempen were characterised for geochemical and hydrogeological purposes. Major and trace elements were analysed in waters sampled from 15 boreholes by four different methods: directly from the head of the artesian wells, with an immersed pump, by a specially designed constant pressure sampler or using a device that brings a water column from the desired depth to the surface by applying a nitrogen pressure. pH was measured either with a flow-through cell or in situ with a home-made spectrophotometric sensor fixed at the end of a 350 m long optical fibre. Methods are described and the analytical results are examined putting in evidence the perturbing effects of the different techniques. Conclusions are drawn on the “best” methods to get precise and accurate data for geochemical interpretation. 1 INTRODUCTION The measurements described here were carried out within the frame of the European Community’s project PHYMOL, dedicated to investigate the past and present groundwater flow in the deep aquifers of a clayey formation (Marivoet et al. 2000, Pitsch & Beaucaire 2000). Some results of previous programmes are also considered. The site of investigation consists of an area of ca 10,000 km2 around the Belgian nuclear research centre at Mol (SCKeCEN) where the underground research facility HADES (acronym for high activity disposal experimental site) is located at - 223 m in the Boom clay massif. This plastic clay belongs to a system of successive layers of tertiary clays and clayey sands or silts. The Boom clay and two underlying aquifers separated by the Asse clay were investigated. The objective was to identify relevant tracers and tools to describe the hydrogeological evolution of such an aquifer - aquitard system over a large period of time including the last glacial period, some 20,000 years ago. Even if the most important parameters are found in the isotopic composition of specific elements, the chemical composition of the waters is also necessary to identify the water-rock equilibrium state at each point and understand mixing processes in extended aquifers. The interpretation of the water chemistry helps to assess the significance of isotopic data and gives additional parameters.
The sampling and analytical strategy was therefore focused on getting as precise and accurate data as possible at each point in order to distinguish waters representative of global trends in the aquifer compositions from those that indicate a local anomaly. Consequently a comparison of available sampling and measurement methods was made in order to put in evidence possible artefacts. This allowed to collect reliable data during the present programme and to critically evaluate the water compositions obtained during earlier sampling campaigns, building up a data bank of interpretable values. The results of this analytical work are presented here. 2 WATER SAMPLING TECHNIQUES The interstitial fluid of the Boom clay was obtained from a nearly horizontal piezometer, made of stainless steel, through a filter located 14 m away from the wall of the HADES gallery. Samples were collected under vacuum in stainless steel cells connected to the tubing attached to the piezometer filter (see Pitsch et al. 1995). Waters of the deep aquifers were sampled from drillholes belonging to the SCKeCEN or to the Belgian Geological Survey, using four different techniques. The home made hydrocapturing device (hc) brings a water column from the desired depth to the surface by applying nitrogen pressure, after a series
325
of cleaning cycles. This system has been described previously (Griffault et al. 1996). Waters were also extracted from the drillholes at a depth of 50 m with a Griinfoss inmersed pump (iP>A stainless steel constant pressure sampler (cp) manufactured by MetroMesure was used, particularly when water was also sampled for noble gas analysis. When immersed in the drillhole at the correct depth, a check valve opens and a piston is displaced by the pressure of the sampled fluid, letting it fill the bottle. To ensure slow sampling, a slightly lower counter-pressure is applied before inmersion to a reservoir filled with deionised water located behind the piston, which empties during sampling through a small check-valve. Correct sampling pressure was determined using a pressure gauge immersed at the required depth or calculated from the water column height. Artesian wells were sampled directly at the surface after purging with twice the volume of the drillhole. For chemical analysis, sequential filtration under argon was applied in the field, in order to avoid sampling artefacts and for cation analysis the sample was acidified with ultrapure acid. Samples were kept at 4°C until they were analysed. Detailed description of the filtration and conditioning procedures can be found in Griffault et al. (1996) and Beaucaire et al. (1999). The alkalinity was determined at the site by potentiometric titration: WTW pH-meter, Ingold combined pH electrode + Tacussel Ag/AgCl reference. Major anions and cations were analysed in the laboratory by ion chromatography: Dionex 4500i device with conductimetric detection + CSRS-I or CSRS-I1 (cations) and ASRS-I (anions) electrochemical suppression modules; Columns: Dionex CS12 (cations) and AS4A or AS14 (anions); Eluents: methanesulfonic acid (cations) and carbonate + bicarbonate (anions). For noble gas analysis water was sampled into a copper tube squeezed every 15 cm in order to obtain five equivalent samples at each spot. The analysis was carried out by the Noble gas hydro-geochemistry laboratory at Reading, UK.
electrochemical cell of 125 ml was used when water was sampled from an artesian well (aw), with the immersed pump (ip) or with the hydrocapturing device (hc). In situ pH measurements were carried out with the previously described Optolec H (Motellier et al. 1995) which monitors the light absorption by a resin bead impregnated with a pH indicator and fixed at the end of an optical fibre immersed in the sample. This system was adapted to measurements in drillholes up to 350 m deep by enhancing the amplification of the signal, giving the sensor a thin cylindrical geometry and improving the reflection power of the mirror screw. The indicator was bromophenol red and calibration was carried out on site before and after each measurement, at the temperature measured in the drillhole.
mM Mol 1 5 ~ hc 30.1 Mol 15c ip 30.3
mM mM mM mM meq/l 0.429 0.234 0.296 0.018 31.6 0.540 0.315 0.411 0.014 32.3
-2.1 1.7
3 MEASUREMENT DEVICES
Zoe43b Zoe43b
0.432 0.387 0.263 0.048 0.419 0.394 0.263 0.062
- 0.4
The in situ temperature was measured with a 500 m long probe. A WTW pH-meter and an Ingold combined electrode with a Pt temperature sensor were used for pH measurements in the field. A home made flow cell of small volume (5 ml) containing the electrode sensor was directly connected to the constant pressure sampler (cp). A Metrohm autothermostated
326
4 RESULTS AND DISCUSSION 4. I Chemical composition
In table 1 two examples show that changing the sampling method does not induce a significant difference in the resulting concentrations of major elements. This is also true for trace elements and, according to results obtained by Universite Paris Sud, for the isotopic composition. Table 1. Chemical composition of waters sampled by different methods: hydrocapturing device (hc); immersed pump (ip); artesian well (aw); constant pressure sampler (cp). Drillhole Mol 15c Mol 15c
Cl mM 18.4 18.4
Br mM 0.020 0.021
Zoe43b Zoe43b
aw 14.3 cp 14.4
17.9 17.9
0.023 0.108 0.011 0.148 0.022 0.108 0.012 0.168
K
Ca
Drillhole
Na
F mM 0.080 0.079
sio?
Alk. mM hc 13.2 ip 12.7
Mg
SO4
mM mM 0.311 0.264 0.262 0.203
Li
Anions meall 32.3 31.8 32.4 32.5
Cation Balance s
aw 30.5 cp 30.5
32.2 32.3
%
-0.4
4.2 Alkalinity andpH Alkalinity must be measured in the field immediately after sampling whenever acidic cations may be generated during transport or conservation of the sample for anion analysis. This is the case for waters which contain large amounts of iron: Fe(I1) is
slowly oxidised by oxygen to Fe(II1) that precipitates as hydroxide, two protons are liberated for each iron ion and the alkalinity is lowered by an equivalent amount. Table 2 shows that, with a flow cell connected directly to the constant pressure sampler, the same pH is measured independently of the sampling depth in a purged well and furthermore no difference is observed when the water is sampled directly from the artesian well. Here also the alkalinity values obtained in the field show that the concentration of dissolved species is not affected by the change in sampling method.
PH
Alkalinity
m
Sampling pressure bars
aw
1
8.12
mM 14.3
60 100 150 200
5 9.3 14.6 20
8.08 8.09 8.12 8.12
14.6 14.3 14.5 14.4
Sampling depth
a constant pressure sampler. The decrease of degassing with increasing molecular weight may be related to the increase of the solvent destructuring effect of the noble gas and the consequent decrease in the salting-out effect of the groundwater on the dissolved gas (Samoilov 1965, Pitsch 1986, Castro 1995). Table 3: Noble gas concentration in Zoe 43b drillhole water, sampled directly from the artesian well (Caw) and at 200 m depth with the constant pressure sampler (Ccp). Unit: cm3 STP / g HzO.
The fact that the acidity level of the fluid is not perturbed along the water column rising in an artesian well is confirmed by the results presented in Figure 1. The same pH value is measured in situ with an optode and with a flow cell directly connected to the artesian well. This is probably due to the slow degassing velocity of carbon dioxide in a smoothly moving water and corresponds to a poorly labile bond between CO2 and water. Conversely the pH measured at the surface in water extracted by the hydrocapturing device (hc) is systematically higher than the one measured in situ with an optode (Fig. 1) by nearly half a unit. This was also observed when comparing the pH of the Boom clay interstitial water measured with the same optode as here and the pH determined by a classical technique on an extracted sample, and is due to degassing of carbon dioxide from the sample exposed to the atmosphere (Pitsch et al. 1995). Here degassing is accelerated by the mechanical perturbation due to the sampling device.
Caw ccp Caw/Ccp
Ne(IO-') 2.51 5.21 0.48
Ar(l0"') 4.29 5.35 0.80
Kr(lO-') 10.61 11.47 0.93
Xe(10-') 1.61 1.62 0.99
5 CONCLUSION This comparison study showed that, provided adapted filtering and conditioning of samples is respected, the general chemical composition of the water is not affected by the sampling method except for volatile species or those involved in equilibriums with volatile species, for which methods which induce a degassing effect must be avoided. Only indepth sampling with a constant pressure sampler provides integer samples for noble gas analysis. To get an accurate value that fairly agrees with the one calculated with the geochemical model (Pitsch & Beaucaire 2000) the in situ determination of pH proved most suitable. As an alternative to in situ measurements, precise and accurate pH values are also obtained using a flow cell connected either to an artesian well or to a constant pressure sampler that brings the
4.3 Noble gases
Noble gases are more prone to degassing than carbon dioxide, probably due to weaker interaction with the water molecules. Table 3 shows that concentrations in a sample taken directly from an artesian well are lower than those in the water sampled in-depth with 327
water out of the drillhole from the required depth. This may represent a non negligible economy of time and investment. ACKNOWLEDGEMENTS Authors wish to thank the European Commission for funding and Serge Labat from S C K C E N for his helpful contribution to the field work, as well as Marc Dao from MetroMesures for the design and manufacturing of the constant pressure sampler.
REFERENCES Castro, M. C. 1995. Transfert des gaz rares dans les eaux des bassins sedimentaires : exemple du bassin de Paris. Thbe, Universite Paris 6. Griffault, L., Merceron, T., Mossmann, J. R., Neerdael, B., De Canniere, P., Beaucaire, P., Daumas, S., Bianchi, A. & R. Christen 1996. Projet ((ARCHIMEDE - argile)) : Acquisition et regulation de la chimie des eaux en milieu argileux pour le projet de stockage de dechets radioactifs en formation geologique. European Commission report EUR 17454 FR, 7 1 pp. Marivoet J., Van Keer, I., Wemaere, I., Hardy, L., Pitsch, H., Beaucaire, C., Michelot, J.L., Marlin, C., Philippot, A.C., Hassanizadeh, M. & F. Van Weert 2000. A palaeohydrogeological study of the Mol site (PHYMOL project). European Commission report EUR 19146 EN, 101 pp. Motellier, S., Noire, M.H., Pitsch, H. & Dureault B. 1995. pH determination of clay interstitial water using a fiber-optic sensor. Sensors and actuators “B“29, No 1-3, 345-352. Pitsch H. 1986. Application de la mise en equation des effets de relargage a la modelisation des systkmes de distribution liquide-liquide. Thbe, Universite Paris 6. Pitsch H., Motellier, S., L’Henoret, P. & C. Boursat 1995. Characterisation of deep underground fluids. Part I: pH determination in a clayey formation. Proc. 8th International symposium on water-rock interaction (WRI-S), Vladivostok, Russia, pp 467-470. Rotterdam: Balkema. Pitsch H. & C. Beaucaire 2000. A palaeo-hydrogeological study of the Mol site (PHYMOL project). Topical report 2: Hydrogeochemistry. European Commission report DOC RTD/54/2000 - EN, 42 pp. Samoilov 0. Ya. 1965. Structure of aqueous electrolyte solutions and the hydration of ions. Translation: D. J. G. Yves, Consultant Bureau, New York, 1965.
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Revised thennodynamic properties of malachite and azurite W.Preis & H.Gamsjager Institut fur Physikalische Chemie, Montanuniversitat Leoben, A-8700 Leoben, Austria
ABSTRACT: The reliability of the thermodynamic properties of malachite, C U ~ ( O H ) ~ C and O ~ , azurite, C U ~ ( O H ) ~ ( C Otabulated ~ ) ~ , in important thermodynamic data compilations, have been checked by constructing predominance diagrams for the ternary system Cu2+- H20 - C02. Severe discrepancies between theoretical predictions and experimental observations of phase equilibria have been revealed. The thermodynamic properties of the copper carbonates have been reinvestigated by performing solubility measurements as a function of temperature in perchlorate media of constant ionic strength. The results of the solubility experiments confirm the literature values for the standard entropy of malachite and azurite, whereas the literature data for the standard enthalpy of formation should be rejected and revised values are recommended. 1 INTRODUCTION The thermodynamic properties of the copper carbonates malachite [Cu2(OH)2C03] and azurite [ C U ~ ( O H ) ~ ( C Oare ~ ) ~tabulated ] in important thermodynamic databases (Wagman et al. 1982, Robie & Hemingway 1995). The enthalpies of dissolution of malachite and azurite were originally determined by Roth et al. (1941) employing solution calorimetry. Their results were re-evaluated by Wagman et al. (1982), leading to the standard enthalpies of formation Affl[Cu2(0H)2C03] = - 1051.4 kJ mol-' and = - 1632.2 kJ mol-'. These A&!@[CU~(OH)~(CO~)~] data have finally entered the NBS - tables (Wagman et al. 1982). Furthermore, Richardson & Brown (1974) determined the standard enthalpy of formation of malachite by means of solution calorimetry, A~HO[CU~(OH)~CO~] = - 1054.0 kJ mol-'. Kiseleva et al. (1992) carried out decomposition experiments on malachite and azurite in a drop calorimeter and obtained A~HO[CU~(OH)~CO~] = - 1046.5 kJ mo1-*, = - 1629.5 kJ mol-'. The A~HQ[CU~(OH)~(CO~)~] standard entropies of malachite and azurite selected by Robie & Hemingway (1995) are based on lowtemperature measurements of the heat capacity (Kiseleva et al. 1992). The value for the standard enthalpy of formation of malachite given by Richardson & Brown (1974) was ado ted by Robie and Hemingway (1995), whereas A&$ of azurite is based on solution calorimetry (Roth et al. 1941) as well as drop calorimetry (Kiseleva et al. 1992). Schindler et al. (1968) performed accurate and precise solubility measurements on malachite and
azurite at 298.15 K, resulting in fairly reiiabie values for the standard Gibbs free energy of formation A@[Cu2(0H)2C03] = - 901.3 kJ mol-' and A@[CU~(OH)~(CO~>~] = - 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
337
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
340
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.
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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.
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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: (<63, 63-75, 75-90, 90-106, 106-125, 125-150, 150-180, 180-212, 212-250, 250300, 300-355, 425-500, 500-600 pm). All subsequent samples were then re-constituted from the above sub-samples. To determine the dissolution kinetics of the schist material a series of 3g samples of different particle size ranges, (63-150, 150-300, and 300-600pm), were placed in polystyrene containers with 30ml of de-ionized water. A series of 150-300pm samples were also leached with a simulated monsoonal rain-
353
water. The composition of the rainwater was modelled on that described in Gillet & Ayers (1990), (see Table 2). The closed containers were then placed in a rotating rack that was submerged in a water bath. Each vial was rotated end over end approximately once every 10 seconds. Throughout the course of the experiment the temperature bath was kept at 2 l+/0.5"C. Table 1. Chemical composition of schist starting material. Oxide Trace element
(%I Si02 63.04 Ti07 0.43 A1203 14.19 Fe20: 5.57 MnO 0.04 MgO 9.97 CaO 0.10 Na20 <0.03 K20 1.35 PI05 0.08 so3 0.01 LOF 5.61 Total 100.32
(PPd Ba 57 Ce 37 CO 23 Cr 104 Cu 297 Ga 25 85 Ni Rb 63 U 45 v 73 Y 13 Zn 118 Zr 153
Dominant phase: Quartz Semi-dominant phase: ferroan clinochlore Minor phases: muscovite, illite, phlogopite
Table 2. Composition of simulated monsoonal rainwater. Species Concentration Concentration This study (Gillet&Ayers 1990) [H'! "a 1 [K'I2+
[q1
[Ca 1' [CU
m03-1
[so:-] [COOH-]
(POW 14.2 3.9 0.7 0.6 1.9 6.5 2.8 2.6 8.0
(IrmoW 13 3.8 0.7 0.6 1.7 7.5 3.2 2.6 6.3
Vials were removed periodically for sampling and the pH determined. The pH probe was calibrated between each measurement with the appropriate pH=4,7 or pH=lO commercial buffer solutions, (Orion). Following the pH determination the sample was centrihged then the supematant was drawn off with a plastic syringe. Samples were then filtered through a 0.45pm nylon filter to remove any fine material remaining in the solution and acidified to pH-1 by the immediate addition of 1 drop of high purity concentrated HCl per 1Omls of solution. Inductively coupled plasma - atomic emission spectrometry (ICP-AES) was then utilized to determine the concentrations of Si, Al, Mg, K and Fe. After 14 days of leaching samples fiom each experiment were removed and the solid material dried
before being tested for changes in particle size distribution. The rock samples were also tested for changes in pore size distribution and surface area. Pore size distributions were determined using a mercury intrusion porosimeter, (range=0.03-360pm diameter) and also a nitrogen adsorptioddesorption technique (0.0006 - 0.3pm diameter). The surface areas of the powder samples were also determined via the nitrogen adsorptioddesorption unit (see table 3). 3 RESULTS 3.1 Changes in the leachate solution 3.1.1 Changes in the solution pH. Figure 1 shows the change in solution pH as a hnction of time for the 63-15Opm schist samples leaching in de-ionized water. The trends are indicative of all the systems that were examined in this study. The initial stages of the leaching process are characterized by a rapid increase in pH. In the above case this was from an initial value of pH=5.52 to 6.33 in the first minute of reaction, indicating a rapid exchange of cations and H+ is taking place at the rock surface. This is followed by a more gradual increase over the following weeks until the p H of the solution approaches 9 for each system. Surprisingly there was little difference in the response of the system with the different particle size ranges examined. The response of the 150-300pm schst leached in the simulated rainwater was similar to the de-ionized water systems even though the initial pH was significantly lower (pH=4.85).
Figure 1. The pH of the bulk solution as a function of time for 63-1 50pm particles in de-ionized water
3.1.2 Kinetics of cation release. The dissolution of silicon and magnesium as a function of time is shown in Figure 2 for the system 150300pm schist + simulated rainwater. Once again it was observed that, particularly for Mg, rapid ion exchange reactions occurred resulting in a large increase in the number of cations into the bulk solution
354
during the first minute of contact. This is then followed by a period, (up to -14 days), in which the cation concentration follows a power law relationship. Our data suggests that after 14 days the dissolution rates begin to increase above that indicated by the power law. However this must be verified in further study.
grinding of the original rock. In order to “clean” the rock surfaces of such anomalous sources of cations both ultrasonic cleaning and etching with HF have been required. An examination of the surfaces of the rock particles naturally present at the Ranger Uranium waste rock dump will be undertaken in the future to see if such an approach is warranted for this study. 3.2 Changes in the rock structure.
Figure 2. Dissolved silicon and magnesium concentration versus time. 0-Si, 0-Mg
In addition to studying the changes in leachate composition we sought to examine the effect of the leaching process on the rock structure. The short nature of the experiment undertaken meant that the total material leached from the rock matrix represented approximately 0.01-0.03% of the original rock mass. Given this low level of extraction only subtle differences in the rock matrix were expected to manifest themselves over this time period. Nevertheless visual changes in the leaching solutions were apparent toward the end of the experimental period particularly in the 63-150pm system where the solution became noticeably more turbid and took on a decidedly pale pinMred coloration. 3.2.1 Changes in the particle size distribution. Samples of the rock material that had been leached for 14 days were dried overnight at 105”C, dry sieved and the size distribution compared to that obtained for the unweathered sample. No discernible difference in size distribution could be found in any of the samples. The exercise was then repeated using a Malvern Laser Particle sizer. The obtained results indicate that no statistically significant changes in particle size distribution occurred even in the 0.5 to 10pm scale.
Figure 3. Cation concentration after 10 minutes leaching for various size fractions in de-ionized water.
Data for the effect of the particle size on the rate of cation release was ambiguous with no statistically significant increase in the rate of cation release being observed with decrease in particle size. This is illustrated in Figure 3 for cation levels detected after 10 minutes of leaching in de-ionized water. In all cases it is apparent that a great deal of dissolution activity has taken place in the initial stages of the leaching process. The results shown in Figure 3 also demonstrate that the effect of particle size is considerably less than that expected from a simple calculation of available surface areas using the average particle diameters involved. Both the rapid initial dissolution and lack of particle size effect have, in previous studies such as Holdren and Berner (1979), been established as being due to the presence of extremely fine, highly reactive particles and sites generated during the
355
3.2.2 Changes in the particle surface area and pore size distribution Samples of the starting material for each size fraction were subjected to both mercury intrusion porosimetry and nitrogen adsorptioddesorption analysis to determine both their BET surface areas and pore size distribution. After 14 days of leaching the rock material was sampled, dried and the tests repeated. Table 3. BET surface areas of both starting (t=O) and leached ( P I 4 days) schist materials. Size range Surface Area (m2/kg) S.A. decrease (P-4 63-150pm 150-300pm 300-600pm
T=O days 9750 7680 8340
150-300um (rain) 7680
T=14 days 8850 6630 6690 7440
(%) 9.2 13.7 19.8
3.1
Table 3 lists the BET surface areas of rock samples both before and after 14 days of leaching. The BET surface areas determined are orders of magnitude higher than if a population of solid rock spheres is assumed. For example, 225 pm schist spheres have a specific surface area of approximately 10m2/kg compared to the measured value of 7680m2/kg for the 150-300pm size fraction. This indicates that approximately 99.8% of the particle’s surface area is derived from sources such as internal porosity that is physically connected to the particle surface. It is also likely that the large surface area is due, in part, to the presence of fine particles generated in the grinding process. Table 3 also indicates that despite the short leaching period, a significant change in the rock surface area has occurred with surface areas decreasing by 10 to 20% when the leaching is carried out with de-ionized water. The effect increases as particle size gets larger. Leaching with simulated rainwater still leads to a decrease in surface area but the effect is less marked.
deposition of extremely fine precipitates or by the recrystallization of minerals generated during the leaching process. One further explanation is that the leaching process has induced a shift towards larger pore diameters (as per Sweevers et al. 1998).
4 CONCLUSIONS Leaching of a quartz-chlorite schist material over a 14 day period yielded cation dissolution rates that obeyed a power law relationship with time. It was observed that the dissolution rates did not vary significantly with particle size. This and the high initial cation release rates observed are likely to be the result of extremely fine particles generated during the sample preparation. While little evidence of changes in the particle size distribution was found, there was a significant decrease in the rock surface areas as determined by N2 adsorptiorddesorption. This decrease in surface area is linked to a decrease in the number of -4nm diameter pores observed in the rock matrix. This decrease may be the result of precipitate deposition, mineral recrystallization, or linked to some pore widening process. Future work will include experiments conducted over longer periods of time and will focus on the changes to the rock structure that occur during the leaching process. This paper has presented the techniques to be used in these studies. REFERENCES
Figure 4. Pore s u e distribution for 63-150pm schist before and after leaching in de-ionized water for 14 days. -T=O, 0-T=14 days
The pore size distribution data for the same rock samples provides an explanation for the observed decrease in surface area. Figure 4 shows the pore size distribution for 63-150pm schist leached in deionized water both before and after 14 days of leaching. The pore size distribution is plotted in terms of incremental surface area. It is clear from the plot that a population of pores of approximately 4nm in diameter contributes about half of the total surface area of the rock particle. The pore distribution peak at 4nm was present in each rock sample regardless of the particle size. After 14 days of leaching the number of the 4nm pores as well as those <4nm in diameter declined significantly leading to a subsequent decrease in the surface area of the sample. Unfortunately the tests do not shed any light on the fate of those pores. It is possible that they have been plugged either by the 356
Busenberg, E & C.V. Clemency 1976. The dissolution kinetics of feldspars at 25°C and 1 atm CO2 partial pressure. Geochimica et Cosmochimca Acta: 40: 41-49. Gillet, R.W. & G.P. Ayers 1990. Rainwater acidity at Jabiru, Australia, in the wet season of 1983184. The Science of the Total Environment: 92: 129-144. Hancock, G.R., K.G. Evans, G. Willgoose, D.R. Moliere, H.J. Saynor & R.J. Loch 2000. Medium term erosion simulation of an abandoned mine site using the SIBERIA landscape evolution model. Australian Journal of Soil Science: 38: 249-263. Holdren, G.R. & R.A. Berner 1979. Mechanism of Feldspar weathering - I. Experimental studies. Geochimica et Cosmochimica Acta:43:1161-1171. Pickering, R.J. 1962. I. Some leaching experiments on three quartz-free silicate rocks and their contribution to an understanding of laterization. Economic Geology: 57: 1185-1206. Sweevers, H., F. Delalieux, & R. van Grieken, 1998. Weathering of dolomitic sandstone under ambient conditions. Atmospheric Environment: 32(4): 733-748. Willgoose, G.R. & S.J. Rdey 1998. An assessment of the longterm erosional stability of a proposed mine rehabilitation. Earth Surface Processes and Landfom:23:237-259. Wollast, R. 1967. Kinetics of the alteration of K-feldspar in buffered solutions at low temperature. Geochimica et Cosmochimca Acta: 3 1: 635-648.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Rate of mineral dissolution during granite-hydrothermal alteration P.Zuddas & F.Seimbille Lab. G&ochimieEaux, Univ. Paris 7 & Inst. Phys. du Globe, 2 place Jussieu, case 7052, 75005 Paris, France
ABSTRACT: The significant gap between field and experimental estimation of mineral dissolution rates during granite fluid interaction can be attributed to the use of single minerals during experimental evaluation. We propose an approach to evaluate simultaneously the dissolution rate of K-feldspar, biotite and plagioclase during the hydrothermal alteration of granite rock by fluids having an initial composition close to the saturation with neogenic minerals. The methodology, based on the isotope doping technique, is suitable for granite systems where minerals have different strontium isotopic signature and Rb/K ratio. establish an isotopic mineral end-member for every mineral containing Sr and Rb because the (84Sr/S6Sr) and (s8Sr/86Sr)ratios are constants in all terrestrial rock and equal to 0.0565 and 0.1 194 respectively (Steiger & Jager 1977).
1 INTRODUCTION The chemical composition of fluids in the crust depends on a gradual transformation of instable primary minerals into more stable neogenic phases. In this irreversible transformation, primary minerals are source of dissolved elements while neogenic minerals control their activity in the fluid phase. The chemical composition of the fluids reflecting the global process may reach equilibrium with neogenic phases. In these systems, where the variation of the affinity reaction is very low it is difficult to evaluate rate and mechanism of mineral dissolution because fluid composition is quite stationary. We report a new approach to evaluate the rate of mineral dissolution during the hydrothermal alteration of a granite by fluids having an initial composition close to the saturation of neogenic phases. The methodology is based on the isotope doping technique. 2 ISOTOPIC DOPING TECHNIQUE GRANITE-FLUID INTERACTION
2.1 Experimental The experimental procedure consists of introducing powered granite of a known specific surface area and solution into a gold cell which is put into a steel autoclave. The initial composition of the fluid was similar to that found in granite geothermal area and corresponds to the saturation with respect to quartz, adularia, prehnite, low temperature albite, kaolinite, calcite at a given temperture (Zuddas & Michard 1993, Zuddas et al. 1995). For a given temperature and mineral association, the fluid equilibrium composition depends thus, on chloride content (mobile ion). Initial solutions were prepared with chemical reagents spiking with s4Sr (75%) and 39K (99%) added as strontium nitrate and potassium chloride respectively. The granite chosen in this study was composed mainly of oligoclase (An20, Kfeldspar, biotite and quartz with accessory sphene, metallic oxides, apatite and zircon. The chemical and isotopic composition of the rock and separated single minerals is reported in Table 1. The water rock ratio was 20. Experiments were carried out at different temperatures.
IN
This method applied to granite fluid alteration is based on the properties of both natural rock minerals and fluids artificially enriched in 84Sr and 39K. The addition of these multiple tracers that have anomalous composition facilitate experimental description because tracers added to the fluids and material of natural isotopic composition liberated from the rock mix and are incorporated into neogenic phases. Since the amount of 87Rb is different in every granite mineral it is possible to
2.2 Results Fi ure 1 shows the linear correlation between (' BSr/86Sr)solutjon and (84Sr/86Sr)solufjon isotopic ratios 357
dissolved plagioclase is 3 times higher then the mass of biotite and 10 times the mass of K-feldspar.
Table 1. Normative abundance (wt %), K, Sr, Rb concentrations (ppm) and strontium isotopic composition of whole rock and separated minerals.
Sr
Min.
Norm.
K
Biotite Oligocl. K-feldsp.
10 44 18
837 2.665 80 7 857 122 0.7074 12 127 698 363 0.7122
Whole R.
100
30
468
Rb
(87S1/86Sr)
193 0.7111
of the Sr in solution at 180°C. This correlation implies that granite fluid interaction can be described as a variable mixing proportion between two end-member having a constant isotopic ratio. Since the (84Sr/86Sr)ratio of the rock end-member is constant and e ual to 0.566, the dissolution endmember (87Sr/%Sr)d , representing the isotopic contribution of the dissolved minerals to the isotopic solution composition, can be evaluated. Figure 1, shows also that the rock end-member remains constants during the 400 days of investigation. The primary source of strontium is then constituted by a mineralogical assemblage that dissolve with constant proportion from the first 20 da s of interaction. Despite the variation of the (’ S~Y’~Sr)soiutionand
Figure 1. (s7Sr/86Sr)ratio as a function of (s4Sr/86Sr)ratio in solution at 180°C (Bio = biotite; K-F = potassium feldspar, T.R. = total rock; PI = plagioclase).
2
(87Sr/’6Sr)solution,the (87Sr/86Sr)d, is constant suggesting that the proportion of dissolved primary minerals remains constant from the early stage of the granite alteration. Fi ure 2 shows that (Rb/39K)so~ution and (41K/3K)solution are linearly correlated. Consequently Rb and 41Kisotopes are covariant with respect to 39K isotopes. Since the two potassium isotopes do not fractionate during terrestrial process, both Rb and K do not fractionate during the interaction and the variation of (Rb/K)solution ratio during time reflects, the variable mixing proportion between two endmembers with different (Rb/K) ratio: the initial solution (with no rubidium) and a dissolved mineral assemblage containing both Rb and K. The existence of both (87Sr/86Sr)dand (Rb/K)d end-members allows to estimate the proportion of dissolved minerals assuming that K, Sr and Sr are released stoichiometrically and that hornblende, apatite and sphene are negligible in the budget. The mass of biotite, K-feldspar and plagioclase dissolved during alteration can then be estimated assuming: a) the mass conservation law, b) the mixing equation for both strontium and potassium isotopic ratios, c) the mass conservation law with respect to the strontium isotopes and, d) the constancy of the (84Sr/86Sr)and (“Sr/%r) ratios (Seimbille et al. 1998). Irrespective of the temperature (in the range between 80 and 2OO0C), we found that the mass of
B
Figure 2. (Rb/39K)ratio as a function of (41K/39K)in solution at 180°C (PI = plagioclase; Bio = biotite; T.R. total rock; K-F = potassium feldspar).
3 MINERAL DISSOLUTION RATES To evaluate the dissolution reaction of every mineral we estimated the amount of potassium released to solution during the alteration of the silicate assemblage assuming that at every time of the reaction the (41K/39K)ratio in solution is in isotopic equilibrium with neogenic phases. Neogenic phases are mainly zeolite and clay minerals with an ‘open structure’. We assume also that the (41K/39K)ratio in
358
solution results from the mixing proportion between the amount of dissolved spike and the amount of dissolved potassium. We found that the potassium released from the rock, Kd, increases linearly as a function of time fiom the first 4-5 days of interaction with a rate of about 30 nmol/days. The contribution of plagioclase, biotite and Kfeldspar to the total released potassium can be estimated by the following equation:
4 DISCUSSION AND CONCLUSIONS
where Kbio, Kplag.and Korthcorrespond to the product between the amount of potassium in minerals and the dissolved fraction mass evaluated in the previous section 2.2. The rate of mineral dissolution is also a function of the reacting mineral surface area. The correct evaluation of this parameter is problematic and represents one of the main sources of discrepancy in the evaluation of kinetic rates of geochemical processes (White 1995). The reactive surface area is related to the density of sites on the mineral surface which are really exposed to the action of the aqueous solutions and could also change during the progression of the mineral dissolution reaction (Stilling et al. 1995). Such a condition in not easily verified in our molar scale measurements thus, we assume that the surface area of every mineral does not change significantly during the dissolution reaction. This assumption is consistent with the constancy of the proportion of dissolved mineral observed in the larger part of the investigated alteration. The mass fraction of every mineral was evaluated by the chemical composition of the bulk rock and is equal to 0.44, 0.10, 0.18 for plagiocalse, biotite and orthoclase respectively. We assume that for the same size of grinding, tectosilicates have similar specific surface area, while biotite have larger specific surface area (Holdren & Speyer 1985; Acker & Bricker 1992). The total surface area of the grain measured by BET is equal to 0.83 m2 g-'. Given the low mass proportion of biotite in our granite rock and that tectosilicates have half specific surface area then biotite, normalizing to the specific surface area assumed constant, we estimated that the rate of biotite and K-felspar dissolution is 10-13 mol m2 s-', while the rate of plagioclase dissolution is 10-I2mol m2 s-'.
359
The evaluation of dissolution rates of plagioclase, Kfeldspar and biotite during the hydrothermal alteration of a granite by fluids saturated with respect to the common mineral phases found in geothermal area is based on the multi-spiked system where a granite reacts with fluids of initial composition enriched or depleted in the less abundant Sr and K isotopes. Given the properties of both (Rb/K) and Sr-K isotopic ratios we describe the interaction by a mixing line between the initial solution and the rock end-member representing the mass proportion of dissolved minerals. Our dissolution rates estimated by the mixing properties of K and Sr are independent on the initial exchange reaction as this process is dominant in the first 2-3 hours of interaction (Lasaga 1995). Since diffusion of both oxygen and strontium in the silicate framework is negligible in our experimental conditions, the introduction of a new isotope in solution coming from feldspar dissolution is related to the destruction of K, Rb and Sr cationic sites. However, in biotite, potassium and strontium are located in the interlayer site and their release is not completely dependent on the dissolution of the alumino-silicate mineral. We found that the proportion of dissolved mineral remains constant from the first days of interaction even if the rock dissolves during the whole year of investigation. Since the precipitation of neogenic phases influences the budget of mineral dissolution reactions, we estimate the rate of potassium release assuming the isotopic equilibrium of the (4'W39K)ratio between solution and neogenic phases. This means that the amount of potassium coming from the mineral dissolution and trapped by neogenic phases corresponds to a maximum value. The proportion of dissolved mineral associated to the overall rate of potassium release allows to estimate that plagioclase dissolves only one order of magnitude faster than K-feldspar and biotite and that there is not a strong difference between dissolution rates of K-feldspar and biotite. REFERENCES Acker, J.P. & O.P. Bricker 1992. The influence of pH on biotite dissolution and alteration kinetics at low temperature. Geochim. Cosmochim. Acta 56: 30733092.
Holdren, G.R. & P. Speyer 1985. pH dependent changes in the rates and stoichiometry of dissolution of an alkaly feldspar at room temperature. Am. J. Sci. 285: 994- 1026. Lasaga, A. 1995. Fundamental approach in describing mineral dissolution and precipitation rates. In White A.F. and Brantley S.L. (eds), Chemical weathering rates of silicates minerals. Reviews in Mineralogy 3 1 : 23-81. Seimbille, F., Zuddas, P. & G. Michard 1998. Granitehydrothermal interaction: A simultaneous estimation of the mineral dissolution rate based on the isotopic doping technique. Earth and Pla. Sci. Lett. 157: 183191. Steiger, R.H. & E. Jager 1977. Sub-commission on geochronology: convention on the use of decay constants in geo and cosmo-chronology. Earth and Pla. Sci. Lett., 36: 359-362. Stilling, L.L., Brantley, S.L. & M.L. Machesky 1995. Proton adsorption at an adularia feldspar surface. Geochim. Cosmochim. Acta 59: 1473-1482 White, A.F. 1995. Chemical weathering rates of silicate minerals in soils. In White A.F. and Brantley S.L. (eds.) Chemical weathering rates of silicates minerals. Reviews in Mineralogy 3 1: 407-46 1. Zuddas, P. & G. Michard 1994. Experimental mineral fluid interaction in the Na-Ca-(Sr)-AI-Si system. Eur. J. Mineral. 5: 807-818. Zuddas, P. , Seimbille, F. & G. Michard 1995. Granite fluid interaction at near equilibrium conditions: Experimental and theoretical constraints from Sr contents and isotopic ratios. Chem. Geology 121: 145154.
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Mineral su$aces and weathering
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Glauconite Dissolution Rates and the Chemical Evolution of Vadose Waters in the Hornerstown Formation, Hornerstown, New Jersey Jill Betts Hahn and Associates, Inc., Portland, Oregon, 97209 USA
D .E.Grandstaff Geology Department, Temple University, Philadelphia, PA 19122, USA
ABSTRACT: The composition of vadose pore waters was measured as a function of depth in the glauconiterich Hornerstown Formation near Homerstown, New Jersey to determine the weathering rate of glauconite. The residence time of pore water within the vadose zone is 1 to 2 months, as determined by both temporal chemical variations and transport calculations. The calculated rate of glauconite weathering is 5 4.6 x 1O-I4 mol m-2s-',based on the geometric surface area of glauconite pellets. Based on net loss of all species, the rate of chemical denudation in the vadose zone is approximately 0.7 cm/lOOO yrs. 1 INTRODUCTION A number of studies have been conducted using mass-balance calculations to determine weathering rates of minerals or rocks in natural environments (e.g. Paces 1983, Velbel 1985, Swoboda-Colberg & Drever 1993, Sverdrup & Warfvinge 1995). Most of these studies have been conducted in catchments or drainage basins and have focussed on weathering rates of igneous or high-grade metamorphic minerals or rocks. In this study we have investigated chemical evolution of pore waters in a vadose zone in the glauconite-rich sediments of the Hornerstown Formation. The site is lithologically and texturally uniform, simplifying study of chemical evolution of the vadose waters. Mass-balance calculations have been used to determine reactions and processes affecting the evolution of the pore waters and the rate of glauconite weathering and chemical denudation.
2 FIELD AREA The field site is located in the northeastern United States, approximately 25 km southeast of Trenton and 2 k m west of Homerstown, New Jersey, along a tributary to Crosswicks Creek at 40' 06' 17.66" N and 74' 31' 35.49" W. The area is thickly forested. Beds are almost horizontal. The vadose zone is about 4.5 m thick, and is entirely within the Homerstown Formation (HTF). The HTF is upper
Cretaceous and lower Paleocene in age, and is a texturally uniform, massive, dark green, unlithified, intensely bioturbated glauconitic sand. The HTF consists of: clay 66% (dominantly glauconite with minor smectite and kaolinite), quartz 13%, muscovite 1%, and heavy minerals 20% (dominantly hematite and goethite). Minor amounts of carbonate apatite and calcite are also present. The glauconite major element composition is: (K0.6C~.02)(Mg0.32A10.29Fel.42)(A10.55Si3.45)0 10(OH)2. Total exchangeable cations are 24 meqA00 g. The glauconite is pelletized into sand-sized grains. The HTF is uniform coarse to fine sand, with minor silt and clay, and may be classified as a sandy loam. The porosity is fairly constant, ranging from 43 to 47%. Macropores occur in the upper part of the vadose zone. Because of the subdued topography and virtually horizontal permeable beds, ground water flow is primarily vertical. 3 RESULTS Precipitation, throughfall, vadose, ground, and stream water samples were collected 18 times from June 1996 to June 1997. Samples show some temporal variability. For example, chloride data from lysimeters at 0.5 and 4 m depth, from near the top and bottom of the vadose zone, are shown in Figure 1. Chloride concentrations from the 0.5m lysimeter reflect the amount and origin of precipitation and dry deposition. Because HTF minerals do not contain significant chloride, the
363
Figure 2. Concentrations of major species (January 7, 1997). The land surface is marked by horizontal lines in each panel. Depth scale in meters below the surface for lysimeter data. Pptn = precipitation, Tf = Through fall.
Figure 1. Temporal variations in C1 concentration from 0.5 and 4 m depth in the Hornerstown vadose zone. The variation pattern from the 4 m lysimeter data "lags" behind that from the 0.5 m lysimeter by about 1 to 2 months, suggesting that this is the residence time of fluid in the vadose zone
difference in C1 concentration between the 0.5 and 4 m lysimeters results from evapotranspirative loss of water. C1 concentrations (Fig 2) increase approx 3.7 times from 0.25 to 4 m depth. This suggests that almost 80% of water is lost within the soil profile. Most water loss (> 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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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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
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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.).
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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
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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.
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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
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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
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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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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
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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
5 CONCLUSIONS Significant relationship was observed between the uranium content and the amount of sand fraction in soil horizons developed over sandstone from Itarari Sub-Group, Parani basin, Brazil, that increases towards the more superficial soil horizons. 234U/238U disequilibria values of 0.8 and 0.9 found at the saturated zone allowed to suggest times between 4 10 and 660 ka for the development of the more superficial soil horizons during weathering of the sandstone parent rock.
410
REFERENCES Allman, M. & D.L. Lawrence 1981. Geological laboratory techniques. New York: Arc0 Publishing Co., 335 pp. Balasubramaniam, K.S., Moorthy, V.K. & B.R. Vyas 1983. Significance of engineering properties in understanding the proper utilization of laterites from western India. Irit. Seminar on Laterisation Processes 11: 577-590. Cherdyntsev, V.V. 197 1 . Uranium-234. Jerusalem: Israel Progr. For Sci. Transl., 144 pp. Colman, S.M. 1982. Chemical weathering of basalts and andesites: evidence from weathering rinds. U. S. Geol. Surv. Prot Paper 1246,51 pp. Colman, S.M. & D.P. Dethier 1986. Rates of chemical weathering of rocks and minerals. New York: Academic Press, 603 pp. Faure, G. 1991. Principles and applications of inorganic geochernistry. New York: MacMillan Publishing Co. Fleischer, R.L. 1975. On the dissolution of respirable Pu02 particles. Health Physics 29: 69-73. Hansen, R.A. & R. Stout 1968. Isotopic distributions of uranium and thorium in soils. Soil Science 105: 44-50. Ivanovich, M. & R.S. Harmon 1982. Uranium series disequilibriuni: applications to environmental problems. Oxford: Clarendon Press, 57 1 p Kigoshi, K. 1971. Alpha-recoil 2’iTh: dissolution into water and the 234U/2”Udisequilibrium in nature. Science 173:47-48. Langmuir, D. 1978. Uranium solution-mineral equilibria at low temperatures with applications to sedimentary ore deposits. Geochirn. Cosmochim. Acta 42: 547-569. Latham, A.G. & H.P. Schwarcz 1986. Models of uranium etching and leaching: their applicability to estimating rates of natural uranium removal fiotn crystalline igneous rocks using U-series disequilibrium data frotn the Eye-Dashwu Lakes pluton. Canadian Atomic Energy (Rep. TR-353),60pp. Latham, A.G. & H.P. Schwarcz 1987. On the possibility of determining rates of removal of uranium from crystalline igneous rocks using U-series disequilibria-1: a U-leach model, and its applicability to whole-rock data. Applied Geochenzistry 2: 55-65. Michel, J. 1984. Redistribution of uranium and thorium series isotopes during isovolumetric weathering of granite. Geochim. Cosmochim. Acta 48: 1249-1255. Moreira-Nordemann, L.M. 1980. Use of 234U/238U disequilibrium in measuring chemical weathering rate of rocks. Geochim. Cosnzochim. Acta 44: 103-108. Osmond, J.K. & J.B. Cowart 1976. The theory and uses of natural uranium isotopic variations in hydrology. At. Etiergy Rev. 14: 621-679. Schellmann, W. 1979. Considerations on the definition and classification of laterites. Itit. Seminar on Laterisatiori Processes I: 1-10, Szalay, A. 1964. Cation exchange properties of humic acids and their importance in the geochemical enrichment of U O F and other cations. Geochirn. Cosmochinz. Acta 28: 1605-1614.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Water-Rock Interaction and the Water Chemistry of a Small Sierra Nevada Lake Dana Sue Kimbal Newinont Mining Co., Twin Creeks Mine, HC 66, Box 69 Golconda, Nevada, U S A .
Ross WSmith University of Nevada, Reno, Metallurgical and Materials Engineering, MS 388, Reno, Nevada, U S A .
ABSTRACT: The interaction of water of initial pH 4, 5, and 7 with ground minerals typically present in a granite or granodiorite plus rock samples taken from the drainage of a small California Sierra Nevada lake was studied. The dissolution of metal ions and the change in pH as a function of time was monitored. The pH and metal ion concentrations thus determined were those that would be expected from the weathering and dissolution of the minerals present in the rocks of the drainage. Furthermore, they were consistent with measurements previously taken by Stoddard (1987) of the composition of the lake water at various times of the year.
natural occurrence of sulfur oxides from natural volcanic activity, atmosphere water should be at about pH 5.4 (Galloway et al. 1982). More acidic precipitation should usually have an anthropogenic source.
1 INTRODUCTION Gem lake is a small (surface area 2.7 ha), shallow (mean depth 3.4 m), high elevation (3,341 m) lake in Inyo County in the California Sierra Nevada. It is located at tree line in a rocky, barren environment mainly consisting of granodiorite talus. There is very little biomass in the drainage, in part because of the high altitude and in part due to the highly seasonal, although substantial, precipitation pattern in the Sierra Nevada of California. The watershed area is approximately 250 ha. This is a larger area than previously reported by Stoddard (I 987) since it was found, by on site inspection, that a higher, alpine lake, Dade lake, actually drains into Gem lake. There are several, small, remnant glaciers at the head of the drainage into Dade lake. The alkalinity dynamics of the lake were studied in detail by Stoddard (1987, 1988). Lake water samples were taken during every month of the year. It was found that the pH in the lake increases dramatically during snowmelt to much greater values than that of the snow itself (about pH 5.5). In addition, the metal ion dynamics of the lake were studied as a function of time. Measurements made in the present study indicated that the pH values in the lake and the stream flowing into the lake were consistent with the measurements of Stoddard (1987). The pH of water in equilibrium with carbon dioxide of the atmosphere should be at pH 5.6-5.7. There is evidence, however, that because of the
2 EXPERIMENTAL In the present paper the reaction of rock samples gathered from the watershed and individual minerals found to be present in the rock samples were reacted with water of initial pH values of 4.0, 5.0 and 6.8. The minerals and rock samples studied are listed in tables 1 and 2. The acidic pH waters were obtained by adding H2S04. The pH and various metal ion concentrations were then studied as a function of time (up to 200 days). Two different size fractions were studied (- 250, + 150 pm and - 150, +75 pm). In the experimentation 25.0 k 0.1 g of the mineralhock sample was placed in 250 ml of water (open to the atmosphere) in a 300 ml Nalgene plastic bottle with a lid at the selected pH and allowed to age under quiescent conditions. At selected time intervals water samples ( 5 ml) were withdrawn for analysis. Then, with hornblende and albite an additional experiment was performed. After reaction with the aqueous solutions, the mineral samples were filtered, and dried and the experiments were repeated with 250 ml fresh water open to the atmosphere. 41 1
Table 1 Water pH as a function of time in the presence of added mineral and rock samples
Mineral or Rock
Particle Size (pm) -210, +150
I
I -210, +150
I -210, +I50
I Albite (rewash)
-210, +I50
-210, +I50
pH (initial) 4.0 rt 0.1 5.0 rt 0.1 6.8 f 0.2
4.0 f 0.1 5.0 f 0.1 6.8 f 0 . 2 4.0 rfr. 0.1 5.0 f 0.1 6.8 f 0 . 2 4.0 rt 0.1 5.0 f 0.1 6.8 f 0.2 4.0 f 0.1 5.0 rt 0.1 6.8 rfr. 0.2
I
I I I
pH (24 hr) 9.2 9.4 9.2
7.8 9.4 9.5 9.2 9.2 9.2 7.3 9.3 9.3 9.9 10.0 10.2
I
I I I
pH (400 hr) 8.5 8.5 8.8
8.9 9.2 8.8 9.0 8.6 8.8 8.7 9.4 8.6 9.8 9.8 9.9
I
I I I
pH (2400 hr) 8.3 8.4 8.4
8.5 8.6 8.0 9.2 8.3 8.3 8.0 8.2 8.2 9.5 8.3 8.3
5.0 f 0.1 Enstatite MgSi03
I
Muscovite KA12Si3A1010(0H)2
-150, +75
Quartz aSi02
-150, +75
Diorite (rock sample from drainage) Olivine (mineral sample from drainage) Granite (rock sample from drainage)
-150, +75
-150, +75
-210, +150
5.0 f 0.1 6.8 f 0 . 2 4.0 rt0. 1 5.0 k 0.1 6.8 f 0.2 4.0 f 0.1 5.0 +_ 0.1 6.8 f 0.2 4.0 f 0.1 5.0 rt 0.1 6.8 f 0.2 4.0 f 0.1 5.0 f 0.1 6.8 f 0.2 4.0 f 0.1 5.0 f 0.1 6.8 rt 0.2
9.9 9.9 9.9 9.1 9.2 9.1 5.2 6.3 7.0 6.9 7.9 8.2 8.8 9.4 9.4 7.9 8.7 8.8
I
I
9.9 9.7 9.7 9.3 8.9 9.0 8.9 6.1 6.5 6.8 7.3 7.5 7.5 8.9 9.1 9.1 7.7 7.9 8.0
I
8.6 8.1 8.2 8.6 8.2 6.0 6.6 7.2 7.0 7.1 7.0 7.4 8.3 8.0 7.0 7.4 8.1
3 RESULTS
4 CONCLUSIONS
Table 1 shows pH of the solutions as a function of aging time after addition of the mineralhock samples. Table 2 shows the metal ion concentrations of the solutions at selected aging times.
It was found that the minerals are able to substantially and quickly raise the pH of mildly acidic water, both freshly fractured minerals and those that had been in contact with water, dried and re-reacted with water. The pH and metal ion 41 2
Table 2 Metal ion concentrations (ppm) in solution as a function of initial pH values in the presence of the minerals and rocks studied Mineral or
(rewash)
(rewash) Augite
Diopside ~
r Muscovite
sample from drainage) (mineral sample from
sample from drainage)
concentrations are consistent with actual field measurements of Stoddard (1987). The data confirm the findings of Stoddard that rock-water interaction and natural weathering processes are able to neutralize acidic water falling in the form of snow and rain (1987, 1988). In addition, the data obtained are consistent with granite weathering data (of recently glaciated areas) as tabulated by Stauffer (1990) and with the dissolution reactions of minerals with water and the regulation of the chemical composition of natural waters (Stumm and Morgan 1970, Hem 1985). The findings are also consistent with measurements of the chemical composition of spring waters from granitic regions of -the Sierra Nevada (Feth et al. 1964). The effect of organic material within the lake on the chemical composition ofthe lake was not investigated because ofthe rapid turnover of water in the lake at peak snowmelt (Stoddard 1985).
5 SUMMARY The ability of the silicate minerals and the rock samples to neutralize acidic water (pH 4.0 - pH 6.8) has been documented. Furthermore, both alkali and alkaline earth cations appear in the leach solutions. In a geologic environment where both weathering due to high rain and snow fall, although seasonal, and rock fracture due to rockslides and freezing and thawing, the buffering capacity of the granitic rocks should be substantial and rapid. In particular, the active fracturing should extend the buffering capacity of the rock minerals to long periods of time. REFERENCES Feth, J.H., Roberson, C.E. & W.L. Polzer, 1964. Sources of Mineral Constituents in Water from Granitic Rocks, Sierra Nevada, California and Nevada. U.S.Geologica1 Survey Water Supply Paper 1535-1: 70 pp.
413
Galloway, J.N., Likens, G.E., Keene, W.C. & J.M. Miller, 1982, The Composition of Precipitation in Remote Areas of the World. JGeophys. Res. 87: 8771-8786. Hem, J.D., 1985, Study and Interpretation of the Chemical Characteristics of Natural Water, 3rd Edition. U.S. Geological Survey Water Supply Paper 2254. Stauffer, R.E., 1990. Granite Weathering and the Sensitivity of Alpine Lakes to Acid Deposition. Limnol. Oceanogr. 35: 1112-1134. Stoddard, J. L., 1987. Alkalinity Dynamics in an Unacidified Alpine Lake, Sierra Nevada, California. Limnol. Oceanogr. 32: 825-839. Stoddard, J. L., 1987. Are Sierran Lakes Different? A Reply to the Comments of Kelly and Schindler. Limnol. Oceanogr. 33: 164 1- 1646. Stumm, W. & J.J. Morgan, Aquatic Chemistry. Wiley Interscience. Chapter 8 (1-4): 383-341.
414
Water-RockInteraction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The S. Antioco of Bisarcio Basilica (NESardinia, Italy): water-rock interaction in ignimbrite monument decay G .Macciotta, G .Bertorino, A Caredda, S .Columbu, M .Franceschelli & M. Marchi Dipartimento di Scienze della Terra, Cagliari, Italy
S .Rescic Centro Studio Cause di Deperimento e Conservazione delle Opere d’Arte, CN.R., Firenze, Italy
R .Coroneo Dipartimento di Scienze Archeologiche e Storico-Artistiche, Cagliari, Italy
ABSTRACT: The Basilica is constructed of pyroclastic rock dating from the calcalkaline-affinity volcanic cycle occurring in Sardinia between 32 and 11 Ma ago. Very small amounts of epiclastic products have also been detected. Petrographic and geochemical data indicate that the composition of these volcanic rocks ranges from dacite to rhyolite. Variations in total open porosity in the external part of the hewn rock permitted the singling out of two extreme trends pointing to reprecipitation within the pores and surface disaggregation. Visible signs of weathering include exfoliation, enucleation and alveolation, as well as chromatic surface alterations, with organic films and salts efflorescence formed as a result of water rise and evaporation.
1 INTRODUCTION The S. Antioco of Bisarcio Basilica, built between the late XI and early XI11 century, was diocese cathedral from the end of the XI to the beginning of the XVI century, when the see was suppressed. Longitudinal in plan, the church has an apse and three naves separated by pillars topped with Romanesque capitals. The original facade was overlain with a two-storey portico. Its Romanesque construction can be distinguished from its walls, built of medium to large ashlars, carefully hewn and lain using the technique adopted by Tuscan builders, employed in the Kingdom of Torres since the XI Century. The hewn stones were dry laid and carefully aligned, but did not display the refined features distinguishing Cistercian buildings. The church, situated a few kilometres outside the village of Chilivani in NE Sardinia, was built with pyroclastic rock of the Oligo-Miocene volcanic cycle, whose products outcrop in vast areas of Sardinia.
Volcanic activity began around 32 Ma ago producing basaltic and andesitic lavas and ended about 13-11 Ma ago, climaxing between 23 and 17 Ma. From 23-22 Ma onwards (Beccaluva et al. 1985 and references therein), highly explosive ignimbritic fissure emissions of dacitic-rhyolitic composition occurred simultaneously and with alternating basaltic and andesitic lavas in various parts of the Island, mainly along the western graben. The pyroclastic rock used for constructing the Basilica can be attributed to this latter volcanic activity (21.2*0.8 Ma, Lecca et al. 1997, locality S. Antioco of Bisarcio). Epiclastic rock, occasionally used in masonry for decorative purposes along with small amounts of green sandstone, occurs to a lesser extent in the area. 2. I Petrographicfeatures Field surveys conducted for the purpose of identifying the rocks used in construction pointed to the presence of vast pyroclastic flows (Savelli et al. 1979, Lecca et al. 1997, Morbidelli et al. 1999) of extremely variable composition. This material was also compared with the pyroclastic rocks occurring in the nearest and easternmost area, Oschiri. Based on their volcanological, petrographic and geochemical features, the samples were divided into three main groups: 1) welded pyroclastic rocks (including lava-like ignimbrites) from the Chilivani area (WPC); 2) unwelded pyroclastic rocks, again from Chilivani (UPC); 3) pyroclastic rocks from the Oschiri area (OP).
2 OLIGO-MIOCENE VOLCANICS Sardinian Oligo-Miocene volcanism has a calcalkaline affinity 1.s. and is generally related with a subduction zone dipping N-NW along the continental European paleomargin that produced the Oligocene rift between Sardinia and Provence (Cherchi & Montadert 1982). 41 5
varies from 5 to 15%, even within the same cooling unit. Welded pumice is rarely found or absent altogether. More or less welded lapilli are always present and angular or rounded cognate fragments vary from 2 to 7%. Gabbro nodules or granoblastic fragments of the basement are occasionally observed. The phenocrysts are, in order of segregation spinel, plagioclase, f orthopyroxene, clinopyroxene, ISfeldspar, quartz and, rarely, horneblende. Figure 1. RI vs. R2 diagram (after De La Roche et al. 1980, detail). Stars = WPC (field campaign); triangles = UPC (field campaign); diamond = WPC (church); squares = UPC (church); circles = OP.
2.1.1 Classijkation of volcanic rocks Approximately one hundred chemical analyses have been performed by means of XRF spectrometry on major and trace elements of samples collected in the field and taken from the church itself. These have been classified according to De La Roche et al. 1980 (Fig. 1). The samples from Chilivani are composed of quartzlatite (WPC), rhyolite, rhyodacite and rare quartztrachite (UPC, field samples) and of rhyodacite, rhyolite, alkali-rhyolite and secondarily of quartzlatite and quartz-trachite (UPC, samples from the church). Note that the so-called alkali rhyolite shows an agpaitic index < 1 (A.I. 0.86-+0.90). The pyroclastic rocks from the Oschiri area (OP) vary from secondary rhyodacite to abundant rhyolite, with minor amounts of quartz-latite and quartz-trachite. Some differences in terms of geochemistry emerged among the different rock types and areas. In particular, WPC f OP exhibit, on average, higher Ti02, A1203, FeO,,,, Na20 (Fig. 2) and lower K20 contents than the UPC as well as slightly lower Zr, Nb, Y, La and Ce concentrations.
2.2.2 Unweldedpyroclastic rocks Unwelded pyroclastic rocks are more varied as regards the incidence of phenocrysts, lithic fragments and, above all pumice. Phenocrysts include, in order of segregation, spinel, plagioclase, f orthopyroxene, clinopyroxene, i biotite, K-feldspar and quartz. Biotite, when present, is often deformed and partially converted into minerals belonging to the chlorite group. 2.3 Late-stage alteration processes Late-stage alteration processes transformed the original paragenesis to a greater or lesser extent, altering mineral assemblage with the occurrence of zeolite, silica phases (e.g. opal-CT), hematite, etc.. Preliminarily XRD analyses indicated that zeolite is present solely in the form of mordenite, while clinoptilolite is absent, contrary to what has been observed in other parts of Sardinia (Ghiara et al. 1997). 3 PHYSICAL PROPERTIES 3.1 Analytical methods Real density (a) and bulk density (a,) have been calculated on the dry weight, real volume (helium picnometer) and bulk volume (mercury picnometer). The latter measurements have been used to calculate open porosity (PO).The imbibition coefficient (IC,) and saturation index (SI) have also been calculated based on the weight of soaked specimens.
2.2 Microscopicfeatures 2.2.1 Weldedpyroclastic rocks Welded pyroclastic rocks have extremely variable petrographic characteristics. The porphyritic index
3.2 Data interpretation
Figure 2. D.I. (differentiation index) vs Na20 diagram showing overlapping between WPC + OP and wide scatter for UPC. Symbols as in Figure I .
As a whole, data obtained for the samples analysed show highly scattered values, which may be attributed to the great heterogeneity of the rock as the result of different eruption conditions and emplacement temperatures (high- to medium- to low-grade ignimbrites). Scattering also depends on the different incidence of pumice, crystal and lithic fragments and matrix. Aside from the late-stage syngenetic alteration processes mentioned above, the weaklywelded pyroclastic rocks are affected to a greater extent by epigenetic alteration, while the majority of strongly welded ones are only weakly altered and major differences between those parts more or less exposed to atmospheric agents are seldom observed. 416
The fact that the weakly-welded pyroclastic rocks suffer weathering more lies in their greater porosity, and hence, weaker welding and lower vitrification. By correlating certain physical properties (a, a, and PO)with geo-volcanological and petrographical data, it was possible to divide the pyroclastic rocks used for constructing the Basilica into three main groups for interpreting physical parameters: a) strongly welded, seldom used; b) moderately welded, widely used; c) weakly welded, used in small quantities. The mean bulk density (2.13t2.20 g/cm3) of pyroclastic rocks belonging to the first group is far higher than the other groups (Table 1). Examination of thin sections cut parallel to the direction of weathering penetration show that the patina usually forms only on the outer surface of the ashlar ( 1I mm), where open porosity is lower and bulk density higher (Table 1). A two-stage model may be advanced to explain this behaviour: initially the surface of hewn blocks exposed to the atmosphere undergoes chiefly physical breakdown, namely slight devitrification of the irregularly arranged vitreous matrix and microfiactures produced by thermal stress. Initially the increase in porosity facilitates the penetration of percolating meteoric water resulting in reprecipitation of substances including plasters, lime milk, etc. and subsequent reduction in porosity. A significant example is sample A55 (removed from the east-facing outside wall of the apse). XRD analyses showed the outermost part to contain considerable amounts of gypsum, reflected in the significant increase in CaO and LOI (Table 2). The different real density values determined for the dry material also indicate the presence of different phase assemblage (reprecipitation) between the outer crust and the inner matrix.
Table 2. Sample A55: chemical analyses of inner part (I) and outer part (E) and relative percent variations A; nd = not detected.
I 65.20 0.48 12.47 3.64 0.55 0.08 0.30 2.15 1.23 8.81 2.09 2.99
a,
PO
I
2.59
2.13
17.9
13.8
78.0
E
2.55
2.20
13.7
10.8
85.1
3.29 -23.5 -21.7
9.1
d
strongly welded
-1.54
variations
IC,
SI
-
moderately welded
I
POI < P OE
E
2.68
1.66
37.6
38.7 102.8
2.61
1.54
41.0
34.7
-2.61 -7.23
variations
84.3
8.8 -10.2 -17.9
moderately welded
1
2.65
1.52
43.1
37.3
86.9
POI > 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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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
Water-Rock lnferaction 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
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.
437
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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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
'Averaged over 5 points; "Averaged over 3 points.
Fig. 5). Minor alteration was also found along cracks and in small patches. SEM-EDX analyses of the unaltered and altered areas are given in Table 1. The data show that the unaltered brannerite contains approximately 4 wt% CaO, 3 wt% PbO, 1 wt% each of Fe0 and Th02, and minor amounts of Al2O3, Y2O3, and lanthanides. The oxides P2O5, Si02, and MnO are generally near or below detection limits, i.e. of about 0.1 wt%, in the unaltered brannerite. Compared with the unaltered brannerite, the altered areas have generally lost U02 and CaO and have gained P2O5, Si02, Th02, Al2O3, and FeO. The chemical U-Th-Pb age was > 400 Ma from SEM-EDX analysis of the unaltered areas. TEM examination has shown that the brannerite structure was rendered amorphous by radiation damage. 3.4 Dissolution of natural brannerite The elemental releases from the natural and synthetic samples over 56 days are summarized in Table 2. For pH 4 solution, the elemental releases for U, Y, A1
Fig. 6. SEM secondary electron micrographs of natural brannerite samples before and after testing (non-replacement in 56 days), a) before; b) in pH 4 solution; c) in deionised water.
44 1
Table 2. Elemental releases (%) from natural and synthetic brannerites in 56 days at 90°C without leachant replacement. Natural Element A1 Ca Th Ti Pb U Y
Synthetic
pH 4
diw'
pH 4
5.3 5.1 0.007 0.002
1.2 0.23 0.0002
-
0.0005 0.002
0.00006
0.002 0.002
0.9
0.03 4.5 6.6
-
and Ca are quite similar, about 5-6%. There is preferential release of U over Ti. The U release from synthetic brannerite to pH 4 solution is -5 times lower than natural brannerite, indicating as a first estimate that the effect of radiation damage on U release is less than an order of magnitude. Compared to pH 4 solution, the elemental releases to deionised water are generally lower, particularly for U and Y, by a factor of (2-3)x103.
3.5 Surface characterization of natural samples after testing SEM images of the samples before and after dissolution tests are shown in Figure 6. The surfaces of the natural samples after 56 days of testing (Fig. 6 ) show a much higher amount of Ti-rich phases in pH 4 solution compared to deionised water, consistent with the greater release of U at pH 4. 4 CONCLUSIONS
REFERENCES Grandstaff, D.E. 1980. The dissolution rate of Forstentic Olivine from Hawaiian sand. 3rd International Symposium on Water-Rock Interaction, Edmonton, Canada, July 14-24, 72-74. Ifill, R.O., Cooper, W.C. & A.H. Clark 1996. Mineralogical and process controls on the oxidative acid-leaching of radioactive phases in Elliot Lake, Ontario, uranium ores: IIBrannerite and allied titaniferous assemblages. CZM Bulletin, 89(1001), 93-103. Jostsons, A., Vance, E.R. & B. Ebbinghaus 1999. Immobilization of surplus plutonium in titanate ceramics. Global 99, Jackson Hole, Wyoming, USA, Aug. 29 to Sep. 3. Knauss, K.G. & T.J. Wolery 1986. Dependence of albite dissolution kinetics on pH and time at 25OC and 7OOC. Geochim. Cosmochim. Acta, 50,2481-2497. Lasaga, A.C. 1998. Kinetic theory in the earth sciences, Princeton University Press. Luce, R.W., Bartlett, R.W. & A.P. George 1472. Dissolution kinetics of magnesium silicates. Geochim. Cosmochim. Acta, 36,35-50. Lumpkin, G.R., Leung, S.H.F. & M. Colella 1999. Composition, geochemical alteration, and alpha-decay damage effects of natural brannerite. Mat. Res. Soc. Symp. Proc. 608,359-365. Lumpkin, G.R., Smith, K.L., Blackford, M.G., Giere, R. & C.T. Williams 1994. Determination of 25 elements in the complex oxide mineral zirconolite by analytical electron microscopy. Micron, 25(6), 581-587. McGlinn, P.J., Hart, K.P., Loi, E.H. & E.R. Vance 1995. pH dependence of the aqueous dissolution rates of perovskite and zirconolite at 9OoC. Mat. Res. Soc. Symp. Proc. 353, 847-854. Thomas, B. 2000. Brannerite dissolution modeling report. Unpublished data at ANSTO. Vance, E.R., Day, R.A., Zhang, Z., Begg, B.D., Ball, C.J. & M.G. Blackford 1996. Charge compensation in Gd-doped CaTi03. J. Solid State Chem., 124,77-82. Vance, E.R., Watson, J.N., Carter, M.L., Day, R.A., Lumpkin, G.R., Hart, K.P., Zhang, Y . , McGlinn, P.J., Stewart, M.W.A. & D. Cassidy 2000. Crystal chemistry, radiation effects and aqueous leaching of brannerite, UTi206. Ceramic Trans., 107,561-568.
The dissolution of brannerite is incongruent and pH dependent with a minimum at near pH 8. Preferential release of U leaves TiOz on the surface with different surface morphologies depending on solution acidity and alkalinity. The measured apparent activation energies at pH 5.6 to 9.8 suggest surface reaction-controlled dissolution mechanisms. The natural brannerite sample has been shown to be resistant to dissolution over geological time. Amorphisation leads the release of U to increase by less than an order of magnitude at 9OoC and pH 4. ACKNOWLEDGEMENT We acknowledge the partial financial support of this work by the USDOE in conjunction with LLNL, and thank E. Drabarek for surface area measurements, E. R. Vance, J. Bartlett and A. Jostsons for helpful comments.
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Groundwater environments
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Shallow groundwater in the Sebou basin (Northern Morocco) T.Bahaj, M.El Wartiti & MZaharaoui Earth Sciences Department, University Mohamed r! Rabat
R.Caboi & R.Cidu Earth Sciences Department, University of Cagliari
ABSTRACT: Different water types have been distinguished in the shallow aquifers of the southern part of the basin of Sebou river in Morocco. The low salinity Ca,Mg-HC03 type is related to the interaction with dolomite-limestone fiom the Middle Atlas; in the Miocene and Pliocene marls and sandstones, the waters evolve to the more saline Na(Ca)-C1 type. Sulphate ions contribute significantly to the anionic composition in the most saline waters, and may be related to the dissolution of gypsum from evaporite levels. Some minor and trace elements such as Br, F, B, U, alkaline and earth-alkaline elements, are linked to salinity, while the heavy metals follow trends which appear related to pollution sources. The chemical quality is in general good enough but it is often lowered by pollution due to the uncontrolled discharge of waste from factories. Middle Lias age, and marl-limestone of the Upper Lias-Dogger. Tertiary sediments occupy a large part of the Oued Sebou basin. They mainly consist of grey marl of Mio-Pliocene age, which may reach a thickness of 2000 m in the Gharb plain. The Pliocene is characterised by sands, with thicknesses of 60 m in the Sais, lacustrine limestone and conglomerate outcroppings in the lower part of the Sebou that decrease in thickness from the Gharb to Fks-Meknks. The Quaternary sediments (sands intercalated with clay and rnarl levels) are particularly important in the Gharb, where the dominant sand and gravel levels may reach a thickness of 250 m, thus forming an important aquifer.
1 INTRODUCTION The Oued Sebou basin in north-western Morocco extends for 40,000 km2; its total population is approx. 5 million inhabitants, 40% of whom live in towns, such as Fes and Meknes. With the growth of the population and of economic activities in the past decades, the demand for water, mainly for irrigation in agriculture and for urban purposes, has increased. Recently, during the long periods of dry weather, the region has oRen resorted to groundwater, and shallow groundwater and streams are used quite extensively by small villages and farms. This paper is aimed at assessing the quality of shallow groundwater in the Sebou basin, with particular attention to the southern sector, which is characterised by the presence of important “imperial towns” and by a more intense agricultural and industrial activity.
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2
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3 HYDROGEOLOGY The different lithologies occurring in the Sebou basin give rise to several distinct hydrogeological sectors. In the Middle Atlas, the Jurassic sequence has thickness of 250-300 m (Essahlaoui 1997), and represents an important aquifer in the area, the fractured and karstified dominant carbonate rocks allowing water infiltration at depth; the phreatic level may vary depending on the morphology, structural features, and precipitation. In the Fks-Meknks plateau it is possible to distinguish several reservoirs; the water-table hosted in sand, conglomerate, and lacustrine limestone is unconfined, but locally a multi-layer aquifer occurs due to intercalated marl and clay levels (Chamayou et al. 1975). The multi-layer aquifer is recharged by
GEOLOGY OF THE STUDY SECTOR
The Oued Sebou basin shows a complex geology with lithological formations ranging from the Palaeozoic to the Quaternary (Fig. 1). The Palaeozoic basement is made up of metamorphic rocks, such as schist and quartzite, and magmatic rocks, dating fi-om the Cambrian to the Permian (Michard 1976), and outcrops in the Middle Atlas and in the Meseta. Mesozoic sediments, lying discordant to the basement, outcrop in the Middle Atlas and consist of Triassic marl and clay containing gypsum and salt, dolomite and limestone of Lower445
Figure 1. Schematic map of the geology of the Sebou basin (modified fiom Atlas du Bassin du Sebou 1970).
rainfall and by groundwater fiom adjacent deep or shallow aquifers. Faults trending NE-SW divide the deep aquifer into separate blocks. In the Zemmour, Gharb, and Mamora sectors, the aquifer is hosted in sand and gravel sediments. The Miocene-Pliocene marl sequence f o m the aquitard, while the Quaternary clay and marl levels intercalated in the sands give rise to a multi-layer aquifer (Kacimi 1992, Aberkan 1989). The coastal dunes allow infiltration of rainwater, which feeds the shallow and deep aquifers. Additional contributions to the aquifer derive fiom oueds and drains. In this area the water table level may vary greatly, due to irregular rainfall and intense exploitation (Kacimi 1992). 4
SAMPLING AND ANALYTICAL METHODS
were prepared in the field using high-purity water and following the same acidification procedure as for the samples. Temperature, pH, redox potential (Eh), conductivity and alkalinity were measured at the sampling site. W+/NH3,and NO2- species were measured by spectrophotometry within a few hours after collection. In the laboratory, anion species were determined by ion chromatography on filtered, nonacidified samples. Metals were analysed by ICP-OES and ICP-MS. Mercury, antimony and arsenic were determined by ICP-MS after flow-injection Hg-vapour or hydride generation. 5 WATERTYPES
Sampling campaigns were performed in April 1999 and March 2000. During the campaigns, 52 shallow groundwater samples were collected randomly, but the dserent lithologies and hydrogeological characteristics, as well as the potential influence of human activities were taken into account. The sampling location is shown in Figure 1. Water samples were filtered in situ through a 0.4 pm pore-size membrane filter, using all-plastic equipment, and collected into acid-cleaned, high-density polyethylene bottles. Sample stabilisation with suprapure acid was carried out in situ upon filtration. Samples were acidified to 1 % HNO3 for the analysis of most metals, 0.2 % HC1 for As and Sb, and 0.2 % H2S04 plus KMnO4 crystals for Hg. In order to monitor potential contamination during either sampling or analysis, blank solutions 446
The main cationic and anionic characteristics are shown in Figure 2. We can distinguish different water types. The shallow groundwaters collected in the Middle Atlas sector appear very homogeneous: the total dissolved solids (TDS; mean: 292 mg/l, standard deviation: 34 mg/l) are low with earth-alkaline bicarbonate ions by far prevailing over other dissolved ions; the Mg/Ca ratio is very close to 1, and is referred to the interaction with the dolomites of the area. Shallow groundwaters in the Miocene of the Fes-Meknks plateaux show an evolutionary trend (TDS in the range of 250-1400 mg/l) that depends both on the length of the interaction with the rocks and on the lithologies of the water-bearing layers. In the lacustrine limestone we found a bicarbonate calcium type, whilst, by flowing into the sandstones,
Figure 2. Main cation and anion composition in waters from different areas of the Sebou basin. Middle Atlas: 1; Meknes plateau:2; Fbs plateau: 3; Gharb plain: 4; Zemmour region: 5; Mamora plain: 6 .
salinity increases, and the HC03 and C1 anions and the Na and Mg cations dominate, as calcium concentrations are controlled by saturation with respect to calcite. Two types of shallow groundwaters were characterised in the Mamora-Gharb plain and in the Zemmour region: i) calcium bicarbonate water of a lower salinity (TDS: 0.3 w 1 g/l), and ii) chloride (sulphate) alkaline water with salinity increasing up to 3.4 g/l. The calcium bicarbonate water is related to the aquifer of sand and gravel of the Mamora sector, while the chloride (sulphate) alkaline water is related to chaotic Quaternary sediments (clay, mark pelites, and salt domes) in the Gharb plain. Acquisition of calcium and bicarbonate ions during the evolution of the waters reaches = 200 mg/l Ca and 600 mg/l HC03, respectively. Sulphate ions contribute significantly to the anionic composition of the most saline waters (Zemmour region), and may be related to the dissolution of gypsum fiom evaporite levels. Ammonia and nitrite ions are lower than the detection limits, while NO3 values are fiequently higher than 50 ppm and at times reach 164 ppm in the studied water. These concentrations are probably due to the extensive use of fertilisers in the more developed agricultural areas (El Wartiti et al. 2000).
concentration of minor and trace elements is very low (even below the ICP-MS detection limit). Higher values both of rare alkaline and alkaline earth elements are found in the chloride (sulphate) alkaline water related to the Quaternary sediments (clay, marls, pelites and salt domes) in the Gharb plain. Elements such as Cd, Ag, Sb, and Be in most samples are lower than the detection limits for the analytical methods used. Other metals such as Zn, Cu, Cr, Ni and Pb follow different trends that are not linked to salinity: they appear related to pollution sources (Fig. 4 and 5). With the exception of chromium, the concentration of heavy metals does not usually exceed the recommended values for drinkable-water. However, a very strong pollution has been locally found; one sample, which was not included in the Figures, shows the influence of polluting sources on shallow groundwater. This water receives the discharge of an olive-oil mill waste, and shows a very high salinity
6 MINOR AND TRACE ELEMENTS Some minor and trace elements such as Br, F, B, U, rare alkaline, and earth-alkaline elements, are linked to salinity (Fig. 3), and therefore appear to be derived fiom rock weathering. Lower values are found in the carbonate aquifer of the Middle Atlas where, owing to the low degree of interaction between water and carbonate rocks, the
Figure 3. Rare alkaline and earth-alkaline elements in shallow groundwater of the Sebou basin.
447
REFERENCES Aberkan, M. 1989. Etude geologique des formations quaternaires des marges du bassin du Gharb (Maroc NW), These d'Etat, Univ. Bordeaux I, pp. 290. Atlas du Bassin du Sebou 1970. Livret explicatif, M.A.R.A., Maroc, pp. 143. Chamayou, J., Combe, M., Genetier, B. & C1. Leclerc 1975. Le bassin de Fbs-Meknks. Plaines et bassin du maroc atlantique. Notes et Mdm. Sew. Gdol. Maroc 23 1: 4 1-7I. El Wartiti, M., Zahraoui, M., Bahaj, S., Saddiki, N., Bahaj, T., Caboi, R, Cidu, R. & A. Carcangiu 2000. New hydrogeochemical data on the Oued Sebou basin (Morocco). Abstracts 5th International Conference on the Geology of the Arab World, Cairo University, Egypt, February 2000: 226-227. Essahlaoui, A. 1997. Etude par prospection geoeletrique dam le plateau de Meknes et essai de reconnaissance du bassin hydroghlogique de Sais. Mem. C.E.A., U.F.R. Geol. Appl.: Environment et ressources naturelles, Rabat. Kacimi I. 1992. Hydrogeologie et hydrologie dans le bassin du Gharb: probleme du fer et de manganese. Master's UQAM; pp. 120. Michard, A. 1976. Element de geologie marocaine. Notes et Mkm. Sen. Gkol. Maroc 252.408.
Figure 4. Heavy metal trends in the shallow groundwater of the Sebou Basin.
Figure 5. Heavy metal trends in the shallow groundwater of the Sebou Basin.
(1 1 g/l) fiom C1 (6.6 g/l), Na (2.8 g/l), Ca (1.25 gA), and Mg (0.37 g/l). High concentrations of Sr, Ba, and heavy metals were also observed. 7 CONCLUSIONS The shallow ground waters of the Sebou basin are diffusely employed by the rural population. The chemical quality is in general good enough as regards salinity and heavy metal concentrations, but is often lowered by nitrate pollution fiom agricultural nitrogen fertilisers and by the uncontrolled discharge of waste fiom factories, which causes a strong rise in salinity and toxic heavy metals. ACKNOWLEDGEMENTS This study was supported by funds from the Sardinian Regional Government (RAS).
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Pore waters in Mesozoic rnudrocks in southern England A .H.Bath Intellisci, Loughborough, LE12 6SZ, England
ABSTRACT: Chemical and isotopic data for pore waters in Mesozoic mudrocks from sites in southern England are presented. Pore waters were extracted by squeezing drillcore samples of the mudrocks. Variations of chloride and stable isotopes with depth indicate the extent of dilution of ancient saline water. Salinity changes reflect the different geological histories of the sites and the diffusive movement of solutes and water. 1 INTRODUCTION
2 DESCRIPTIONS OF SITES
Sedimentary rocks, mostly of Jurassic age, were sampled at seven sites in centralhouthern England (Fig. 1) in the 1980s for site investigations and generic research studies connected with disposal of low level radioactive waste. Hydrochemical studies were carried out by the author and other staff at the British Geological Survey (Brightman et al. 1985, Ross et al. 1989, Bath et al. 1989, Metcalfe et al. 1990). This paper describes the compositions of pore waters in the clay-rich sedimentary rocks (for which the term ‘mudrocks’ is used). These reflect the hydrogeological and geochemical evolution of the mudrocks over considerable periods of their postdepositional histories.
The Harwell site (HW: 125 m above sea level) is on the northwestern rim of the London Basin. Jurassic beds underlie 140 m of chalk and sand/clay beds of Cretaceous age and comprise: Kimmeridge Clay (30 m thickness), Corallian sandstones/limestones (35 m), Oxford Clay (110 m) and Oolite hmestones/sandstones. The chalk aquifer is unconfined and deeper saline aquifers occur in the Corallian beds and Oolite formation. Other sites around Hanvell from which drillcores were obtained are at Steventon (EE: 100 m asl), Milton (M: 59 m asl), and Marcham (MF: 68 m asl). These locations are aligned over 10 km northwards from Harwell towards the valley of the River Thames. Cretaceous Upper Greensand and Gault clay were sampled at Steventon (to 45 m depth). Kimmeridge Clay was sampled at Milton (to 54 m) and at Marcham (to 14 m). Basin inversion and erosion took place during the late Tertiary. Incision since then has produced the present arrangement of river valleys. Most of the present land surface has been exposed to meteoric water penetration since the late Tertiary. The Down Ampney site (DA: 78 m asl) is situated 40 km west of Harwell. River terrace gravels (4 m) overlie Jurassic Oxford Clay. Below the Oxford Clay, which varies in thickness from 18-70 m, is the Oolite limestone aquifer. The site contains a major fault which displaces of the base of the Oxford Clay by about 50 m. The fault is about 6 m wide overall and has probably been present since basin subsidence in the Mesozoic. Hydrogeological testing in the faulted zone indicated a slightly higher range of permeabilities.
Figure 1. Outline map showing the locations of sites in southern and central England and abbreviations for each site.
449
of the pore water sample was preserved for chemical and stable isotopic analyses.
The Elstow site (EBB: 35 m asl) is situated 80 km north-east of the Harwell site in the valley of the River Ouse. The Jurassic sequence is about 120 m thick, dipping at low angle to the south-east: 15-30 m of Oxford Clay is underlain by Oolite limestone/sandstone and Lias siltstones and clays. The Oolite limestones form localised aquifers. The Fulbeck site (FB: 17 m asl) is situated 38 km north-east of Nottingham in the Trent river valley. Jurassic strata rest on the Eastern England Shelf, dipping eastwards at c l " towards the North Sea. The Lower Lias sedimentary sequence comprises clays, siltstones and limestones, each unit between 5-30 m thick. Uplift during the Cretaceous resulted in erosion and exposure to meteoric water. Jurassic mudrocks at all sites are overconsolidated stiff clays with hydraulic conductivities typically c10-l' ms-' and moisture contents 15-25%. The conceptual models for groundwater conditions depend on topographic setting of the sites and on flows in minor aquifers formed by limestone/sandstone strata. Groundwater movement, if any, in the mudrocks is upwards or downwards depending on the prevailing hydraulic gradient. At HW and EE the gradients are thought to be downwards across the Gault and Kimmeridge Clays to the Corallian aquifer and upwards across the Oxford Clay from the Oolite aquifer. The nearby M and MF sites are on lower ground and the gradients are likely to be upwards from the regional flow system in the Oolite aquifer. Similar conditions of 'scarp and vale' hydrogeology exist at the EBB site, although the nearby scarp is Oxford Clay so it is uncertain whether there is sufficient flow to sustain an upwards gradient from Oolite limestone through mudrocks in the river valley. The FB site is in a similar topographic setting with a limestone scarp a few kilometres east.
4 HYDROCHEMICAL DATA The results of pore water extraction tests are summarised in Table 1. High sulphate concentrations in many pore water samples originate from oxidation of pyrite during sample storage or squeezing - these correspond with low pH in some but not all cases. Acidity produced by sulphide oxidation was neutralised in situ by reaction with carbonate. Ca and Mg released by carbonate dissolution have exchanged with Na in clays. Sequential sub-samples were taken for HW, EE, M and MF samples as the squeezing load was increased stepwise. Concentrations of K, Na and C1 decreased in successive sub-samples (Brightman et al. 1985). The composition of the first sub-sample of pore water is generally used for geochemical interpretation because it is most representative of 'free' pore water, i.e. is least influenced by mineral surface interactions. All pore waters have Na-C1 or Na-S04-C1 compositions. They are fresh to brackish with the highest chloride concentrations occurring at greatest
Table 1. Summary of mudrock formations and C1, SO4 and pH data for corresponding pore waters.
W IOxClav
3 METHODS
Drillcores samples were preserved and trimmed to remove contaminated material. Pore water was extracted by squeezing in a stainless steel cell to a maximum load of 75-110 MPa (Entwisle & Reeder 1993). Between 10 and 50% of the pore water contents was recovered. Flowing groundwaters were also sampled from interbedded limestones and from the limestone and sandstone strata that occur below the mudrocks at each site. Samples were obtained by pumping from between packered intervals in the boreholes. Extraction of pore waters by squeezing perturbs labile chemical properties, namely pH, alkalinity/pCOz and redox potential. pH and alkalinity were measured immediately after extraction if the sample size was adequate. The rest
1224-3
IM IGault 13 1411-518* 12020-2390* (7.2-8.51 M 134-55 I1540-3170* 196-6900*" 17.0-9.5 bmm I10-456*" 17.5-8.0 MF Kimm 17.0-14 1178-838 5.2-43.3 108-581 8.0-2890 7.6-9.2 DA- OxClay 3100-3390 8.0-9.0 DA- KellBeds 21.5-30.1 223-361 66-126 8.9-9.0 DA- ForMarb 39.3-48.2289-363 8.2-8.7 11.9-63.4700-2600 50-4620 DAD OxClay 7.4-8.7 DAD KellBeds 68.5-78.4 1400-2600 163-4120 1790 206 8.9 DAD ForMarb 87.3 136-7990 6.7-8.6 FB LLFerrLst 1.8-25.2 20-667 FB LLIBuckC132.5-44.6 1120-2500 1200-10300 6.7-9.1 259 9.1 42 FB LL/ObOx 5.25 458-1250 7.4-8.7 3.5-15.4 46-118 EBB OxClay IEBB 1IKellBeds 118.3-23.d 125-605 11550-3670 17.8-8.8 I 1 EBB ULias 140.3-53.l/749-2040 1314-3370 17.7-8.5 Uo-faulted block: Down-faulted block at DA I * Includes range of sequential sub-samples I Sulphate concentrations calculated by charge balance
A
450
I
I
depth at HW: 11,600 mg/l at 318 m depth in the Oxford Clay formation. Stable 180/160 and 2H/'H isotope ratios range from -8.4 to -5%o6l8O and -56 to -34%062H. The lightest values occur at FB at about 30 m depth. The heaviest values are in shallow Kimmeridge Clay at MF and in deep (220-290 m) Oxford Clay at HW. Potassium reaches 160 mg/l in the Oxford Clay at HW and is about 60 mg/l at 54 m in the Kimmeridge Clay at M and at 23 m in the Kellaways Beds at EBB. Magnesium and calcium are generally correlated. The Br/Cl ratio is 0.005 by weight (cf. 0.0035 in sea water), suggesting that Br has increased possibly due to diagenesis of organic matter (there are no Br, F, total organic carbon or thiosulphate data for HW, EE, M and MF samples). Total organic carbon concentrations are mostly in the range 15-40 mg/l, with a maximum of 135 mg/l in the Lower Lias at FB. The rather narrow range of TOC values suggests that there is a limiting process, perhaps the rate of degradation or solubility of solid organics. In common with other cases where rapid sul hide oxidation has occurred, thiosulphate ( S 2 0 3 ?) is present in significant amounts: several hundreds mg/l in DA and FB pore waters and up to 1885 mg/l in one sample.
5 INTERPRETATION AND DISCUSSION 5.1 Salinities: dilution of ancient saline water The change of chloride with depth in the mudrock sequences (Fig. 2) indicates that ancient saline water, originating from marine conditions at and after deposition, is being diluted. Surface infiltration (in the case of outcropping mudrocks) or adjacent aquifer formations are sources of diluting water. Stable isotope compositions are further evidence of mixing between waters originating from different sources (Fig. 3). Geographic variation of meteoric
Figure 2. Chloride concentrations versus depth in pore waters from Jurassic mudrocks at various sites.
Figure 3. Stable oxygen isotope compositions versus depth in pore waters from Jurassic mudrocks.
water compositions accounts for some of the spread of 6'*0 values for shallow pore waters. Lower 6"O values at FB, DA and HW are characteristic of colder climate recharge. Higher 6l8O values in the Oxford Clay at HW are closer to the marine composition that prevailed for much of the Mesozoic burial history. It is likely that dilution of ancient marine pore waters has occurred by diffusion of solutes and water between mudrocks and adjacent aquifers. Hydraulic gradients generally would have been insufficient to drive significant flow through intact mudrocks though advection might occur through fractures and interbedded limestones. The role of diffusion can be illustrated by comparison of Figure 2 with the transient concentration profiles that are calculated with the diffusion equation for progressive dilution of marine pore waters by an overlying fresh water aquifer (Fig. 4). The profile in Figure 2 corresponds to a diffusion time of >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.
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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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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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.
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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
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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 <0,05 <0,02 < 0,05 30 - 50 B 0,05 - 0,2 0,02 - 0,2 0,05 - 0,5 > 50 C >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
463
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.).
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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 (
465
Cd, Pb, Se, Bi, Te, and Sb were not included in the process because their concentrations were below the detection limit in most samples. The frequency distributions of concentration values are non-normal, most of which displaying log-normal characteristics. Multivariate statistic analyses (cluster and principal component) were carried out on the normalised data. Hierarchical cluster analysis was performed using Squared Euclidean distances and Ward’s method as a similarity criterion and Principal Component Analysis (PCA) by Varimax rotation. Geostatistical study was carried out by linear Kriging and contouring of the factor scores. 4 RESULTS AND DISCUSSION 4.1 Hydrogeochemical classlJication
Figure 1. Geological schematic map of the Flumendosa river basin with location of the water samples. 1: Quaternary deposits; 2: Pliocene-Pleistocene volcanic rocks; 3: Tertiary sediments; 4: Jurassic limestone-dolomite formation; 5: Permian-Carboniferous volcanic complex; 6: Hercynian granites; 7: Middle Ordovician volcano-sedimentary sequence; 8: Paleozoic metasediments (Cambrian to Devonian); 9: water samples (Barca et al. 1996, modified).
The water samples were filtered at 0.4 pm pore-size and collected in pre-cleaned bottles. One aliquot was acidified to 1% HN03 for the analysis of most metals, and another to 0.2% HCl for As and Sb. Unfiltered samples were acidified to 0.2% H2SO4 (with addition of a KMn04 grain) for Hg. Anions were determined by ion chromatography on unacidified aliquots. Metals were determined by ICP-OES and ICP-MS. Mercury, antimony, and arsenic were determined by ICP-MS after Hg-vapour or hydride generation respectively. Cr was determined by GFAAS; NH, by a spectrophotometry. 3.2 Statistical data processing Geostatistic methods have been widely used in geochemical data (Marini & Ottonello 1997a, b; Ceron et al. 2000, and references therein). Statistic data processing was performed on 30 out of 41 analysed elements. NI&, NO2, Fe, Cr, Al,
The water samples show total dissolved solids (TDS) ranging from 62 to 896 mg/l; only the waters in two wells close to the coast have TDS higher than 1 g/l; pH values are in the 5.9 - 8.0 range, and dissolved O2ranging from 1.4 to 8 mg/l. Hierarchical cluster analyses allowed to separate water samples into three groups. The chemical composition of waters in each group is ascribable to the different zones of the hydrological basin. The first group (av. TDS = 128 & 39 mg/l) refers to samples from springs at higher altitude in the Paleozoic schistose and granitic formations, prevalently located in the northern sector of the basin. The second group (av. TDS = 340 k 106 mg/l) is related to waters in the Paleozoic schistose formations and in the Tertiary sediments at lower altitude in the basin, generally in the fractured southern area. The third group (av. TDS = 407 k 65 mg/l) is composed of waters interacting with Jurassic limestonedolomite formations in the central part of the basin. The chemical composition of each group is shown in the triangular plots of Figure 2. These plots show that chemical molar ratios of the first two groups are similar; the parameter that best distinguish the waters is the TDS. In fact, very dilute waters of the first group present generally a HC03-Na(Ca) composition, while the waters of the second group, with higher TDS values, have a bicarbonatecalcic-sodic character showing a slightly more evolute composition. Waters of third group are of Ca-Mg-HC03 type with a Ca/Mg molar ratio approaching unity; their chemical composition is typical of waters interacting with dolomites. Cluster analysis rejected three more saline samples. Two samples, from wells in the coastal alluvial plain of Muravera, show a dominant sodium chloride chemical composition due to marine intrusion in a overexploited aquifer, another sample with a calcium sulphate character derives from the oxida tion of sulphide minerals. 466
Figure 3. Plot of HC03 vs pH for the waters of the Flumendosa river basin (symbols as in Figure 2). The two solid lines refer to calcite saturation at 0 "C and 25 "C.
undersaturated with respect to carbonate mineral phases, is ascribable to the initial stages of waterrock interaction showing a rain similar chemistry. The second group of waters with HC03-Ca-Na composition shows calcium concentration typical of water at a more advanced interaction, and these waters reach equilibrium with respect to calcite. 4.2 Hydrogeochemical maps
Figure 2. Ternary plots of cationic and anionic molar ratios. Open and solid circles represent respectively dilute (1' group) and more evolved (2ndgroup) waters from the basement; solid rectangle represents water from Jurassic carbonatic formation (3rd group). Stars and cross show respectively anomalous samples for marine intrusion contamination and for sulphide minerals oxidation.
Speciation with the PHREEQC computer program (Parkhurst 1995), shows that the less evolved waters of the first group are undersaturated with respect to all mineral phases, except chalcedony, while slightly more evolved waters of the second group approach equilibrium with respect to calcite and dolomite when TDS and pH increase. The samples with a TDS higher than 300 mg/l are at equilibrium with these two mineral phases. The waters of the limestone-dolomite formations (3rd group) are close to equilibrium with respect to chalcedony and calcite and slightly oversaturated to dolomite. Equilibrium with barite is reached due to the increase in the barium andor sulphate components in the more saline samples. The plot in Figure 3 reveals an evolutive trend from dilute to more saline water on increasing pH. The first group with HC03-Na-(Ca) composition,
The factors extracted by PCA are representative of the hydrogeochemical properties. Although PCA yields six factors that account for 79% of the total variance, only the first most representative three principal components will be here discussed. The loading by variable on the first three principal components (factors) are given in Table 1 and the maps by factor score contouring of each factors are showed in Figure 4. Factor 1 associates elements deriving by interaction with Paleozoic and Tertiary sediments. In fact, the most positive factor scores are associated with the waters hosted in the southern area of the basin, where Tertiary sediments and most fractured Paleozoic rocks are located (Fig. 4a). Factor 2 associates elements from the interaction with carbonate rocks, and presents the highest factor Table 1. Factor loadings of the three factors that explain most of variability in the principal component analysis. Percentage of variance explained ("AV) by each factor is also reported. FACTOR 1 Ba 0.81 K 0.80 Na 0.79 Ga 0.78 Br 0.77 V 0.73 Rb 0.69 C1 0.68 SiOz 0.64 Sr 0.61 B 0.61 %V 40
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FACTOR 2 HC03 0.94 Ca 0.92 U 0.84 Mg 0.81
54
FACTOR 3 Zn 0.80 CO 0.71 Mn 0.71 Ni 0.66 Cu 0.61
62
Figure 4a, b, c. Factor scores contouring maps for Factor l(a), 2 (b) and 3 (c) of the PCA. The darker and lighter colours in the grey scale correspond to higher and lower values, and represent the more or less intense processes of water rock interaction, respectively. Factor 1 groups chemical elements mainly fiom the leaching of silicatic rocks; higher values are located in the area of the lower course of the Flumendosa, where Tertiary sediments and most fiactured Paleozoic rocks are located. Factor 2 groups elements fiom the interaction with carbonate rocks and presents the highest incidence in the central part of the basin where Jurassic limestonedolomite formations are located. Factor 3 groups Zn, CO, Mn, Ni, and Cu, and represents the leaching of mineralised rocks. Higher values were found near the most important mineralised areas: copper minerals at Gadoni, Zn, Ag, and As sulphides at Mt.Narba, Pb sulphide at M.Gibas, and Fe-oxides at Perdasdefogu.
scores in the central part of the basin, where Jurassic limestone-dolomite formations are located (Fig. 4b). Factor 3 groups the typical toxic metals of the sulphide mineralisations of the sector. Figure 4c shows how the anomalies related to the interaction processes of the waters with the known mineralisation of the area (Gadoni, Mt. Narba, M. Gibas, Escalaplano, and Silius) are effectively highlighted.
ACKNOWLEDGWNTS This research was supported by the financial contribution of MLTRST, Italian Minister for University and Scientific Technologic Research.
REFERENCES Ardau, F. & G. Barbieri 1994. Evolution of salt water intrusion phenomena in the coastal plain of Muravera (South-Eastem Sardinia). Proceedings 13thS m G. Barrom (ed): 305-312. Barca, S., Carmignani, L., Oggiano, G., Pertusati, P.C. & I. Salvadori 1996. Carta Geologica della Sardegna, Scala 1:200,000. L.A.C., Firenze. Bertorino, G., Caboi, R., Caredda, A.M., Fanfani L., Gradoli, M.G. & P. Zuddas 1987. Prospezione idrogeochimica mineraria nell'area di Gadoni - Seulo (Sardegna Centrale): il significato di solfati e fluoruri disciolti. Rend Soc. It. di Min. e Petrol., 42: 47-58. Ceron, J.C., Jimenez-Espinosa, R. & A. Pulido-Bosch 2000. Numerical analysis of hydrogeochemical data: a case study (Alto Guadalentin, southeast Spain). Applied Geochemistry 15: 1053-1067. Manq L. & G. Ottonello 1997a. Atlante degh acquiferi di Genova. 1: Alta Val Bisagno ed Alta Val Polwera Pia:P d . Manq L. & G. Ottonello 1997b. Atlante degh acquiferidi Genova. 2: La falda di subalveo della bassa Val Bisagno. Pisa: Pacini. Parkhurst, D. 1995. A Computer program for speciation, reaction-path, advective-transport, and inverse geochemical calculations-User's guide. US.G.S. Water resources Investigation report 95-4227: 151 p.
5 CONCLUSIONS The main chemical characteristics of the shallow waters in the Flumendosa river basin point out a generally good quality. In fact, most waters are relatively low in TDS with a prevalence of alkalineearthy bicarbonates in the dissolved salts. At times the chemical quality of the waters is lowered by the presence of toxic metals from the leaching of mineralised rocks. The water resources available are scanty but very important for human activity (rural and touristrecreational) because they are widespread in the territory and of good quality. The representation method of the scores relating to the factors extracted by PCA appears adequate to prepare hydrogeochemical maps. 468
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrogeochernical characteristics of surface water and groundwater in areas underlain by black shales and slates of the Okchon zone, Korea Hyo-Taek Chon & Seok Young Oh School of Civil,Urban and Geosystem Engineering, College of Engineering, Seoul National University, Seoul 151-742, Korea
ABSTRACT: The purpose of this study is to investigate the hydrogeochemical characteristics and the potential contamination of surface water and groundwater in areas underlain by black shales and other metasedimentary rocks in the Okchon zone, Korea. Surface water and groundwater samples were collected from the Duk-Pyung, Chu-Bu, Bo-Eun, I-Won and Geum-Kwan areas of the Okchon zone, and the in-situ measurement and chemical analysis of dissolved constituents were performed. All water samples have Ca2’-HC0,and Ca”-(Cl--SO%-)compositions due to the dissolution of calcite and Mg-carbonates and the oxidative dissolution of sulfide minerals from black shale and slates. The increase of concentrations of Ca”, Mg”, HCO;, Na’ and SiO, from surface water to groundwater indicates progressive water-rock interaction as supported by hydrogeochemical modeling and phase stabilities of related minerals. High concentrations of K+, C1- and NO,are also found in groundwater due to the inflow of external contaminants. From the results of statistical analysis, it could be concluded that the factors controlling the hydrogeochemical characteristics of surface water and groundwater are the inflow of external contaminants, the dissolution of carbonates, the weathering of silicates and the oxidation of sulfides. tamination of natural water, and (3) to interpret the obtained results in comparison with the previous results of soil contamination in the Okchon zone underlain by black shales and metasedimentary rocks.
1 INTRODUCTION The Okchon black shale in Korea provides a typical example of natural geological materials enriched in potentially toxic elements such as As, Cd, Cu, MO, Ni, Pb, U, V, and Zn (Kim & Thornton 1993, Lee et al. 1998). This black shale formation is a member of the so called Guryongsan Formation or Changri Formation of Cambro-Ordovician age, which are part of the Okchon group outcropping in the central part of the southern Korean Peninsula (Lee et al. 1998). In 1990’s, many researchers have reported high levels of trace elements in residual soils derived from black shales and slates of the Okchon zone in Korea (Kim & Thornton 1993, Chon et al. 1996). Dispersion patterns and chemical forms of toxic elements in the Okchon black shale areas were also investigated (Lee et al. 1998). Recently, an hydrogeochemical study on the quality and contamination of groundwater around the Duk-Pyung area was also performed (Kim et al. 1998). In this study, five areas of the Okchon zone (the Duk-Pyung, Bo-Eun, Chu-Bu, Geum-Kwan and Iwon areas) were selected in order (1) to investigate the hydrogeochemical characteristicsof groundwater and surface water, (2) to identi@ the potential con-
2 GEOLOGY The Okchon zone, located in the central part of Korea, has an average width of 80 km and a northeastern trend (Fig. 1). Black shales in the Okchon zone occur in the Guryongsan formation, which is also named the Changri Formation in the southern part of the Okchon zone (Lee & Kim 1972). In the five study areas, these formations are mainly composed of metasedimentary rocks such as black shale, slate, phyllite, limestone and chlorite schist (Lee et al. 1998). Black shales are mainly found in the DukPyung, Bo-Eun and Chu-Bu areas and black slates in the Geum-Kwan and I-Won areas. According to the previous mineralogical study (Lee & Kim 1972), black shale and slate in the Guryungsan formation consist of quartz, feldspar, pyrite, coal, chlorite and plagioclase with sericite. Black slate in the Changri formation is composed of quartz, muscovite, biotite, chlorite, plagioclase and graphite as primary miner-
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that surface water samples experienced various pHchanging interactions with geological media, such as dissolution, hydrolysis and precipitation. One surface waters in the I-Won area collected around a black slate quarry shows a strongly acidic pH value of 4.1 and 130.9 mg/L of SO,:' which can be attributed to the oxidation of sulfide minerals. Most groundwater samples show a range of pH values (6.0-8.7) similar to that of surface water samples, which indicates that similar interactions occur in local aquifers. Surface waters and groundwaters with TDS contents of 53-445 m g L are classified as fresh waters (Fetter 1994). However, a significant difference in average TDS is observed between surface waters (114 mg/L) and groundwaters (194 mg/L). Most surface waters and groundwaters have Ca"HCO,- or Ca"-(Cl--SO,"-) compositions. All Ca2'HC0,- groundwaters originate through dissolution of calcite, whereas Ca2+-(Cl--S0,2-)groundwaters and surface waters are generated by oxidation of sulfide minerals. 4.2 Behavior of dissolved constituents by depth
Figure 1. Locations of the 5 study areas situated in the Okchon zone of Korea.
Groundwaters (54 points) and surface waters (34 points) were sampled along the stream line at five study areas. All samples were collected in a polyethylene bottle (1,000 mL) and transferred to a laboratory in a cold storage. Then, the samples for ion analysis (200 mL) were filtered through a 0.45 pm cellulose nitrate membrane filter using a hand vacuum pump. Samples for cation analysis (100 mL) were acidified with HNO, to pH<2. All samples were preserved in refrigerator before analysis. Eh, pH, electrical conductivity (EC) and temperature were measured in-situ. Concentrations of Al, B, Bay Ca, Cd, CryCu, Fe, Mg, Mn, Pb, Sr and Si, Zn were determined by ICP-MS, K, Li and Na by AAS, F, C1, Br, NO,, NO,, PO, and SO, by ICYand alkalinity by titration. From the balance of cationic and anionic charges, it was found that overall analytical errors were less than 1% for all samples.
Groundwater samples were subdivided by well depth as shallow groundwater (loo m) to verify the behavior of dissolved constituents by depth. Average concentrations of Ca2+(37.42 mg/L), Mg2' (5.41 mg/L) and HCO,- (93 mg/L) in shallow groundwaters increase abruptly in comparison with those of surface waters and deep groundwaters. The dissolution of carbonate minerals evidently affects the increase of such ions in shallow aquifers. Average concentrations of SiO, (21.33 mg/L) and Na' (9.19 mg/L) in deep groundwaters are much higher than those of surface waters and shallow groundwaters, indicating that the dissolution of silicate minerals such as feldspar and mica can occurs in deep aquifers. However, average K' concentration of shallow groundwaters (2.94 mgL) is higher than that of surface waters and is less than that of deep groundwaters. Average concentrations of NO,- (25.3 mg/L) and C1- (13.9 mgL) in shallow groundwaters is much higher than those of surface waters, due to pollution related with agricultural and anthropogenic activities. Average F- concentration (1 .OS m a ) in deep groundwaters is higher than that of surface and shallow groundwaters, due to dissolution of silicate minerals.
4 RESULTS AND DISCUSSION
4.3 Statistical Analysis
als. Pyrite, microcline, tourmaline, hematite and limonite are also found in black slates. 3 MATERIALS AND METHODS
4.1 Hydrogeochemical characteristics Most surface water samples show a pH range from slightly acidic to alkaline (6.4-8.8). This indicates
Factor analysis, a kind of multivariate data analysis, was used to perform the statistical interpretation of water-rock interactions in areas underlain black shales and slate of the Okchon zone. R-mode factor 470
analysis was performed with SPSS/PC+ to reduce the complexity of analytical data and group the numerous chemical variables in a limited number of factors. Major dissolved constituents were treated as input variables. Factors with eigenvalues greater than 1 were retained. Results of factor analysis show that the factors controlling the chemical compositions of surface waters and groundwaters reflect the following associations (Fl) Cl--K+-NO,-,(F2) HC0,-Ca2'-SI2'-Mg2+, (F3) F--Na+-SiO, and (F4) SO,"(Table 1). The first factor is explained by inflow of pollutants in waters. The second factor reflects dissolution of carbonate minerals. The weathering of silicate minerals to clays governs the third factor. The fourth factor is related to oxidation of sulfide minerals. 4.4 HydrogeochemicalModeling Hydrogeochemical modeling by WATEQ4F (Ball & Nordstrom 1991) was performed to obtain the saturation indices of specified minerals which might enlighten the interactions of surface water and groundwater with rocks. Most water samples are undersaturated with respect to calcite and dolomite, which indicates the possible dissolution of these carbonate Some groundwater however, are oversaturated with calcite, indicating that precipitation of calcite might locally occur in aquifers. Most water samples are undersaturated with respect to albite, anorthite, and tremolite, indicating possible dissolution of these silicate minerals. In order to verify the possibility of dissolution of silicate minerals and identify the weathering products, ion activities calculated by WATEQ4F were plotted on activity diagrams (Figs. 2 and 3). Most
Figure 2. Activity diagram for the system Na,O-Al,O,-Si0,H,O at 298K and 1 atm. All thermodynamic data are from Helgeson (1969).
surface water and groundwater samples are plotted in the field of kaolinite or in the field of kaolinite silicate minerals, and illite, suggesting that
Table 1. Varimax factor matrix of chemical constituents and factor scores for waters collected in the study areas. Variable
Factor 1
c1 K NO, HCO, ca Sr Mg F Na Si02
2
Percentage* 3
0.868 0.805 0.789
0.463 0.595
4
45.3 0.950 0.877 0.673 0.530
15.2
0.901 0.662 0.632
10.5
so4 0.939 8.8 Eigenvalue 4.88 1.88 1.15 1.08 * Percentageof variance explained by facor ('Yo), F: Factor ** Factor loadings less than 0.4 are omitted.
Figure 3. Activity dlagram for the system K20-Al,03-Si02-H20 at 298K and 1 atm. All thermodynamic data are from Helgeson (1969), but that of muscovite from Nesbitt (1977).
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such as plagioclase, microcline and mica, can be dissolved and weathered to kaolinite and illite. This is also in agreement with previous result of XRD analyses on residual soils in the Duk-Pyung area (Chon et al. 1997).
K I ~ H.D., , Woo, N.C. & M.J. Choi 1998. Groundwater quality and contamination in Dukpyung area. J. of the Korean Society of Groundwater Environment 5: 141-148. Kim, K.W. & I. Thornton 1993. Influence of Ordovician uramferous black shales on trace element composition of soils and food crops in Korea. Applied Geochemistry, Suppl. Issue no.2, 249-255. Lee, C.H. & J.H. Kim 1972. Explanatory Text of the Geological Map of Goesan. Geological Society of Korea. Lee, J.S., Chon, H.T. & K.W. Kim 1998. Migration and Qspersion of trace elements in the rock-soil-plant system in areas underlain by black shales and slates of the Okchon zone, Korea. J. Geochemical Exploration 65: 61-78. Nesbitt H.W. 1977. Estimation of the thermodymtuc properties of Na-, Ca-and Mg-beidellites. Canadian Mineralogist 15: 22-30.
5 CONCLUSIONS Main conclusions are summarized as follows: (1) Surface waters and groundwaters collected in the Okchon zone have Ca2'-HC0,- or Ca2'-(Cl-SO:-) composition, due to either dissolution of carbonate minerals or oxidation of sulfide minerals in black shales and slates. (2) Shallow groundwaters ( 4 0 0 m depth) have high concentrations of Ca", Mg2' and HCO,- due to dissolution of carbonate minerals. On the contrary, deep groundwaters (>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-
474
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
475
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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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
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 . $, .g 5 G
2.47pj 3.11,Jj 1.57pj 3.17iIj lid 1.54(,, na 1.37,lj 3.45Clj nd 1.76(,, na 1.43,,, 3.15,,, lid I1 1.23(,, 4.63,,, na 2.45(1, nd 4.12,,, na iia 1.98(1! lid I11 lid 4.87,,, nd 2.33(1j lid IV 1.12,2, 2.24,,, nd 2.45(,, 1.67(,, nd 2.84(1, nd 3.1 O.68,,, V lid 2.14,;j na 2.440j 1.39(lj na 5.25(3, 1.23(,, 2.32,,, 1.44il, I. 7-15 3-8 <5 nd modal 11. 25-32 3-5 <5 nd % of 111. 17-25 <3 <5 nd <3 <5 <3 minerals IV. 5- 15 --+ V 5-20 <3 <5 <3 *see Figure 1 for sample aiid locality; ( pniber oftarget minerals; lid, not determined; na, iiot analysed;
1 2 4 9 10 14 16 18 19 20
2.2
P
-cd
I
64
IIC
4,2 a u
IIC IIC
-$ cdg
I1
‘u
4-J .-
ea
111
E
up
2 h
.z 2 ‘2
t-
IV
-
U
3 2 Sprrllg-lwler5
V
IIC 0 ?% k 6
nc gz *see Figure 1 for sample aiid locality: **fromCoban ct a1 2000; IIC, iiot calculated; P, potassic; UP, ultrapotassic. 22
0.10
I
5.45 0.22
hasized that the F anion has nearly the same ionic radius (1.33AO) as that of the (OH) anion. Thus, minor amounts of F are stored in (OH) sites of hydrous minerals at shallow levels of the upper mantle. Furthermore, micas are very efficient at removing flourine from a melt (Munoz & Eugster 1969). The high F content of phlogopite in ultrapotassics from some localities of the world are as follows: 3 100-4806 ppm, West Kimberley, Australia (Jaques et al. 1986), 2380-4520 ppm, Leucite Hills (Aoki et al. 198l), up to 7500 ppm, Alban Hills, Central Italy (Gaeta et al. 2000). Although F- values in apatite, amphibole and titanite in our samples are noticeable, and there has been a positive correlation between F and PzOS,they provide lower amounts of flourine because of its less abundance in the rocks. Moreover, apatite alone can not supply all of F, and the greater part of F in basaltic magma, especially high K ones, is supplied by breakdown of minerals other than apatite (Aoki et al. 1981). The recent study supports that the other mentioned minerals are phlogopite, biotite, amphibole and titanite. Among these, the most eficient minerals are micas.
To determine the rock-water interaction, spring-water samples discharging from the volcanics were collected periodically i n different seasons (Table 3) F- values ranged between 0 31 - 3 35 mglL in the Isparta springs, 0 1 - 1 5 mg/L in $hut (S of Afyon) springs The data show two interesting points ’l’heseare 1;changes in different seasons and regions, and Fcontent in the springs decreases i n the rainy seasons while it increases in dry seasons. The reason of less F content in rainy seasons is thought to be a short interaction time between water and rock. Table 3. F contents (mg/L) of springs related to the volcanics in 1999.
IGolcuklake, 1 0.79 0.93 1.22 1.95 1.32 1.24 I1 Andik** 2 I1 Isparta** 3 IV Cakirozu***4 IV Aydin*** 5 IV Kavakli*** 6 IV Yiprak*** 7
2.80 3.12 3.35 2.61 2.92 0.31 0.46 1.12 0.26 0.49 0.42 0.64 0.71 0.46 0.53 0.50 0.91 1.13 0.48 0.68 0.83 1.38 1.56 0.85 1.07 0.14 0.24 0.22 0.10 0.16 *see Figurc 1 for localit).:** stream names. ***spring names 479
2.75 0.33 0.43 0.38 0.75 0.1 I
F-contents of spring-waters related to potassic and pyroclastic rocks in the Isparta region are higher than those of springs from ultrapotassic and pyroclastic rocks in south of Afyon. However, the rock analysis show that ultrapotassics contain higher F than potassics (Table 1). So, springs from the south of Afyon are expected to have relatively higher F. We consider two reasons for this contradiction; the volume and outcrops of volcanic lavas and the rock type influencing the water. The lavas of potassic volcanics from Isparta cover wider areas than Afyon ultrapotassics that are present in small volumes as dykes and cones. Additionaly, the rock type influencing water is dominantly lavas in the Isparta region while it is pyroclastics in the south of Afyon.
4 CONCLUSIONS Fluorine data signify that high fluorine is a prominent aspect of Afyon-Isparta volcanic district in SW Turkey. It is more abundant in ultrapotassic rocks than potassics, and pyroclastics contain the lowest amounts of fluorine. Mica (phlogopitehiotite), apatite, amphibole and titanite were found as the main F carrying minerals in this volcanic realm. The flourmicas (predominantly fluorphlogopite) were determined the main F source minerals. So, high F is attributed to abundance of micas. F anomalies were also determined in springs discharging from volcanics. According to obtained data, the high F of groundwater was assumed to have been influenced by volcanism, particularly ultrapotassic/potassic type rocks which include fluormicas.
REFERENCES Aoki, K. & S. Kanisawa 1979. Fluorine contents of some hydrous minerals derived from upper mantle and lower crust. Lithos 12: 167-171 Aoki, K., Ishikawa, K. & S. Kanisawa 1981. Fluorine geochemistry of basaltic rocks from continental and oceanic regions and petrogenetic application. Contrib. Mineral. Petrol. 16: 53-59 Ayhan, E. 1983. Sulardaki Fluorun muhtemel kokeni, A.U., I. Miih. Hft.Smp.Turkey, 15: 49-65 Besang, C., Eckhardt, F.J., Harre, W., Kreuzer, H. & P. Muller 1977. Radiometris. Alterbestim. and Neogen. Erupt. der Tiirkei. Geol. Jb. 25: 3-36 Bilgin, A., Koseoglu, M. & G. Ozkan 1990. Isparta-Golciik yoresi kayaqlarinin mineraloji, petrografi ve Jeokimyasi. Doga, 14: 342-361, Ankara. Coban, H., BOZCU,M. & K. Yilmaz 2000. Petrogenesis of ultrapotassic rocks from Afyon-Isparta volcanic suture(SW Turkey), Abs. Internal. Earth Sc. Congress on the Eagen Regions, IESCA, Izmir/Turkey.
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Foley, S.F., Taylor, W.R. & D.H. Green 1986. The role of fluorine and oxygen fkgacity in the genesis of the ultrapotassic rocks. Contrib. Mineral. Petrol. 94: 183- 192 Gaeta, M., Fabrizio, G. & G. Cavaretta 2000. F-phlogopites in the Alban Hills volcanic district (Central Italy) indications regarding the role of volatiles in magmatic crystallization, , J. Volcan. Geotherm. Res. 99: 179-193 Jaques, A.L., Lewis, J.D. & C.B. Smith 1986. The kimberlites and lamproites of Western Australia. Geol. Sun. Western Aust. Bull. 132. 268 pp. Kazanci, N. 1995. Egirdir go1 canaginin olu$um zamanina ili3kin gozlem, Tiirkiye Jeoloji Miihendisligi Dergisi 47:3233, Ankara. Keller, J. 1983. Potassic lavas in the orogenic volcanism of the Mediterranean area. Volcanol. Geotherm. Res. 18: 321-335. Koqyigit, A. 1984. Intraplate neotectonic development in SW Turkey and adjacent areas. Bull. Geol. Soc. Turkey 21: 1-16. Kuehner, SM, Edgar, A.D. & M. Arima 1981. Petrogenesis of the ultrapotassic rocks from the Leucite Hills, Wyoming. Am. Min.. 661663-671 Lefevre, C., Bellon, H. & A. Poisson 1983. Presence de leucitites dans 1e volcanism Pliocene de la region d’Isparta (Taurides occidentales, Turquie). C.R.Acad. Sc. Paris 297: 367-372. Munoz, J.L. & H.P. Eugster 1969. Experimental control of fluorine reactions in hydrothermal systems. Am. Mineral. 54: 943-959. Ozgur, N., Pekdeger, A., Schneider, H.J. & A. Bilgin 1990. Pliocene Volcanism in the Golciik area, IspartdW-Taurides. Proc. Internal. Earth Sc. Congress on the Aegean Region 41 1-419. Pekdeger, A., Ozgur, 0. & H.J. Schneider 1992. Hydrogeochemistry of fluorine in shallow aqueous systems of the Golcuk area, SW Turkey. Pr0c.7‘~Int..Sym. On Water-Rock Interaction, Utah, U.S.A., S. 821-824 TS-266 1986. Drinking waters, TSE (Institute of Turkish Standart), Ankara. Yagmurlu, F., Savasqin, M.Y. & M. Ergun 1997. Relation of alkaline volcanism and active tectonism within the evolution of Isparta Angle, SW Turkey. J. Geol. 105, 7 17-728
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Groundwater composition of perched-water bodies at Azores volcanic islands J.V.Cruz & Z.M.Franga Department of Geosciences, University of the Azores, Portugal
ABSTRACT: A data set of the groundwater composition from perched-water bodies of the Azores archipelago was compiled, in order to characterize water chemistry and to assess differences due to lithologic controls. About 253 analyses from the nine islands of the Azores are studied and springs are divided for this study as discharges from aquifers of basaltic nature and discharges from more acidic volcanics as trachytes. Waters are cold, with a pH range from 4.70 to 8.21 and electrical conductivity measurements range from 36 pS/cm to 725 pS/cm. Groundwater chemistry is mainly expressed in equivalent units as Na'>Mg2+>Ca2+>K+for cations and Cl'>HC03->SO?' for anions. Two majors trends can be defined: (1) waters that evolve mainly by CO2 solution and water-rock interaction and (2) waters composition modified by sea-salts spraying, which chemistry is strongly controlled by chloride and sodium. 1 INTRODUCTION The Azores is an archipelago, made of nine islands of volcanic origin, scattered along a 500 Km length strip trending NW-SE. Located in the North Atlantic Ocean, near the junction between the American, Euroasian and Afkican plates, the Azores archipelago is characterized by an intense volcanic and seismological activity, well marked by over 30 inland eruptions since the settlement in the XV century. Groundwater is a key natural resource at the Azores archipelago, where about 98% of the total domestic water supply comes from springs and wells exploitation (Cruz & Coutinho 1998). In the majority of the islands the supply is entirely of groundwater origin. Due to this framework knowledge about groundwater chemistry is rather important, supporting the water resources management, through the identification of the major trends as well as the contamination problems arising from the increasing agriculture activities and the overexploitation, that lead in a few wells toward salinization problems. The hydrogeologic conceptual models that have been described in the Azores fit in the general literature trend, and can be grouped into two main types: (1) the Hawaiian model, considering a low-lying basal water body linked to inland dike-impounded and perched-water bodies (Peterson 1972, 1993; Macdonald et a1 1983); and (2) the Canary Islands model, that considers a continuous basal water body (Custodio 1975, 1978). The springs studied in the present paper corresponds to the discharge of the perched-water bodies,
in the islands slopes. Discharge is higher in springs related with lava flows aquifers comparing to aquifers made of pyroclastic deposits. For the major island of the archipelago (S5o Miguel) a study show that the avera e recession discharge in lava flows is about 484 m /day, higher than the 33 m3/day for springs in pyroclastic layers (Paradela 1980). From the hydrogeological framework groundwater short residence times are expected, which implies high water-rock ratios (Langmuir 1997).
F
2 DATA PRESENTATION 2.1 Sources A vast hydrogeochemical data base was compiled from the literature, made off 253 cold spring waters composition from all the nine islands from the archipelago. Therefore, this paper corresponds to a preliminary overall study on groundwater composition at the Azores, on a comparison basis, allowing to stress the main characteristicsand processes. The literature survey was made in order to observe homogeneity of methods, relying whenever possible in one reference per island. The data is from Santa Maria island (Cruz 1992, Cruz et al. 1992), Silo Miguel (Coutinho 1990, Lob0 1993, Cruz et al. 1999, Carvalho 1999), Terceira (Lobo 1993), Siio Jorge (Lobo 1993), Graciosa (Lobo 1993), Pico (Cruz 1997, Cruz et al. 2000), Faial (Coutinho 2000) and Flores and Corvo (Lobo 1993). The analytical methods carried out by the several authors were atomic-absorption and ionic chro-
48 1
mathography. Sampled waters were kept in doublecapped polyethylene bottles and pH, temperature and electrical conductivity were measured in the field immediately after sampling, as well as dissolved CO2 and alkalinity. For the present study, after the original data set completion, all the analysis with a charge-balance error higher than 10% were rejected. The selected waters are cold, with an average temperature of 15°C. They are acid to slightly alkaline, as the pH ranges from about 4.7 to 8.21, and presents a median equal to 7.20. Alkalinity ranges from 4 to 170 mg/L of CaC03. Electrical conductivity ranged in value from 36 to 725 pS/cm, with an average equal to 188 pS/cm, which suggest that the springs related with the perched-water bodies corresponds to low mineralised waters. The waters can be classified as soft to very hard, as total hardness varies between 2 and 155.7 mg/l CaCO3 (Freeze & Cherry 1979), but the average value (33.7 mg/l CaC03) is on the soft waters range.
Figure 2. Relationship between water electrical conductivity and chloride.
Groundwater composition is dominated by the major ionic species in solution, bicarbonate, chloride, potassium and sodium. In the Figure 1 the maximum, the minimum and the median concentration is plotted, in order to show the actual range of the content of each specie in solution. Chloride and sodium can account respectively for 13.8% to 84.9% of the total anions in solution and 38% to 93.2% of the total cations. The water mineralisation control by these species can be shown by the close positive correlation between chloride and the water electrical conductivity (Fig. 2). Chloride and sodium show a well-defined linear relation (Fig. 3), which is thought to be caused by the influence of seawater spraying. Some springs discharge waters with a bicarbonate enrichment, and this specie can account for as much as 84.9% of the total anions content in solution. Therefore, bicarbonate also controls partially the wa-
ter mineralisation, as can be shown by the positive linear correlation in the plot with electrical conductivity (Fig. 4). The bicarbonate enrichment is interpreted as a result of the reaction that involves carbon dioxide dissolved in soil-water and groundwater and the silicate minerals. Nevertheless, it should be mention that in active volcanic centers, as occurs in the majority of the islands from Azores, the partial pressure of CO2 can be increased due to volatile release which also contributes to the observed bicarbonate enrichment, as observed for example at Furnas volcano (Cruz et al. 1999). The groundwater chemistry of these springs with higher bicarbonate content can be expressed in equivalent units as Na+>Mg2+>Ca2+>K+ for cations and HC03'>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.
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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.
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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<%mol MgC03 <0.05 do not produce significant changes (log Kss=-8.49k0.005). For this reason inverse computing of CaMg ratio in solid from Ca/Mg ratio in aqueous solution is possible only for Mg-rich calcite. Further detailed mineralogical investigations of MdCa ratios on carbonate rocks forming Hyblean aquifer are required to better un-
Figure 3. Stability diagram phase of Kaolinite, Albite, Gibbsite and Na-Montmorillonite in aqueous solution at 25°C. Close circles represent VD groundwaters.
507
eralogical phases. Starting from the sample representing the local fresh water (BB) and considering the processes suggested above the computer program NETPATH (Plummer et al., 1991) was used to calculate the mass transfer models, and was able to explain the chemical composition of most evolved VD groundwaters (PA). On the bases of the constraints derived from the achievement of water-rock chemical equilibrium some of the models were rejected. Water-rock interaction process involving VD groundwaters during their flow-path can be summarized as follows: BB waters + 0.89 C02, + 5.53 Albite + 0.14 KFeld + 1.41 Ca/Na EX + 0.41 Mg/Na EX= PA water +8.48 Kaolinite (6) 3.4 S04-rich waters
Figure 4. Relationship between isotopic composition of TDIC and amount of total carbon.
Groundwaters belonging to this group show the higher calcium and sulphate contents. Calculated saturation index with respect to gypsum close to 1 indicates that these samples have reached the chemical equilibrium with this mineralogical phase, as favoured by the its rapid dissolution rate. However, these waters show a Ca-excess with respect to SO4 ion concentration. This seems to suggest that gypsum not is the only source for calcium, but a dissolution process of carbonate minerals can also be inferred.
5 CONCLUSIONS
4 SOURCES OF CARBON DIOXIDE GAS The calculated values of partial pressure of carbon dioxide (pCO2) range between 10-2.5and 10aO.'atm and are higher than the atmospheric CO2 partial pressure (10-3.5). As inferred from mass balance model, a carbon dioxide gas phase seems to interact with groundwaters in the Hyblean Plateau. In order to individuate the origin of CO2, the isotopic composition of the total dissolved inorganic carbon has been determined. The 6 l 3 C ~ ~values l c range between -15 and -8 per mil vs V-PDB. The plot of Figure 4 suggests the existence of a mixing between two endmembers. The first end-member shows 6 1 3 C ~val~2 ues close to -25 per mil vs V-PDB, thus suggesting the presence a CO2 gas sources deriving from processes involving organic matter (transpiration, oxidation of organic compounds). The second one is characterized by more enriched 6l3Cc02 deriving from dissolution of C-bearing minerals (carbonates), andor CO2 solubilization coming from inorganic reservoir (e.g. magmatic or crustal). All the samples display different degrees of mixing between these two end-members.
Proposed geochemical models for groundwaters hosted within the Hyblean Plateau indicate that tiesh waters deriving from meteoric recharge become more saline due to mineralization processes involving water-rock interaction. Some waters show a large marine contribution. Carbonate dissolution and weathering of silicate minerals, together with ionexchange reactions are the most important geochemical processes. These processes are also favored by the presence of a CO2-gas phase deriving mainly from two end-members: one organic source depleted in heavy isotopes and an inorganic one deriving from inorganic reservoirs. REFERENCES Appelo C.A.J. & D. Postma. 1993. Geocheniistry groundwater and pollution. Rotterdam: Balkema Busemberg, E. & N. Plummer. 1989. Thermodynamics of magnesian calcite solid-solutions at 25°C and 1 atm total pressure. Geochim. Cosmochim. Acta 53: 1189-1208 Favara, R., Grassa, F. & M. Valenza. Climate and isotope features of precipitation in a semi-arid area (South-Eastern Sicily) and relationship with local groundwaters. In prep. Guggenheim, E.A. 1937. The theoretical basis of Raoult's law. Trans. Faraday Soc. 33:15 1-1 55 Parkhurst, D. 1995. User's guide to PHREEQC-a computer program for speciation, reaction path, advective-transport, and inverse geochemical calculations. USGS Water Res. Inv. 95-4227, 143pp Plummer, L.N., Prestemon, E.C. & D.L. Parkhust 1991. An Interactive code (NETPATH)for modeling NET geochemical reactions along afrow PATH. USGS Water Res. Invest. Rep. : 9 1-4078
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Monitoring of groundwater quality in Umbria (Central Italy) F.Frondini , G .Marchetti , A .Martinelli , L .Peruzzi & R .Crea ARPA- Umbria, Perugia, Italy
ABSTRACT: In Umbria (Central Italy) the Regional Environmental Protection Agency (ARPA) is monitoring the groundwater quality of the principal alluvial and carbonate aquifers of the region. The results of the monitoring activity point out that diffuse nitrate contamination is the main problem affecting the alluvial aquifers followed by local and diffuse contamination events with pesticides, hydrocarbons, AOX and heavy metals. In contrast, the waters discharged by the springs of the carbonate-evaporite aquifers are practically unpolluted and their protection is a primary objective of the local environmental authorities. 1 INTRODUCTION In Umbria, where more than 98% of the region's drinking water comes from groundwater and more than half of this total comes from wells, the protection of the quality of groundwater resources is a high priority for the local environmental authorities. Since the early eighties several groundwater monitoring projects and aquifer vulnerability studies were carried out in the critical areas of Umbria (Giaquinto et al. 1991, Marchetti 1995) and from 1997 started the PRISMAS project, funded by the Italian Ministry of Environment and by the Region of Umbria. The main purpose of the PRISMAS project was the definition of standard procedures for designing and exploiting a groundwater-monitoring network at a regional scale. Within this project the regional environmental protection agency of Umbria (ARPA) is monitoring the principal aquifers of the Umbria region with the following specific objectives: - the organisation of a permanent groundwater monitoring network at a regional scale in Umbria; - the definition of groundwater chemistry of each aquifer; - the identification of the seasonal and longterm variations in recharge and their influence on groundwater quality; - the study of the spatial and temporal variations of the chemical composition of groundwater related to diffuse and local contamination processes. In this work we present the principal results of the
PRISMAS monitoring project. In particular we will illustrate the natural geochemical characteristics of the studied aquifers and the effect of diffuse pollution on groundwater chemistry. Diffuse aquifer Contamination has become of increasing concern in Umbria because most of groundwater-quality problems are areally dispersed, including problems associated with the application of agricultural chemicals, human and animal wastes and the aggregate effects of many different point source contamination events.
2 GEOLOGY AND HYDROGEOLOGY The geology setting of the study area (Fig. 1) is the result of two main tectonic phases. A primary compressive phase produced, during the OligoceneMiocene, overthrusting of the Tuscan Nappe onto the Umbria-Marches Mesozoic carbonate-evaporite units, and the uplift of the Apennine chain (Miocene-Plio-Pleistocene). A subsequent tensional tectonic phase (Upper Miocene-Upper Pleistocene) overprinted the compressive structures and resulted in several NW-SE graben systems bounded by westward-dipping master faults and filled by Neogene sediments (Lavecchia 1990). The hydrogeological setting of the study area (Fig. 1) reflects the geological features described above (Boni et al. 1986, Giaquinto et al. 1991). The NW-SE graben systems are the seat of five alluvial aquifers, with thickness ranging from few meters to more than a hundred meters, composed of Quaternary sediments (sand, gravel and clay). The eastern 509
acteristics. We selected preferably domestic wells and public-supply wells rather than agricultural wells because they are not in use during the autumn and the winter. In order to correctly describe both the seasonal and the long-term variations we decided to use a seasonal sampling frequency of four campaigns a year.
3.2 Sampling procedures and analytical methods
Figure 1. Geological sketch map of Umbria and location of the monitoring points of the PRISMAS-Umbria Project.
sector of the study area (Apennine chain) is characterised by the presence of large carbonate-evaporite aquifers hosted by the Mesozoic permeable formations of the Umbria-Marches sedimentary sequence. These aquifers have large recharge areas and few highly representative springs discharging most of the ground water. Waters circulating in the upper part of the Umbria-Marches sedimentary sequence, composed of limestones and marly limestones, are characterised by a relatively rapid flow mechanism. Waters circulating in the lower part of the sequence, where dolostones and evaporites are the predominant rock types, are characterised by slower flow paths and longer residence times. 3 HYDROGEOCHEMISTRY 3.1 Selection of the monitoring network Starting from a database of about 1400 points sampled during previous studies, we selected a monitoring network made of 220 wells and springs. For the carbonate aquifers, the monitoring points have been selected on the basis of their discharge rate and considering their position with respect to the main geological structures. For the alluvial aquifers, the selection of the monitoring points has been done considering their statistical significance, the land use, the local geological heterogeneity and the well construction char-
Temperature, pH, conductivity, redox potential, dissolved oxygen and alkalinity were measured at the sampling site together with the measurement of the depth to water level. Water samples for chemical analysis were collected in two polyethylene bottles of 250 ml and 500 ml, respectively. The 500 ml bottle was acidified with 1 ml of concentrated HNO3 after filtration with a 0.45-pm membrane. The anions were determined by liquid chromatography, Ca and Mg by AAS, Na and K by AES, Fe, Mn, Zn, Pb Ni, Cd by GF-AAS using the acidified sample, Hg and As by AAS after hydride generation. The concentrations of pesticides, hydrocarbons, phenols and organic micropolluttants were determined by gaschromatography and gas-cromatography-mass spectroscopy using a SPE technique for the preconcentration of the samples. 3.3 Chemical composition of groundwater 3.3.1 Alluvial aquifers The waters of the alluvial aquifers are characterised by a variable composition and a salinity ranging from 0.4 to 1.6 g 1-I. From the Piper diagram of Figure 2a it is possible to distinguish three main groups: a) Ca(Mg)-HC03 waters, which represent the largest part of the sampled waters; b) Ca(Na)-HC03 waters, including only few samples from the Media Valle del Tevere and Valle Umbra aquifers; c) Ca(Mg)-S04(Cl) waters including some samples from the central part of the Valle Umbra aquifer. Group a) represents groundwaters with shallow circulation into the permeable levels of the alluvial aquifers where the most important geochemical process is the fast dissolution of carbonate minerals. Waters belonging to group b) evolve to higher Na concentrations because of cation exchange processes in non-steady state situations related to local pollution events, or to local circulation into organic matter-rich levels. The increasing Mg-SO4 character for waters of group c) is due to long residence times of groundwater into the lacustrine sediments outcropping in the central part of Valle Umbra. These sediments, characterised by a low permeability, are composed by clay, silt and silty sands containing sulphate minerals. 510
Table 1. Ranges of the saturation indexes and log PC02 values of the Apennine springs. water type
Ca-HC03
log PC02 SI CAL _ _ _ max min max min
_
SI DOL SI GYP _ _ _ _ max min max min
~
-1.5 -1.8 +0.1 -0.4 -1.0 -2.2 -2.0 -3.7
Ca(Mg)-S04 -1.6 -1.7 +0.1 -0.1 -0.3 -0.6 -1.0 -1.2
(1995), this process affects the waters circulating in the lower part of the Umbria-Marches sequence. A further process occurring in the springs of the Apennine area is the mixing between the Ca-HC03 waters and the Ca(Mg)-S04 waters. This process, clearly shown in the Piper diagram of Figure 2b, occurs when groundwaters of different types mix along fault planes crossing two different aquifers and feeding springs. 3.4 Geochemical indicators of difuse pollution The most important group of diffuse pollutants of regional importance in Umbria are agricultural chemicals, expecially fertilisers and pesticides. These compounds have been extensively monitored during the PRISMAS project. Other compounds have been used as diffuse pollution indicators (i.e. heavy metals and AOX for urban and industrial areas) or specific markers of point source contamination (hydrocarbons). The monitoring of the Apennine springs indicates that NO3, ammonia, organic pollutants and heavy metals are practically absent in groundwater of the carbonate-evaporite aquifers. In contrast, the waters of the alluvial aquifers show several signs of diffuse pollution. Groundwater nitrate concentrations are considerably higher than the natural background values and in several wells exceed the public health standards. Table 2 shows the median and the range of variation of NO3 in the alluvial aquifers: Valle Umbra, Media Valle del Tevere and Conca Eugubina show the higher contamination levels, while Conca Ternana and Alta Valle del Tevere are characterised by lower NO3 contents. The comparison of the NO3 concentrations measured during the period 1998-2000 with the data of earlier monitoring campaigns (Giaquinto et al. 1991, Marchetti 1995), suggests a general increase of the NO3 load in Umbria during the last 15 years. A clear example of nitrate long-term increase (Fig. 3) is given by the northern part of the Valle Umbra aquifer, where the mean NO3 concentration increased from 38 mg 1-' in 1985 to 57 mg 1-' during the 19982000 period. Considerin that in the northern sector of Valle Umbra (160 km ) the average total recharge
Figure 2. Piper diagrams: a) alluvial aquifers; b) springs of the Apennine aquifers.
3.3.2 ,4pennine springs The Piper diagram for the Apennine springs (Fig. 2b) shows that approximately 70% of the samples are Ca-HC03 in composition, 8% Ca(Mg)-S04 in composition, and the remaining samples have intermediate compositions and lie along a line between the Ca-HC03 and the Ca(Mg)-S04 end-members. The springs are characterised by discharge temperatures close to the average temperature of the recharge areas (9-12°C) and salinity lower than 1 g 1-'. In order to describe the processes governing groundwater composition in the Apennine aquifers, PcO2 and saturation indexes of calcite, dolomite and gypsum, were calculated for each spring sample with the aqueous speciation model PHREEQC (Parkhurst 1995). A summary of the saturation indexes and PCOZ values is reported in Table 1. The saturation indexes of spring samples suggest that a major process controlling groundwater composition is the dissolution of calcite close to equilibrium conditions (SI CAL between -0.4 and +0.1, with a mean value of 0). Calcite dissolution affects all the studied springs and is strictly controlled by the PC02 values. All the groundwaters are undersaturated with respect to gypsum (or anydrite). However, dissolution of gypsum (or anhydrite) is evident in the springs with the higher SO4 concentrations. With dissolution of CaS04, calcite precipitates and dolomite dissolves in order to maintain water-mineral equilibrium, shifting groundwater toward a Ca(Mg)-S04 composition (dedolomitization process). According to Frondini
5
51 1
~
Table 2. Nitrate content of the alluvial aquifers during the period 1998-2000. Concentrations are expressed in mg 1-I. aquifer
median
max
min
34 25 45 43 23
138 67 168 327 71
Conca Eugubina Alta Valle del Tevere Media Valle del Tevere Valle Umbra Conca Temana
Figure 3. Mean nitrate concentration and 95% confidence interval of the mean of 50 wells in the northern sector of the Valle Umbra aquifer during the period 1985-2000 (empty squares). The same diagram illustrates also the NO3 variations of the Petrignano-2 public supply well, for which a very detailed sequence of data is available (full squares - data from Cortina et al. 1999).
Table 3 shows the percentages of samples with detectable contents of pesticides, hydrocarbons and chlorinated hydrocarbons (AOX) of each aquifer. These compounds are not naturally dissolved in groundwater and their presence in detectable concentrations, even if lower than the limit of the Italian drinking water regulations, is an unequivocal indicator of aquifer contamination. The Valle Umbra and Media Valle del Tevere aquifers show the higher detection percentages of pesticides, Conca Ternana show the higher presence of AOX, and of these aquifers are related to the use of agricultural chemicals. The other alluvial aquifers show lower, but still frightening concentrations of nitrate and pesticides. Additionally, some areas are characterised by diffuse contamination processes with AOX and hydrocarbons due to the aggregate effect of many local sources. Finally, only a few cases of heavy metal contamination have been detected in the last two years. On the other hand, pesticides, AOX hydrocarbons, NO3 and heavy metals are practically absent in the waters discharged by the Apennine springs. Groundwater circulating in the carbonate-evaporite Apennine aquifers is a strategic virtually unpolluted resource and its protection should be a primary objective of the regional and national environmental authorities. REFERENCES
is 2.1 m3 s-' (Falcone et al. 1991), the NO3 quantity that is leached from the soil to the aquifer amounts to about 23.6 1O3 kg y-l k n ~ - ~ . Conca Eugubina has the maximum number of hydrocarbon detections. 4 CONCLUSIONS In Valle Umbra and Media Valle del Tevere the increasing NO3 content of groundwater and the large number of wells contaminated with pesticides indicate that the most important water-quality problems
Table 3. Detection percentage of pesticides, hydrocarbons and chlorinated hydrocarbons (AOX). aquifer
pesticides W O )
Conca Eugubina Alta Valle del Tevere Media Valle del Tevere Valle Umbra Conca Temana
hydrocarbons W O )
5 9 12 13 0
25 10 12 7 3
AOX
YW 25 3 11 17 48
Boni, C., Bono, P. & G. Capelli 1986. Schema idrogeologico dell'Italia Centrale. Mem.Soc.Geo1.h. 35: 991-1012. Cortina, C., Facchino, F., Ficiara, R., Giuliano, G., Marchetti, G., Martinelli, A. & F. Pennacchi 1999. Individuazione di in piano di intervento tecnico-normativo per il recupero di un acquifero alluvionale contaminato da nitrati di origine agricola. In Quaderni di Geologia Applicata; 3" Convegno Nazionale sulla Protezione e Gestione delle Acque Sotterranee per il 3' Millennio, Parma, 13-15 October 1999: 2.4 1-2.5 1. Bologna: Pitagora Editrice. Falcone, M., Marchetti, G., Martini, E. & G. Pizzi 1991. It model10 matematico di flusso dell'acquifero alluvionale della Valle Umbra. In S. Giaquinto, G. Marchetti, A. Martinelli & E. Martini (eds), Le acque sotterranee in Umbria: 6794. Perugia: Protagon. Frondini, F. 1995. Geochemistry of ground water in SouthCentral Umbria, Plinius 13: 79-83. Giaquinto, S., Marchetti, G., Martinelli, A. & E. Martini (eds) 1991. Le acque sotterranee in Umbria. Perugia: Protagon. Lavecchia, G. 1990. The Tyrrehenian-Apennine system: structural setting and seismotectogenesis. Tectonophysics 47: 263-296. Marchetti, G. (ed.) 1995. Studi sulla vulnerabilitri degli acquiferi, la Conca Ternana. Bologna: Pitagora Editrice. Parkhurst, D.L. 1995. User guide to PHREEQC - A computer program for speciation, reaction-path, advective-transport, and inverse geochemical calculations. U.S. Geol. Surv., Water-Resour. Invest. Rep, 95-4227, 15 1 pp.
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Salt water intrusion in the Pisa coastal plain (central Italy) F.Frondini ARPA- Umbria and Dipartimento di Scienze della Terra, Universitd di Perugia, Italy
A Zanzari & S .Giaquinto Dipartimento di Scienze della Terra, Universitd di Perugia, Italy
ABSTRACT: Three aquifers are identified in the Pisa coastal plain, central Italy: (1) a shallow phreatic coastal aquifer; (2) a sandy confined aquifer; (3) a deeper aquifer made up of gravel. The fraction of salt water in the coastal phreatic aquifer is calculated using a two component mass balance model of C1 and Br as between 0.1 and 8%. Saline intrusion occurs near the coastline and is associated with to the Arno river, the Scolmatore channel and the Lamone channel. The composition of the few samples of the deeper aquifers is described by a three component model. The three component mixing model shows that the salty contamination affects also the deeper aquifers (mainly in the southern sector) producing a progressive worsening of groundwater quality. 1 INTRODUCTION
2 HYDROGEOCHEMISTRY
The Pisa coastal plain is located west of Pisa (central Italy) and is bounded by the Serchio river to the north, the Scolmatore to the south and the Tyrrhenian sea to the west (Fig. 1). It is a large plain characterised by the presence of a dense surface drainage network. Sediments of marine origin (Upper Miocene-Upper Pleistocene) and continental sequences of fluvial and aeolian origin (Upper Pleistocene-Olocene) occur in outcrops. Three aquifer systems occur in the Pisa coastal plain (Rossi & Spandre 1994, Baldacci et al. 1994): (i) a shallow phreatic aquifer occuring in coastal dunes and commonly containing brackish water. This aquifer is separated from the deeper aquifers by a confining clay layer; (ii) a confined aquifer made up of sand and partly recharged by groundwater from the coastal dunes in the southern sector of the study area (Rossi & Spandre 1995, Baldacci et al. 1994); and (iii) a deeper confined aquifer made up of gravel. Two campaigns of groundwater sampling have been carried out in September 1997 and June 1998, during which 29 and 42 samples have been collected respectively, to: (i) define the characteristics, and amount, of the salt water intrusion in the shallow aquifer; and (ii) assess if the salt intrusion affect also the deeper aquifers. The location of the sampling sites is shown in Figure 1.
Most of the samples (Fig. 2) have a composition ranging from Ca-HC03 to Na-C1 and only few samples, belonging to the deep aquifers, are slightly shifted towards a Ca-SO4 composition. Salinity is generally lower than 1 g I-' for the Ca-HC03 samples and increase up to 3 g I-' with the increasing Na-C1 content. Saturation indexes indicate that all the samples are practically at equilibrium with Calcite and Dolomite
Figure 1. Study area and location of the sampling points.
513
Figure 2. Langelier-Ludwig diagram. Symbols refer to the September 1997 and June 1998 campaigns.
and undersaturated with respect to Gypsum. The samples of the phreatic aquifer are the result of a binary mixing process involving infiltration water and sea water (or fossil seawater). The few samples of the deeper aquifers cannot be described by a simple mixing of two pure terms so a third component should be considered. In fact, the Ca-Mg-Na+K and HC03-SO4-Cl diagrams (Fig. 3a,b) clearly indicate the presence of a Ca-SO4 component in the samples of the confined aquifers. The shift of these samples towards a Ca-SO4 type water is probably related to mixing with groundwaters from the carbonate-evaporite formations of the Monti Pisani area (north-east of the Pisa Plain). The waters discharged by the springs of San Giuliano Terme, located few kilometers north-east of Pisa (Grassi et al. 1992), are typical of the Ca-SO4 type. The Langelier-Ludwig diagram also shows that the (Na+K)/(Ca+Mg) ratio is not simply fixed by the seawatedfreshwater ratio but is characterised by variations which are probably due to secondary processes such as adsorption, cation exchange (Appelo & Postma 1994), and/or local pollution events. 3 MIXING MODELS 3.1 Coastal phreatic aquifer For each sample of the phreatic coastal aquifer the amount of salt water has been calculated considering the mass balance of C1 and Br. These chemical species have been selected because their concentration in groundwater is not directly controlled by any equilibria with a mineral phase. The C1 and Br contents of groundwater depend only on the fractions of salt water and infiltrating water and on the C1 and Br content of salt water and infiltrating water:
where X, and X m are the fractions of salt water and infiltrating water and Cis and Cim are the concentrations of the chemical species i (C1 or Br) in the salt water and in the infiltrating water respectively. The compositions of seawater and of the sample with the lowest salinity have been considered as end-members of this mixing model. The fractions of salty water computed using the C1 mass-balance are very similar to the fractions estimated using the Br mass-balance and range from 0.1 to 8%. The geographic distribution of the saline intrusion (Fig. 4) show that: (i) the intrusion occurs near the coastline and from lateral encroachments connected to the river Arno estuary, the Scolmatore channel and the Lamone channel; (ii) the relative maximum of salinity, not directly connected to the sea, which is present in the south of the study area roughly coincide
Figure 3. Ca-Mg-Na+K (a) and H C O ~ - S O ~ -(b) C ~ternary &agrams.
514
Figure 6 . Triangular diagram showing the results of the three component mixing model. Xm, Xc-ev and Xs represent the shallow, the carbonate-evaporite and the seawater components.
3.2 Deep conjned aquifers Figure 4. Map of the average salt water content in the shallow aquifer during the period September 1997-June 1998.
with the area where the phreatic aquifer is connected to the confined sandy aquifer (Rossi & Spandre 1995, Baldacci et al. 1994); (iii) a relative minimum is present in the northern sector of the study area, close to the Serchio River; (iiii) a large intrusions is located west of Pisa. The comparison of the September 1997 and June 1998 analyses indicate that in the wells close to the coastline the fraction of saline water is greater at the end of the summer when groundwater levels are lower.
The triangular diagram Mg-SO4-Cl (Fig. 5 ) , where the samples of the phreatic and confined aquifers are compared to the composition of the San Giuliano thermal spring (data from Grassi et al. 1992), allows to differentiate the SO4 and Mg contribution due to the seawater intrusion from the contribution related to the carbonate-evaporite aquifers. Samples of the deep aquifers are shifted towards the composition of the San Giuliano thermal spring (close to the SO4 corner) supporting the hypothesis of a partial recharge fiom the carbonate-evaporite formations. The relative amounts of the meteoric, salty and carbonate-evaporite components in the samples of the deep aquifers are given by:
(3)
x, + x, + XC-,,
Figure 5. Triangular diagram Mg-S04-C1. San Giuliano spring and seawater compositions are reported.
=
I
(4)
where the notations are the same of equations (1) and (2), and X,.ev refers to the component deriving from the carbonate-evaporite formations. The equations have been solved for X,, Xm and Xc-ev, using C1 and SO4 and considering the San Giuliano spring as the end-member of the sulphate deep component (carbonate-evaporite component). The results of this three component mixing model are shown in Figure 6 and indicate that: (i) the component related to the recharge from the carbonate-evaporite formations range between 10% and 40% in the samples of the confined aquifers; (ii) the salt water intrusion occurs in the confined aquifer in the southern part of the study area (3%o<xs<5%o).In this area the phreatic aquifer is connected to the confined sandy aquifer (Rossi & Spandre 1995, Baldacci et al. 1994) and the expansion of the saline intrusion to the deeper aquifer is due to cross contamination between the two aquifers. 515
4 CONCLUSIONS The shallow groundwaters of the Pisa coastal plain show evidence of salt water intrusion. The fraction of saline water, estimated for each sample through a mass balance of C1 and Br, range from 0% to 8%. The composition of the samples of the deeper aquifers is described by a three component model (meteoric water, salt water and a Ca-SO4 component from the carbonate-evaporite formations). The results of this three component mixing show that the salty contamination affects also the deeper aquifers in the southern part of the study area where the confined sandy aquifer is recharged by groundwater from the shallow phreatic aquifer. Cross contamination between aquifers play an important role for the expansion of the saline intrusion to the deeper levels and the resulting worsening in chemical quality of groundwater may affect also wells used for drinking-water supplies. The protection of the deep confined aquifers represent one of the main problems and should be a primary objective in managing the withdrawals of water in the study area. REFERENCES Appelo, C.A.J. & D. Postma 1994. Geochemistry, groundwater and pollution. Rotterdam: Balkema. Baldacci, F., Bellini, L. & G. Raggi 1994. Le risorse idriche sotterranee della Pianura Pisana Atti Soc. Tosc. Sci. Nut. Mem. 101 : 241-322. Grassi, S., Carosi, R., Marroni, M. & D. Ancora 1992. I1 sistema idrotermale di S. Giuliano Terme (PI) : note di geologia, idrogeologia e geochimica. Boll. Soc. Geol. It. 1 1 1 : 303313. Rossi, S. & R. Spandre 1994. L’intrusione marina nella falda artesiana in ghiaia nel litorale pisano. Acque Sotterrunee 43: 51-58. Rossi, S. & R. Spandre 1995. Caratteristiche idrochimiche della l a falda artesiana in sabbia nei dintorni della citta di Pisa. Acque Sotterrunee 48: 27-36.
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Salinization in coastal plain of Grosseto: hydrochemical study E.Gim6nez Forcada Department of Environmental and Agroforestal Sciences. Catholic University of Avila (Spain)
A.Bencini & G.Pranzini Department of Earth Sciences. University of Florence. Florence (Italy)
ABSTRACT: The main purpose of this paper is to study the chemical mechanisms causing groundwater salinization in the Grosseto plain (Southern Tuscany) and identify the geochemical processes involved in the behaviour of dissolved constituents. Composition of waters shows an important salinization in both coastal and inner areas. Marine intrusion is a likely mechanism producing water salinization in coastal areas, while in inner zones groundwater salinization seems to be due to sulphate waters that characterises the thermal manifestations of the area. Hydrochemical study clearly shows that the plain is affected by the influence of both saline fluxes and also by an influx of bicarbonate-fresh water, and thus we can think that Grosseto plain is affected by mixing of these three end members. In addition to the mixing phenomena, cation exchange processes participate to the variability of chemical facies, and these water-rock reactions seem to play an important role in controlling cation concentrations. 1 INTRODUCTION In coastal areas differentiation of salinization sources is particularly complex due to the large hydrochemical variability caused by the overlapping of different processes. Recent hydrogeochemical studies (Bencini & Pranzini 1993; Pranzini 1995; Pranzini & Bencini, 1996), based on chemical and stable isotopic data, stated that chemical characteristics of Grosseto plain groundwater show at least two main salinization processes: marine intrusion and mixing with sulphate thermal waters. Similarly, these studies suggested the importance of ionic exchange processes in the genesis of alkaline-earth chloride waters. The present work examines and identifies each one of the geochemical processes involved in the behaviour of dissolved constituents in Grosseto plain groundwaters.
2 GEOLOGICAL AND HYDROGEOLOGICAL BACKGROUND The coastal plain of Grosseto (Fig. 1) corresponds to a tectonic depression where a notable thickness of sediments has been accumulated. The rock formations surrounding the plain belong to Tuscan Complex (Triassic-Oligocene): sandstones, siltstones, limestones, marls, marly limestones and shales (Pranzini & Bencini 1996). The Grosseto plain
mainly consists of silts, clays with gravels and sands. Gravel and sand layers are the most important aquifers and are exploited mainly for municipal water supplies. The transmissivity ranges from 10-2m2/sto 10-3 m2/s, and the storage coefficient varies from 1.2 10-4 to 6.1 10-6. The piezometric surface is frequently below sea level. The progressive extension of the surface areas below sea level strongly suggest overexploitation of water resources. (Pranzini 1991, 1995). 3 SAMPLING AND ANALYTICAL METHODS 105 samples were collected in the Quaternary aquifers (Fig. 1). The sampling and analytical procedures were performed according to the American Public Health Association (APHA, AWWA & WPCF 1995) recommendations. 4 GROUNDWATER CHEMISTRY 4.1 Chemicalfacies and area1 distribution The term “hydrochemical facies” defines the sum of all the primary chemical characteristics of a certain water. Accordingly, waters have been classified by using a graphical method based on the expanded Durov Diagram (Bourdom & Mazloum 1958; Lloyd 1967). Figure 2 shows that groundwaters of Grosseto 517
Figure 1. Geological-hydrogeological map of Grosseto Plain. 1: Slightly permeable or impermeable rocks (sandstones, siltstones, marls, marly limestones, shales); 2: Mesozoic carbonate rocks, highly permeable because of fracturing and karst; 3: Dune and beach sands; 4: Alluvial deposits. e : Sampling sites.
Distribution of facies does not allow establishing clearly a zoning of the area according to the distal or proximal salinization source, although the chloride facies dominates in coastal areas and sulphate facies in inland zones. Bicarbonate waters are usually identified with the lateral inflow fiom aquifers of border areas and in several dispersed points in the plain that represent the most superficial level of the aquifer.
plain belong to six facies, though Ca,Mg-HCO,, Ca,Mg-SO,, Na-C1 and Ca,Mg-C1 waters are more abundant.
4.2 Mixing processes
Figure 2. Durov Diagram modified. Hydrochemical facies for samples and end members participating in the mixing phenomena. DEx: Direct Cation Exchange; Ex: Inverse Cation Exchange. (-) Mixing lines.
The chemical character of groundwaters in the Grosseto plain suggests that marine intrusion is not the only factor of salinization in this aquifer. In fact, rS042-/rC1-ratios are higher than that of the sea water, showing that sulphate contribution is not related to the sea water. rS042-/rC1-ratios increase fiom the coastal zone towards the inland. Thus, marine intrusion process is located in a small strip near the shoreline, while almost all aquifers are affected by the influence of non-marine derived sulphate waters. waNorth Of Grosset' the presence Of ters (Roselle) is well known. These waters have a temperature of about 36-37°C and a chemical com518
Figure 3. Relationship between S042- and Cl- contents for all samples. Hypothetical three component mixing between fresh (FW), sea (sW) and sulphate (SW) waters is also represented.
Figure 4. Chloride waters. Trends of ANa+AK deficits and ACa+AMg(-ASO,) excesses. The samples are ordered according to increasing of chloride concentration.
position that classifies them as Ca,Mg-SO,. The origin of these waters has been related to Triassic evaporites, which represent the geothennal reservoir (Bencini et al. 1977). Grosseto plain groundwaters have a temperature of 20°C to 22"C, and locally may reach 27"C, which indicates a contribution of thermal waters in salinization of the aquifer. The existence of two different saline fluxes (sea water and sulphate water) lets us suggest different degrees of mixing into the aquifer. Thus we can say that the composition of Grosseto waters can be related to three component mixing: sea water (sW)sulphate water (SW)-bicarbonate-fresh water (FW). Considering SO,2- as the conservative tracer for sulphate water influence (high rS042-/rC1-ratio) and C1- as a conservative tracer for sea water mixing (low rS0,2-/rCl- ratio), fresh water-sea water alignment represents marine intrusion, fresh watersulphate water alignment relates sulphate intrusion and the field defined by the three end members represents the concurrent effect of mixing of the three end members (Fig. 3). The overlapping of both salinization processes (sulphate waters influx and marine intrusion), together with the dilution effect provoked by fresh waters, produces a great variety in the chemistry of the waters from the Grosseto plain.
cesses. It must be pointed out the predominance of Ca-C1 waters, probably related to the participation of inverse cation exchange phenomena. Likewise, some sulphate waters show Ca-Mg concentrations lower than those determined by hypothetical mixing line, which can be interpreted as a result of direct ionic exchange processes participation. In fact, it is possible to recognize the presence of Na-SO, and NaHCO, facies, though in only two samples. Calculation of ionic deltas (Aim) is useful while identifying ionic exchange processes (GimCnez et al, 1995). Ionic concentrations in the water samples are compared with those resulting from theoretical mixing between fresh and sea water. Differences between the observed and expected concentrations are expressed as Aim (in meq/l). In the Grosseto plain, we found positives values for SO,2-, Ca2+ and Mg2+, and negative for Na+ and K+, which indicate excesses and deficits of these ions regarding the conservative mixing. Nevertheless, Ca2+ and Mg2+ values are much higher than values those expected by single ionic exchange, shown mainly by Na values. These excesses can partly be attributed to the influence of Ca,Mg-SO, waters that provokes an increment of alkaline-earth elements. ANa+AK and ACat-AMg-AS04 values for chloride waters are reported in Figure 4, where water samples are orderd according to the increase of chloride concentration. ACa+AMg-AS04 values are representative of Ca2+ and Mg2+excesses not in relation to the influence of sulphate waters. This figure shows participation of exchange reactions in the chemical character of groundwaters.
5 SECONDARY PROCESSES According to previous studies and hydrochemical methodology about exchange processes in coastal areas (Lloyd & Heathcote 1985; Tellam & Lloyd 1986; Appelo & Postma 1993) the water samples have been analysed in relation to the ionic exchange phenomena. In aquifers with an important clayey fraction, cationic exchange processes tend to compensate the chemical variations of water composition linked to mixing waters. Figure 2 displays the cation concentration variations as a function of mixing pro-
6 CONCLUSIONS Groundwater salinization in Grosseto plain is rather widespread and salinization is controlled by two principal saline flows: sea water and sulphate water. Hydrochemical study clearly shows that sea water 519
intrusion prevails in coastal areas, while sulphate waters influx is more important in the inland. Nevertheless the plain is affected by the influence of both saline fluxes and also by participating of bicarbonate-fresh water, and so we can think that Grosseto plain aquifer suffers a mixing of three end members. rS042-/rC1- ratio is the hydrochemical parameter, which better differentiates both saline intrusions. In fact, waters affected by marine intrusion have ratios similar to that of the sea water, while waters associated to sulphate influx have higher ratios. Chemical variability of waters associated with clay sediments produces cation exchange processes. Inverse cation exchange provokes the occurrence of Ca,Mg-C1 facies, while direct cation exchanges can induces Na-SO, and Na-HCO, facies (only two samples are this composition). More detailed studies about these processes would allow to support the exposed hypothesis.
Tellam, J.H. & J.W. Lloyd (1986) Problems in the recognition of sea water intrusion by chemical means: An example of apparent chemical equivalence. Quaterb Journal of Engi19: 389-398. neering
REFERENCES APHA, AWWA & WPCF (1995). Standard methods for the examination of water and wastewater. 18th edition, APHA. Washington. Appelo, C.A.J. & D. Postma,( 1993). Geochemistry, groundwater and pollution. Ed.: AA Balkema. Rotterdam. Brookfield. 536 pp. Bencini, A., Duchi, V. and M. Martini (1977). Geochemistry of thermal springs of Tuscany (Italy). Chem. Geol., 19: 229252. Bencini, A. & G. Pranzini (1993). The salinization of groundwaters in the Grosseto Plain (Tuscany, Italy). In: Custodio, E. & Galofre, A. (eds). Study and Modelling of Salt Water Intrusion into Aquifers. Proc. 12th Salt Water Intrusion Meeting: Barcelona, 1992: 161-175. Bourdom, D.J. & S. Mazloum(1958). Some chemical types of groundwaters from Syria UNESCO. Symp. Teheran: 73-90. Unesco. Paris. Bravetti, L. & G. Pranzini (1987) L'evoluzione quatemaria della pianura di Grosseto (Toscana): prima interpretazione dei dati di sottosuolo. Geogr. Fis. Dinam. Quat., 10: 85-92. Gimenez, E., Fidelibus, M.D. and I. More11 (1995). Metodologia de Analisis de Facies Hidroquimica aplicada a1 estudio de la intrusion marina en acuiferos detriticos costeros: Aplicacion a la Plana de Oropesa (Castellon). Hidrogeologia, 11 55-72 Lloyd, J.W. (1967) The hydrochemistry of the aquifers of north-eastem Jordan. J. Hydrology, 3: 319-330. Lloyd, J.W. & J.A. Heathcote (1985) Natural inorganic hydrochemistiy in relation to groundwater. An introduction. Ed.: Clarendon Press. 296 pp. Pranzini, G. & A.. Bencini (1996). Groundwater salinization in Southern Tuscany (Italy). Proc. Salt Water Intrusion Meeting SWIM'96: 261-270. Malmo (Sweden). Pranzini, G. (1991) Piano di Bacino del Fiume Ombrone. Studio Idrogeologico. Studi preliminari. Regione Toscana. Unpublished Pranzini, G. (1995) Studio della salinizzazione delle acque di sottosuolo dell 'area costiera fra Castiglione della Pescaia e Orbetello. Regione Toscana (Dipartimento Agricoltura e Foreste). Unpublished.
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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Biogeochemical cycles of chloride, nitrogen, sulphate and iron in a phreatic aquifer system in The Netherlands Jasper Griffioen & Thomas Keijzer TNO Netherlands Institute of Applied Geosciences, Delfi, The Netherlands
ABSTRACT: We studied biogeochemical cycles of a groundwater system defined by a drinking water abstraction with dominantly forest and agriculture in the infiltration area. The average yearly input of N, S and C1 was calculated from data about atmospheric deposition, manuring, plant uptake and harvesting. The output was calculated using data of groundwater abstracted. The area-averaged input of C1 is near-constant in time since 1950, whereas the output increases to this input concentration from 1950 to 1980 and subsequently remains near-constant. The input of N species and SO4 increases strongly after Second World War and decreases since 1985 and 1965, respectively. Nitrate and S04, and to some extent Fe, increase step-wise in groundwater abstracted. Here, NO3 is lower than the input concentration corrected for conservative transport to the abstraction. Sulphate increases to a figure larger than the input concentration. These observations are explained by nitrate reduction associated with pyrite oxidation, which is also indicated by groundwater quality at observation wells. 1 INTRODUCTION Human activities influence the composition of infiltrating groundwater, threatening groundwater as a resource for drinking water. The human impact can be evaluated in terms of the disturbance of biogeochemical cycles: it is evaluated by comparison of the input/output relationship for the natural situation and that for the disturbed situation. Biogeochemical cycles of elements are mainly studied at global scale and local scale. Biogeochemical cycles of elements in groundwater systems at the regional scale have received little attention. Tesoriero et al. (2000) studied the subsurface biogeochemical cycle of nitrate of a hydrological catchment. Unfortunately, the output of this catchment was unknown, which hinders a rigorous input-output comparison (cf. Blanchard & Lerch, 2000). The purpose of this investigation was to compare the long-term input of species at the surface with the long-term output of species at the drinking water production site in order: 1. to identify temporal changes in input and output, including anthropogenic contamination and 2. to characterise the role of subsurface geochemical processes to alter the groundwater composition during flow. The subsurface biogeochemical cycle of four species was determined: chloride, nitrogen, sulphate and iron. Chloride behaves conservatively in the subsurface and serves as reference. The other three species are
involved in redox processes. The cycles were determined for the regional infiltration area of the phreatic drinking water abstraction at Holten, The Netherlands.
2 STUDY AREA The study area is situated in Eastern Netherlands at a location with unconsolidated Tertiary and Quaternary deposits. The altitude lies 20 to 55 meters above sea level. A phreatic aquifer happens which stretches to a depth of 70 m below sea level. The northern part of the infiltration area is composed of a Saalian ice-pushed ridge, that continues subsurface into southwestern direction. The icepushed deposits reach the surface again at the southwestern corner of the infiltration area. The depth of the ice-pushed deposits is at least 5 meters below sea level at the winning location. The aquifer sands are mainly of fluvial origin. Weichselian eolian and fluvio-glacial sands are present in the eastern and southern low-lying area on top of the ice-pushed deposits. These sands are up to 10 m. thick. The drinking water production started in 1957 at 0.14 Mm3/y and gradually increased to about 2.4 Mm3/year in 1969. The amount abstracted is about 2.0 Mm3/y since 1990. Groundwater is abstracted at 20 wells. Small ditches occur in the eastern corner of 521
the infiltration area, which drain eastwards. The 20, 50 and 80 percentile values of the travel time from groundwater table to abstraction well are 8.2, 34 and 96 years, respectively. At present, 34% of the infiltration area is grass land, 6.7% is maize land and 11% is urbanized (including regional roads and rail roads). The rest is dominantly forest. The maize area was used for growth of potatoes and cereals before 1975. The agricultural and urbanized areas occur at the Weichselian sands. 3 DATA COLLECTION AND CALCULATION The mass balance model was set-up on a yearly basis as a two box model. Box 1 is the unsaturated zone from surface to groundwater table and Box 2 is the groundwater saturated zone from groundwater table to abstraction well. The following terms were considered as input at surface: 1. atmospheric deposition, 2. fertilizer salts, and 3. animal manure. Data of yearly precipitation since 1911 was derived from a station 13 km away; calculated data of yearly crop evapotranspiration were obtained for a station 90 km away (KNMI, Royal Netherlands Meteorological Institute). The average annual precipitation is 762 k 135 mdyear and the average annual crop evapotranspiration is 543 & 33 mdyear. Differences among crops and trees and dependency of actual evapotranspiration on depth of groundwater table were neglected. The average groundwater recharge was therefore set equal to the difference between annual precipitation and annual evapotranspiration, which is 219 m d y . Data of precipitation quality since 1978 was obtained from a station 25 km away (RIVM, National Institute of Public Health and the Environment). Wet and dry deposition of S02, NO, and NH, since 1980 were obtained from Bleeker & Erisman (1996). Atmospheric emission of SO2 before 1990 was obtained from Mylona (1996). The data about SO2 emission and deposition were combined for the period 1980-1990 in order to deduce a relationship between SO2 emission and deposition before 1980. Complete conversion of SO2 deposited to SO4 dissolved was assumed. Data about emission of NH, and NO, before 1990 were obtained from Stuyfzand (1993) and also combined with more recent data about deposition to yield data about deposition. Data about fertilizer salts since 1900 and animal manure since 1950 were obtained from CBS, Central Office of Statistics, and LEI, Agricultural Economic Institute. Annual input of fertilizer salts and animal manure per hectare agricultural land was set equal to the total amount of fertilizer salt consumed or manure produced in The Netherlands, divided by the area of agricultural land in the Netherlands in that
year. Here, distinction is made among types of fertilizer salt, types of animal manure and its state (solid or water-based). The elemental composition of fertilizer salts and animal manure was next considered, where the composition was assumed to be constant in time. Data about volatilization of NH3 from manure and salts since 1980 were obtained from CBS. The volatilization from agricultural lands is about equal as the total deposition of NH, expressed per hectare agricultural land. The two terms were therefore neutralised by each other for agricultural land. Complete nitrification of NH4 remaining and NO2 added was assumed within the unsaturated zone. The temporal changes in agricultural land use were obtained since 1980 for the municipality of Holten and before 1980 for an adjacent, identical municipality. Leaching of No3 within the unsaturated zone of agricultural lands was considered as output term for Box 1, using empirical functions (Beekman, 1998). Agricultural land use was combined with national data about crop production per hectare and elemental composition of crops, to yield output of C1 per hectare agricultural land. Uptake of S by agricultural plants other than tuberous plants was set equal to 10 kg S/ha.y, which is a minimum value (Stevenson, 1986; Postma et al., 1999). No simple models exist for nitrate leaching out of natural lands. The leaching depends among others on age of forest or heather, cutting down of trees, and carbon dynamics. Nitrogen uptake by trees was set equal to 6 kg N/ha. y, whereas conservative behaviour was assumed for SO4 in the unsaturated zone (De Vries, 1994), and also for C1. All calculations were performed on basis of kilograms per hectare, taking into account the fractions of land use types within the infiltration area. For C1 and SO4, the sum of input terms minus the output via plants (Box 1) was set equal to the input of Box 2. Here, the figures were converted to concentration basis by dividing with the average yearly net groundwater recharge. The input of Box 2 is compared with the output at the winning location in order to characterise the importance of subsurface processes. The temporal input curve was also recalculated to an output curve for conservative transport, taking into account the travel-time distribution of groundwater flowing from infiltration area to the winning location. The traveltime distribution was calculated with a numerical geohydrological model of the area. The input of Fe to Box 2 was set to 0, because shallow groundwater is near-neutral and oxic or suboxic. At present, sixteen wells abstract groundwater between 10 and 40 meters below surface and four wells abstract groundwater between 40 and 75 meters below surface. These four wells are operative since 1985 to avoid clogging of the shallow wells due to mixing of aerobic and anaerobic groundwater in the vicinity of
522
the well casings. The major groundwater composition was determined after installation of each well and since 1983 on a quarterly basis. For an operational well, Fe is measured at least yearly since 1957, C1 since 1964 and NO3 since 1967. The total amount yearly abstracted is recorded since the start of the winning. The pumping hours per well are yearly recorded since 1976, which allows mass balances per well. Equal pumping hours per well were assumed for the period before 1976. This is reasonable because the groundwater composition varied less in the past than at present. The amount of ion removed per well is set equal to the amount of pumping hours per year multiplied with discharge per hour and yearly averaged concentration. The amounts per well are summed for all wells in operation to obtain the total amount removed. This is divided by the annually abstracted amount to yield average output concentration. Observation wells are present within the infiltration area, that are sampled occasionally.
Figure I . Input of N-species at the surface
to 1980 to the groundwater input concentration, which has remained remarkably constant since Second World War. Figure 3 presents a comparison between the input of N 0 3 , the input corrected for residence time from groundwater table to abstraction wells, and the output at the winning location. The data show that input is substantially larger than the output. This is explained in two different ways: errors in the calculation of the input and nitrate reduction in the aquifer. Firstly, the assumption that the inputs of fertilizer and manure are equally distributed among
4 RESULTS AND DISCUSSION The general tendency for the atmospheric deposition of N and S is that it strongly increased after Second World War and maximum values occurred in the eighties and sixties, respectively. Intensification of agriculture also started after Second World War and continued to the end of the eighties. Subsequently, the intense agricultural use levelled off somewhat. Figure 1 shows the resulting figures of N for the study area. The data show that the contribution of atmospheric deposition is small for agricultural lands. This also holds for C1, but does not hold for SO4 (not shown). Atmospheric deposition was in the sixties the most important input term for SO4 in agricultural lands. The intensification of human activities since Second World War also causes increased leaching to groundwater, because increased yield per hectare and uptake by crops does not follow increased input by deposition and manuring. Figure 2 presents the output via groundwater abstraction. Chloride, SO4 and NO3 increase in time, whereas Fe remains near-constant. The erratic behaviour is due to changes in abstraction well connection among years: wells having high concentrations are used many hours in one year and few hours in the next year. A general increase in concentration occurred in the sixties for all four ions, and a second, sharp increase is evident around 1989 for NO3 and SO4. The concentration increases from the beginning of abstraction to present are three to sixfold. They mean a deterioration of drinking water quality, but the concentrations are still below the EU standards for drinking water. The C1 concentration of groundwater abstracted had increased from 1950 I
Figure 2. Output composition of groundwater at the drinking water abstraction.
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ACKNOWLEDGMENTS The WMO, Drinking Water Company Overiljssel, is gratefully acknowledged to make available the data about the pumping wells, and other data as well. Iwaco Consultancy is thanked for making available the geohydrological flow model.
REFERENCES
Figure 3. Input and output concentration (measured and calculated conservative breakthrough) of NO3 for groundwater.
the agricultural lands of the Netherlands will not hold in reality. Groundwater from several shallow observation wells in the southern area has NO3 concentrations of 150 to 200 mg/l. Extensive leaching of NO3 thus can occur from the agricultural lands. Secondly, groundwater analyses available point out nitrate reduction associated with pyrite oxidation: Fe concentrations up to 54 mg/l are observed and multi-screen wells also show increased SO4 concentration at observation screens below observation screens with high NO3 concentrations. Nitrate reduction associated with pyrite oxidation is thus ongoing within the subsurface of the infiltration area. The small temporal increase of Fe at the winning location compared to the SO4 increase is explained by retardation of Fe mobilized due to cation-exchange with other cations.
Beekman, W. 1998. SPREAD. Prediction of NO3, hardness, C1 and SO., in shallow groundwater. Kiwa, report no. SWE 98.012, 32 pp (in Dutch). Blanchard, P.E. & R.N. Lerch 2000. Watershed vulnerability to losses of agricultural chemicals: Interactions of chemistry, hydrology, and land-use. Environ. Sci. & Technol. 34: 33 15-3322. Bleeker, A. & J.W. Erisman 1996. Deposition of acidifying components in The Netherlands during 1980-1995. RIVM, report no. 722108018,55 pp (in Dutch). De Vries, W. 1994. Soil response to acid deposition at difSerent regional scales. Ph.D thesis, Agricultural University Wageningen, The Netherlands, 481 pp. Mylona, S. 1996. Sulphur dioxide emissions in Europe 18801991 and their effect of sulphur concentrations and depositions. Tellus 48B: 662-689. Postma, R., P.J. Van Erp, & R. Saanen 1999. Quantifying the s u l k supply to agricultural crops. Meststofen 1999: 28-35 (in Dutch). Stevenson, F.J. 1986. Cycles of soil. Wiley Interscience, New York, 380 pp. Stuyfzand, P.J. 1993. Hydrochemistry and hydrology of the coastal dune area of the Western Netherlands. Ph.D thesis, Free University, Amsterdam, The Netherlands, 366 pp. Tesoriero, A.J., Liebscher, H. & S.E. Cox 2000. Mechanism and rate of denitrification in an agricultural watershed: Electron and mass balance along groundwater flow paths. Water Resour. Res. 36: 1545-1559.
5 CONCLUSIONS The temporal input of elements at the surface is compared with the temporal output at a drinking water winning. Anthropogenic disturbance of the subsurface biogeochemical cycles is clearly indicated: the input had increased several times for N-species and SO4. Transport of NO3 is reduced by subsurface redox processes, where Fe and SO4 are mobilized. The approach made shows that biogeochemical cycles can be well studied in a groundwater system that is determined by groundwater abstraction, because the system is closed.
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Magnesium concentration control in groundwaters in Iceland I .Gunnarson, S .Arn6rsson & S .Jakobsson Science Institute, University of Iceland, Dunhagi 3, 107 Reykjavik, Iceland
ABSTRACT: Precipitation experiments involving magnesium silicate carried out in alkaline solutions in the temperature range 53-150°C resulted in an equation describing the solubility constant of poorly crystalline antigorite in the range 0-200°C. The saturation state of the magnesium silicate phase precipitated was evaluated in natural water in Iceland. Rivers and streams, peat soil water and lakes are generally undersaturated with the magnesium silicate but groundwaters in the basaltic terrain of Iceland are close to saturation. It is, therefore concluded that a magnesium silicate phase similar to that precipitated in this study controls magnesium concentrations in Icelandic groundwaters. sure below 100°C and at vapor saturation pressure above 100°C. Experiments above 100°C were carried out in a titanium reactor but below 100°C in polypropylene bottles in temperature controlled water bath. The experimental procedure was as follows: Solution containing dissolved silica was heated to the appropriate temperature. Magnesium chloride solution was injected into the solution to bring about precipitation of magnesium silicate. The pH value was controlled with ammonium or boric acid buffers. Once the experiment was started, the solution was sampled periodically, silica and magnesium concentrations analyzed and the pH value measured. All samples were filtered through 0.2 micrometer filter membrane. Titrations were carried out at room temperature where the pH value of a solution containing aqueous magnesium and silica was raised to produce precipitation of magnesium silicate.
1 INTRODUCTION Magnesium concentrations in natural waters in Iceland vary by several orders of magnitude. Surface waters which have a pH value between 6-8 have magnesium concentration between 0.2-15 ppm. The pH value of surface water is controlled by steady state condition between the production of protons from atmospheric and organic CO;! and their consumption by reaction with the surrounding rock (Gislason & Arn6rsson 1993). When surface waters seep into the ground, they become isolated from the atmosphere and organic matter. Proton-cation exchange reactions with the surrounding rock raises the pH (Arnorsson 1995). In 10-20°C groundwater the pH value is as high as 10 and these groundwaters are depleted in magnesium. It is, therefore, evident that considerable magnesium precipitation occurs when surface water seeps into the ground and reacts progressively with the bedrock and that the precipitating phase controls the aquatic magnesium concentrations. In this study the precipitation of a magnesium silicate was studied experimentally in the temperature range 53-150°C. The purpose was to find out which magnesium silicate phase precipitates easily from alkaline solutions and to obtain a solubility constant for that phase.
3 RESULTS The precipitation of the magnesium silicate was fast. The solutions attained equilibrium with the precipitated phase from within an hour up to a month. This time depends on the initial degree of supersaturation and temperature. Both increased supersaturation and increased temperature enhance the precipitation rate. In the titration experiments the precipitation started when the pH was as high as 10. The solution became milky colored and the drop in aqueous magnesium and silica was simultaneous. In separate ex-
2 EXPERIMENTAL DESIGN Experiments were made in alkaline solutions in the temperature interval 53-150°C at atmospheric pres-
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periments, in which no silica was present, magnesium precipitation started at higher pH than when silica was present.
3.I Chemical composition The drop in magnesium concentration as the precipitation reaction progressed always exceeded that of silica (Fig. 1). Linear regression through the data points in Figure 1 gave a line with a slope of 1.35 which corresponds to the Mg/Si ratio of the precipitate. The Mg/Si ratio in natural antigorite is between 1.35 and 1.425 (Kunze 1961, Shitov & Zvyagin 1966). This ratio is not found in any other magnesium silicate. For this study the Mg/Si ratio of 4/3 (1.33) has been chosen for the precipitate. Because of the small amount of magnesium silicate precipitated in each experiment, it is difficult to measure its water content. With the water content of the magnesium silicate not known, it is necessary to write the chemical formula of the precipitate as Mg4Si30~o-x(OH)~X. 32 Solubility constant The dissolution reaction for the magnesium silicate precipitated in this study can be written as:
4Mg2++ 3H4Si04 The amount of water in the magnesium silicate does not affect its solubility in dilute solution (activity of
Figure 2. Solubility constants for magnesium silicate. The line was obtained by regression through the data points.
water = 1). If the activity of the magnesium silicate is taken as one also, the solubility product for its dissolution reaction is given by: logKsp=
[Mg2']4 [H,SiO;l3
were brackets denote activity. Values for the solubility constant of the magnesium silicate were obtained in experiments at 72", 91", 93", 120" and 150°C. In experiments made at 53°C equilibrium was not attained. Figure 2 shows logK,, versus 1000/T (K). The values for the logK,, vary linearly with the reciprocal of the absolute temperature. Therefore extrapolation is safe for obtaining values for logK,, at higher and lower temperature than the actual experimental conditions. The temperature variation of the solubility constant of the magnesium silicate precipitated in this study is described in the equation:
where T is in K. It is valid in the temperature range 0"-200"C at 1 atm below 100°C and at P,,, at higher temperatures. 3.3 Characteristics of the precipitate
Figure 1. Depletion of aqueous magnesium and silica in precipitation experiments.
X-ray diffraction pattern shows the precipitate to be poorly crystalline. No regular peaks were detectable only distinctive swells. The swells do not rise high above the background but they seem to have fairly constant 2 Theta values at 7, 22, 26, 35 and 60 degrees. IR spectrum of the precipitate is shown in Figure 3. The predominant features of the spectrum are a Si-0 stretching bands at 999 and 1080 cm-' ,a single Si-0-Mg band at 450 cm-', Mg vibration at 555 526
Figure 3. IR spectrum of the magnesium silicate precipitated in this study.
cm-' (Farmer 1974), Si-OH deformation band at 895 cm-' (Farmer 1974, Russell, 1987) a single band at 625 cm-', that could be attributed to inner and surface libration of the layers (Pampuch & Ptak 1970), and shoulder at 780 cm" . All these bands are characteristic of antigorite. In a well crystallized antigorite the Si-0 stretching bands at 1080 cm-' and 999 cm-' are better separated and have a shoulder at 970 cm-', but because of the semi-crystalline nature of the sample the shoulder at 970 cm" is not visible.
Figure 4. Magnesium concentration in natural waters in Iceland.
4 SATURATION STATE IN NATURAL WATERS The magnesium concentration of selected Icelandic surface and groundwater is shown in Figure 4. The waters are from basaltic terrain in N-Iceland and can be divided into four groups. The groups are: rivers and streams, peat soil waters, lakes and groundwaters. The magnesium concentration in the rivers and streams, the peat soil waters and the lakes is high, or between 1 and 12 ppm in the peat soil water, 0.2-12 ppm in the rivers and 0.5-4 ppm in the lakes. The magnesium concentration in the groundwaters is not as high, especially in the groundwaters with elevated pH where it can be as low as 0.002 ppm. The saturation state of the magnesium silicate precipitated in this study in these waters is shown in Figure 5 . The saturation state was calculated with the aid of the WATCH chemical speciation program (Arnorsson et. a1 1982). The peat soil waters, rivers and streams and the lakes are generally undersaturated due to the low pH of these waters. Few waters in these groups that are close to saturation or somewhat supersaturated are those with elevated pH value and represent streams with a high groundwater component. The groundwaters are generally close to saturation with respect to the magnesium silicate precipitated in this study. Those below 30°C a little
Figure 5 . Saturation state of the magnesium silicate precipitated in this study in natural waters in Iceland. The symbols are the same as in Figure 4.
bit supersaturated and those above 60°C are somewhat undersaturated 5 CONCLUSIONS
A magnesium silicate, which is poorly crystalline antigorite according to IR spectra, controls Mg concentrations in <5OoC, high pH groundwater in the basaltic formations in Iceland. By contrast, surface 527
waters which typically have pH value between 6 and 8, are undersaturated with this phase. Titration experiments of solutions containing magnesium and silica cause magnesium silicate to readily precipitate when the solution pH has been raised to a sufficiently high level to saturate the solution with this phase. This kind of precipitation occurs when surface water with high magnesium concentration seeps into the bedrock and reacts progressively with it to attain a high pH.
REFERENCES Arn6rsson S. 1995. Geothermal systems in Iceland: Structure and conceptual models-11. Low temperature areas. Geothermics. 24: 603-629. Arndrsson, S., S. Sigur-sson, & H. Svavarsson 1982. The chemistry of geothermal waters in Iceland. I. Calculation of aqueous speciation from 0 to 370OC. Geochim. Cosmochim. Acta. 46: 15 13-1532. Farmer, V. C. 1974. The infrared spectra o f minerals. Mineralogical Society Monograph 4. London: Mineralogical Society, Gislason, S. R. & S. Arn6rsson 1993. Dissolution of primary minerals in natural waters: saturation state and kinetics. Chem. Geol. 105: 117-135. Kunze, G. 1961. Antigorite. Strukturtheoretische grundlagen und ihre praktische bedeutung fur weitere serpentin forschung. Fortschr. Mineral. 39: 206-324. Pampuch, R. & W. Ptak 1970. Vibrational spectra and the structure of laminar silicates of 1:l type. 11. Octahedral layers. Polsku Akud. Nauk, Oddzial Krakowie, Prace Komisij Ceram. 14: 7-36. Russel J. D. 1987. Infrared Methods. In: A Handbook of Determinative Methods in Clay Mineralogy. Editor: M.J. Wilson.,Glasgow and London: Blackie, 133-173. Shitov, V. A. and B. B. Zvyagin 1966. Electron microdiffiaction study of serpentine minerals. Soviet Physics-Cryst. 10: 71 1-716.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Can major ion chemistry be used to estimate groundwater residence time in basaltic aquifers? Andrew L .Herczeg CSIRO Land and Water,Adelaide, South Australia, Australia
ABSTRACT: In tropical environments where there are high water fluxes and high rates of acid production via plant activity (i.e., CO, plus organic acids), most of the dissolved ions in groundwater from basaltic aquifers are derived from weathering of basaltic minerals. The concentration of dissolved constituents is therefore a function of groundwater residence time, volumetric water content (porosity), mineral grain size and weathering rates. Chemical mass balance of groundwaters from the Atherton Tableland basalts region of tropical north-eastern Australia indicate that almost 100% of alkalinity (as HCO,' + organic anions) is produced as a by-product of silicate hydrolysis. Estimates of groundwater residence times based on HC0,concentrations range from 5 - 120 years. Despite large parameter uncertainties, particularly in estimates of mean weathering rates, these values are generally consistent with those obtained independently using environmental tracers such as chlorofluorocarbons. 1 INTRODUCTION
Chemical weathering rates of basalts have been extensively studied in the laboratory and field to develop models of groundwater chemical evolution as well as rates of soil development. Because of the high reactivity of their constituent minerals, the rates of basalt weathering can be high especially in warm tropical areas where there are high rates of soil CO2 and organic acid production. If aquifer parameters such as porosity and mean grain size can be constrained, then we can utilize the knowledge of chemical weathering rates to estimate groundwater flow rates through their major ion composition. This paper explores such an approach in the high rainfall area of tropical northern Australia. The rate at which dissolved solutes in groundwater accumulate in groundwater that drain basaltic rocks will be a function of the weathering rate(s) of the constituent minerals, the porosity, grain size and rate of groundwater flow or mean residence time. Although one can qualitatively relate higher dissolved solute concentrations with longer groundwater residence times, it is more difficult to derive explicit 'ages' primarily because of large uncertainties in estimating weathering rates. Forty eight groundwater samples were collected and analysed for major ions and stable isotopes of water from the Atherton Tableland region located in far north Queensland, Australia. The field site is located about 60 km west of Cairns (17'18's;
145"32'E), on a plateau about 600 m above sea level. Mean annual temperature is about 24°C and a distinct rainfall gradient (south to north) from 2300 down to 980 mm/yr (mostly during November to April). The aquifers are predominantly basaltic, or the weathered equivalent of the original basalts, with fractured metamorphic and granodiorite aquifers in the southern (higher rainfall) part of the study area. 2 RESULTS Groundwaters in the Atherton region are relatively fresh by world standards, ranging from 40 to 250 mg/L. The major dissolved anion is bicarbonate (HCO,*) with slightly higher proportion of sodium than calcium and magnesium. Dissolved silica (H,SiO,) makes up a large fraction (up to 30%) of the total dissolved load. Positive correlations between major cations (Na, Ca, Mg) and alkalinity (as HCO,) are observed for all groundwaters (Fig. 1). There is a tendency for higher concentrations of solutes and alkalinities in the basaltic aquifers than in the metamorphidgranitic aquifers. There is no correlation between major ions and C1 which is typically observed in drier environments where evapo- transpiration dominates major element distributions and salinity. Given that no carbonate minerals occur within the soils and aquifers of the Atherton Tablelands area, the observed increase in HC03- and cations must be caused by reaction 529
Figure 1. Ca, Mg and Na concentrations versus alkalinity (HC03-) for Atherton metamorphic (closed symbols) and basaltic (open symbols) aquifers.
between carbonic acid (H2CO3) and the reactive silicate minerals within the unsaturated and saturated parts of the aquifer. High concentrations of carbonic acid (pCOz's> 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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Wafer-RockInteraction 2001, Cidu (ed.),02001 Swefs & Zeiflinger,Lisse, ISBN 90 2651 824 2
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.
534
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.
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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-
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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
<0.05 5080
(l)DatabyCm1985
-- NA 60
2.2 Tortugas Mountains Hydrothermal Waters The samples considered here are from the two currently active pumping wells and from three nonthermal wells of the adjacent alluvial groundwater Table 2. Hydrochemical data of selected wells in the TMGA. Concentrations are in mg/l. Sample
NMSU-PG-1 NMSU-PG-4 (1)
MAR SW-1OA HTA-3 (1) (1) (1) 25 20.5 19 8.2 8 7.5 10200 339 839
Sodium Calcium Magnesium Potassium Iron Manganese
drinking water, represented by the samples HTA-3 and SW-lOA, is dominated by calcium and bicarbonate and represents a typical dilute and relatively shallow type of groundwater. The other samples present a dominance of sodium, except for WHSAMW3, where calcium is the major cation, and Pz-4, which is dominated by magnesium. On the anion side these samples are clearly dominated by sulfate, except WHSA-MW2, which is Cl-dominated with slightly lower SO4 content (Table 1 and Fig. 2).
(1)
JP4 (1)
0-1
Temperature (C) PH Cond. (pS/cm)
53.2 6.2 2180
52.6 6.47 2330
27.6 7.96 316
24 7.1 2050
NA 7.8 533
Sodium Calcium Magnesium Potassium Iron Manganese
430 165 34.8 51.4 3 0.04
400 167 33.1 54.5 2.7 0.04
58.8 10.4 6.2 2 0.3 0.0009
62 174 74.8 6.8 1.4 0.46
21 69 2.2 NA NA
Chloride Sulfate Nitrate Bicarbonate
880 200 c0.05 510
640 240
8.1 48 0.55 NA
140 150 3.3 NA
13 120 NA 171
0.59
520
(1) Data by Schulze-Makuch & Kennedy 2000 (2) Data by Wilson et al. 1981 NA = not analyzed
= not analyzed
546
WS
(2)
(1)
13
-
system. Representative analyses of the hydrothermal water and the alluvial groundwater are shown in Table 2, where NMSU-PG-1 and 4 are the thermal waters, and JP-4, 0-1 and W-5 were taken from the alluvial aquifer. The water from the hydrothermal wells is more acidic and has a higher amount of TDS compared to the alluvial groundwater. In the Piper diagram (Fig. 2) the hydrothermal water separates clearly from other waters types representing a Namixed anion type of water. W-5, similar to HTA-3 and SW-IOA, is a typical dilute and shallow groundwater. JP-4 and 0-1 indicate a more mixed composition. Schulze-Makuch & Kennedy (2000) described the hydrothermal waters as a mixture of meteoric water from the alluvial groundwater flow system (W-5, as dilute end-member) with higher mineralized anaerobic water. 3 THERMODYNAMICAL RESULTS Saturation index (SI) pattern for both Tularosa Basin and Tortugas Mountains were calculated using the codes SOLMINEQ/GW (Perkins et al. 1996) and “The Geochemist’s Workbench” (Bethke 1992). Activity coefficients for all samples were calculated using the Debye-Huckel model, except for sample GWBP-D, for which Pitzer equations were used due to its high ionic strength (around 3.3). 3.1 Tularosa Basin The thermodynamical calculations reflect well the mineralogical composition of the study area. Supersaturation was mainly found with respect to clay minerals, carbonates and gibbsite. Soil samples from Lake Lucero show for instance the presence of gypsum, dolomite, magnesite and gibbsite and different types of clay minerals. All waters are slightly supersaturated or at equilibrium with respect to calcite and chalcedony, and
Figure 4. Saturation indices (SI) of carbonates of the TMGA as a fhction of the temperature of the fluid. The SI values were calculated on the basis of the solubility products and fluid chemistry of NMSU-PG-4.
clearly supersaturated with respect to dolomite (Fig. 3). Gibbsite is slightly undersaturated in SW-lOA, MAR and PZ-4, in equilibrium in WHSA-MW3 and lightly oversaturated in HTA3, WHSA-MW2 and GWBP-D, whereas magnesite might only form in the middle and high range of TDS. Unexpectedly, Gypsum, the most important mineral at Lake Lucero (Allmendinger 197I), shows a slight undersaturation for all water samples (Fig. 3). 3.2 Tortugas Mountain Geothermal Area The geochemical modeling of the hydrothermal samples was carried out as a cooling path beginning at sampling temperatures (about 5OOC) and ending at 20°C. The results indicate a steady decrease in SI for most calculated carbonate minerals with decreasing temperature as shown in Figure 4. Some other important minerals found were clay minerals (illite,
Figure 3 . SI values vs. TDS of the examined wells in the Tularosa Figure 5. Activity diagram for some carbonate mineral phases Basin. in the study area. For an explanation of the symbols see Fig. 1.
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kaolinite), feldspars (albite, K-feldspar), quartz and chalcedony. Similar to carbonates, they present an increasing SI with decreasing temperature. Comparing the SI values obtained for thermal waters to those SI calculated for other waters, similar SI values were found with respect to dolomite, calcite and magnesite in HTA-3, W-5 and SW-1OA. An activity diagram for carbonates with the calcium activity at the y-axis and magnesium activity at the x-axis is shown in Figure 5. All waters are apparently in equilibrium with respect to dolomite.
4 CONCLUSIONS The Tularosa basin represents geologically a gypsum playa with Lake Lucero in the lowest part of the basin. The Tortugas Mountain Geothermal Area is a convective low-temperature hydrothermal system. The waters of the Tularosa Basin reach from a dilute end-member (e.g. SW-1OA) to the saline endmember (GWBP-D, a typical brine of a sedimentary basin). The origin of the solutes is related to the mineralogical composition of the recharge area with a dominance of carbonates. Even in the discharge area the sediments are still dominated by the same material, therefore the chemical evolution could be at least partly explained by concentration of the residuents by simple evaporation. The waters of the TMGA indicate that the SI values for the carbonates magnesite, calcite and dolomite at W-5 could be explained by a simple cooling of the geothermal water. The chemical and microbiological parameters indicate that the sampled hydrothermal water derives from a mixture of deep anaerobic water with meteoric water from the adjacent alluvial, non-thermal groundwater flow system. The identification of bacteria only present in hydrothermal environments in waters from Lake Lucero indicates a connection between the Tortugas Mountain water and the Tularosa Basin. In the activity diagram for carbonates all waters are lying in the stability field for dolomite. This might underline the existence of a connection between the two systems. Despite the different geological setting of the two hydrogeological systems, several chemical and thermochemical characteristics are similar, as shown by the mineral stabilities or by the saturation indices calculated for calcite, magnesite, and dolomite. All these observations may indicate a possible genetic relationship between both systems. Further detailed studies (e.g. minor and trace elements, and isotopic composition of the different waters) are going on to verify this hypothesis.
REFERENCES Allmendinger, R.J. 1971. Hydrologic control over the origin of gypsum at Lake Lucero, White Sands National Monument, New Mexico. M.S. Thesis. New Mexico Institute of Mining and Technology. Barrol, M.W. & M. Reiter 1990. Analysis of the Socorro hydrogeothermal system, central New Mexico. J. Geophys. Res. 95(B 13): 2 1949-21963. Bethke, C. 1992. "The Geochemist's Workbench" A user's guide to Rxn, Act2, Tact, React and Gtplot.. Urbana: University of Illinois. Cruz, R.R. 1985. Annual Water - Resources review, White Sands Missile Range, New Mexico. US Geological Survey Open- File Report 85-645: p.25. Morgan, P., Harder, V., Swanberg, C.A. & P.H. Dagget 1981. A ground-water convection model for Rio Grande rift geothermal resources. Tram. Geotherm. Recmrr. Council 5 : 193-196. Orr, B.R. & R.G. Myers 1986. Water resources in the basin-fill deposits in the Tularosa Basin, New Mexico. US. Geological Siimey Water-ResourcesInvestigations Report, Vol. 854219, p. 1-94. Perkins, E.H., Gunter, W.D & B. Hitchon 1996. SOLMTNEQ.GW: a user friendly computer code for ground water studies. b i Proceedings of the Fourth International Symposium on Geochemistry of the Earth 's Siq4ace. Ilkley, Yorkshire,England, (editor S.H. Bottrell), p.7-11. Pray, L.C. 1961. Geology of the Sacrament0 Mountain escarpment, Otero County, New Mexico. MBMMR Bulletin 35 144 PP. Schulze-Makuch, D. & J.F. Kennedy 2000. Microbiological and chemical characterization of hydrothermal fluids at Tortugas Mountain Geothermal Area, southern New Mexico, USA. HydrogeologyJournal 8: 295-309. Seager, W. R. 1981.Geology of Organ Mountains and southern San Andres Momtains, New Mexico. "MR Memoirs M-36, 97 pp. Wilson, C.A., White, R.R., Orr, B.R. & R.G. Roybal 1981. Water resources of the Rincon and Mesilla Valleys and adjacent areas, New Mexico. NiWState Eng. Tech. Rep. 43, 5 14pp.
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Water-Rock Interaction 2001, Cidu (ed.), 02007 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Investigation of the carbonate system in Aquifer Storage and Recovery: an isotopic approach C.Le Gal La Salle, J.Vanderzalm & J.Hutson Centrefor Groundwarer Studies and Flinders University of South Australia, Adelaide, South Australia, Australia
P.Dillon & P.Pavelic Centre for Groundwater Studies and CSIRO Land and Water,Adelaide, South Australia, Australia
R. Martin Centrefor Groundwater Studies and Department of Water Resources, South Australia, Australia
ABSTRACT: In Aquifer Storage and Recovery schemes (ASR), the carbon system plays an important role involving processes such as oxidation of organic carbon and carbonate dissolutiordprecipitation that can impact on aquifer matrix integrity and recovered water quality. The aim of this study is to investigate the use of the carbon isotopes to elucidate carbon geochemistry during an ASR trial. Up to 121 ML of reclaimed water was injected into a confined carbonated aquifer over a 12 months period. Observation wells allowed groundwater geochemistry to be monitored. At 4 m from the injection well, calcite dissolution was suggested by a deficit in the carbon-14 signature. After breakthrough at the 50 m observation well, the signature of the carbon isotopes reflected dissolution of carbonate mineral and oxidation of aquifer organic carbon. The investigation is continuing to encompass the end of the injection and the following recovery phase.
1 INTRODUCTION In arid and semi arid areas where surface water is scarce, mining of groundwater resources is an issue that can threaten sustainable development. Aquifer Storage and Recovery (ASR) of surface water or reclaimed-water is becoming a very attractive option for replenishing depleted resources. In addition, aquifer storage can significantly contribute to water quality improvement by natural attenuation and degradation of contaminants. Important issues in ASR are sustainability of injection rates and quality of recovered water. Injection of surface or reclaimed water into an aquifer can induce numerous geochemical and biochemical reactions which may modify aquifer properties and affect the quality of recovered water. Injection of foreign water into a carbonate matrix aquifer may lead to substantial dissolution of the aquifer material (Rattray 1998). While this can offset the effects of clogging during injection, it has the potential to lead to aquifer erosion and borehole collapse. Carbonate dissolution can be further enhanced by microbial activity after injection of organic matter and nutrient rich water. Through degradation of organic matter and respiration, microorganisms produce carbon dioxide that may induce carbonate dissolution. Column studies (Rinck-Pfeiffer et al. 2000) showed that both carbonate dissolution and precipitation occur during injection of reclaimed water. Dissolution occurred at the inlet end of the column
where biological activity was high while calcite precipitation was observed at the outlet end of the column. Precipitation of carbonate can be beneficial if heavy metals are co-precipitated, but may have detrimental effects, such as decreasing the hydraulic conductivity of the aquifer material through physical clogging of pores. Understanding the geochemical reactions that are induced during ASR is essential to improve management practices of such operations. In contrast to column experiments, observation of the aquifer material is not possible during in situ trials. Geochemical reactions can be inferred only through observations of the geochemical signatures of water pumped from observation bores, the number of which is limited. Multiple tracers are required to elucidate the geochemical reactions induced by injection of foreign water to an aquifer between the injection and observation well. A full-scale injection trial using reclaimed water in a confined carbonated aquifer provided an ideal setting for investigating such reactions. The ASR trial is situated in the north Adelaide Plains, South Australia, an area where groundwater resources are mined for irrigation. The aim is to store non saline treated effluent in a slightly brackish aquifer for reuse in summer for irrigation. The effluent has undergone conventional secondary treatment, in addition to storage in polishing lagoons and dissolved air flotation/filtration (DAFF). The 60 to 70 m thick carbonate aquifer is overlain by an approximately 7 m thick clay layer at a depth of 100 m. 549
Carbonate minerals constitute approximately 70 % of the aquifer material. The aim of this study is to i) characterise the geochemistry of the carbonate system and ii) investigate the potential of carbon isotopes to contribute to the understanding of such a system. Use of isotopes has previously been investigated in one study only, in the reuse of storm water for ASR in South Australia (Rattray 1998, Herczeg 2000).
2 METHODS The injection well was approximately 170 m deep with an open interval from 103 to 170 m. Observation wells were placed at 4 m, 50 m, 75 m, 120 m, 300 m and 600 m down gradient from the injection well. The observation wells are open over the full thickness of the aquifer, from 101 to 170 m, while a nest of four piezometers was installed at 50 m. Injection of reclaimed water occurred in three episodes interrupted by redevelopment and/or acidification of the well. Thirty M1 was injected in the first test over a six week period commencing on the 13/10/99. In the second injection phase 7 ML was injected during 17 days in April. The third injection period started in August 2000 and 121 ML was injected by the time of observation. The average rate of injection was approximately 1ML per day. Groundwater samples were collected on a weekly basis after breakthrough started at a specific observation well. Chloride concentrations were analyzed by ICP-MS. The Dissolved Organic Carbon @OC) concentrations were measured by potassium persulphateAJV oxidation with detection of carbon dioxide as methane by FID (Australian Water Quality Center). The total dissolved inorganic carbon was precipitated by addition of barium chloride. Carbon-13 content of the precipitate was measured by mass-spectrometry with an associated uncertainty is 0.2 %o. The carbon-14 samples were prepared by carbosorb adsorption and the carbon-14 activity was measured by p counting (Leaney et al. 1994). The associated error is 1.5 pmc .
the in'ectant concentration ranged from 420 to 520 mg.L'. The fraction of injectant in a sample, f, is:
where C is sample concentration, C,, is mean initial groundwater concentration and Ci,, is the mean injectant concentration. At the 4 meter observation well, breakthrough was detected within the first hour of injection (Fig. 1). Complete breakthrough, or 100% injectant recovery, was achieved after two weeks. The time scale of Figure 1 commences at the start of the first injection period. At the 50 meter observation well, breakthrough was observed at only one depth interval, 134 to 139 m, after 10 months of discontinuous injection. At the time of observation, the 50 m well water quality had been monitored for 5 weeks following commencement of breakthrough and breakthrough was not yet complete. The recovered water at 50 m was approximately 60 % ambient groundwater and 40 % injectant. The dissolved organic carbon was monitored simultaneously. The ambient groundwater had a minimal DOC of 0.06 mmol.L-' with a maximum of 0.24 mmol.L-' at the 50 m well. The injectant had a DOC content of 1.38 mmol.L-', two orders of magnitude higher than the ambient groundwater. At the 4 m well there was initially a net loss of DOC compared to mixing, while the loss decreased in the fourth week of injection. High microbial activity and oxidation of organic carbon in an aerobic environment occurred at the beginning of injection, while microbial activity decreased as the groundwater became anoxic (Vanderzalm et al. 2000). At the 50 meter well, DOC remained relatively low indicating degradation of dissolved organic matter over the monitoring period.
3 RESULTS 3.1 Breakthrough and mixing Breakthrough of the injectant was monitored using chloride concentration, as chloride can be considered as a conservative ion in this environment. The contrasting chloride concentrations in the groundwater and in the injectant allow estimation of mixing between ambient groundwater and the injectant in the recovered water. The ambient groundwater was brackish with a chloride concentration reaching 1 050 mg.L" while
Figure 1. Evolution of chloride concentration as an illustration of injectant breakthrough at the 4 m well and 50 m well.
550
activity ranging from 94.6 to 101.4 pmc while the ambient groundwater in the confined aquifer showed very low concentration of carbon-14 ranging from 2.9 to 10.5 pmc (Fig. 3). At the 4 m well, after breakthrough, the carbon-14 showed a slight decrease in activity compare to the injectant, with activity ranging from 82 to 94 pmc. In contrast, after breakthrough at the 50 meter well, the carbon-14 activity was very low and, as for the carbon-13 signature, fairly similar to the ambient groundwater carbon-14 activities. Figure 2. Carbon-13 of the TDIC versus chloride, showing the ambient groundwater, injectant and water sampled at both the 4 m and 50 m wells.
4 DISCUSSION 4.1 Ambient groundwater and injectant
3.2 Carbon isotopes
The carbon-13 signature of the TDIC of the ambient groundwater can be explained by dissolution of marine carbonate of the aquifer matrix, with CO2 from resident organic matter following the reaction:
3.2.1 Carbon-13 The carbon-13 content of the Total Dissolved Inorganic Carbon (TDIC) is expressed in per mill as a variation from the Pee Dee Belemnite (PDB) standard following the equation below:
H20 + CO2 + CaC03 -+ 2 HC03 + Ca2+
where R = 13C/'2C. The concentration of the ambient groundwater was relatively depleted with 613C around -11.5 %O while the injectant was much more enriched with 613C ranging from -6.8 to -3.8 %O (Fig. 2). At the 4 m well, after total breakthrough, the carbon-13 signature was similar to that of the injectant. However at the 50 m well, the carbon-13 signature was similar to the carbon-13 signature of the ambient groundwater. 3.2.2 Carbon-14 The carbon-14 activity is expressed in percentage of modern carbon (pmc). As expected, the TDIC of the injectant showed a modern carbon-14 signature with
(2)
The marine carbonate signature ranges from -2.1 to +2.7 %O vs PDB (Deines et al. 1974) while CO2 produced by oxidation of C3-type plant material by micro-organisms leads to a carbon-13 signature around -25 %O (Grossman et al. 1989). If 50 % of the TDIC originates from carbonate mineral and the rest from the biogenic carbon dioxide, the final signature of the TDIC is around -12 %o. The carbon-13 signature of the injectant was surprisingly enriched. This might be explained by methanogenesis occurring during secondary treatment of the effluent in the decantation pond following the reaction: 2CH20 -+ C&
+ CO2
(3)
The carbon dioxide produced from this reaction is usually very enriched in carbon-13 while the methane is very depleted (Games & Hay 1976). The enriched C02, up to 18%0(Baedecker and Bach, 1979) is further dissolved and contributes to carbon-13 enrichment of the TDIC. 4.2 Observation wells
Figure 3. Carbon-I4 of the TDIC versus chloride, showing the ambient groundwater, injectant and sampled water at the 4 m and 50 m wells.
At the 4m well, the carbon-14 activity showed a slight deficit from mixing, while no changes were noticeable in carbon- 13. Although dissolution of calcite has been shown in a geochemical investigation (Vanderzalm et al. 2000), this does not totally explained the isotopic signature since it would be expected that carbon- 13 would also be depleted. Over all, mixing with the injectant had a much greater influence on the carbon isotopes signature than the geochemical reaction processes. 551
guishing between oxidation of injectant organic matter and resident aquifer organic matter. Injection is continuing and ongoing monitoring of the tracers will cover the end of injection and the recovery phase.
AKNOWLEDGMENTS
Figure 4. Carbon-14 against carbon-13 for the ambient groundwater, injectant and observation bores after beginning of breakthrough at 4 m and 50 m.
At the 50 meter well, even after 40 % of breakthrough, the carbon isotope signature of the TDIC was similar to that of the background. This shows that dissolution of calcite, using CO2 produced by organic matter oxidation, completely obliterated the TDIC signature of the injectant. In addition, the very low carbon-14 activity shows that the organic matter oxidised has a fairly low carbon-14 activity indicative of old organic matter. This suggests that organic matter from the aquifer was the main source of organic matter used by micro-organisms as against recent organic matter from the injectant. This contrasts with the results of Herczeg (2000) from a nearby site which suggested an important contribution of modern carbon.
5 CONCLUSION At the 4 m well signs of calcite dissolution were shown by a deficit in the recovered carbon-14 signature. However at 4 m the effects of reaction processes on the carbon isotope signature were slight compared to the effect of mixing. The current observations will need to be corroborated with the geochemical investigation. The signature of the carbon isotopes at the 50 m well was totally dominated by dissolution of carbonate mineral associated with aquifer organic carbon oxidation. Together with DOC data, this may indicate that most modem carbon in injectant has been degraded before reaching 50 m during the period investigated. In this trial, carbon isotopes proved valuable and complemented usefully more traditional geochemical tracers. the carbon-13 was found to be quite useful due to the very enriched signature of the injectant, which makes it quite distinctive from the groundwater. This is believed to be linked with the secondary treatment undergone by the effluent before injection. The carbon-14 signature is useful for tracing the origin of the organic matter and distin-
We thanks all the partner organizations involved in the Bolivar Reclaimed Water Aquifer Storage and Recovery Research Project, a collaborative research project of the Centre for Groundwater Studies, CSIRO Land and Water, the Department for Water Resources, South Australian Water Corporation, United Water and the South Australian Government Project Delivery Task Force. We also thank the CSIRO Isotope Analysis Service for their very valuable analytical support.
REFERENCES Deines P., D. Langmuir & R.S. Harmon 1974: Stable carbon isotope ratios and the existence of a gas phase in the evolution of carbonate groundwaters. Geochimica et Cosmochimica Acta, 38: 1147-1164. Games L.M, and J.M. Hay 1976. On the mechanisms of CO2 and CH4 production in natural anaerobic environments. In Environmental Biogeochemistry. Grossman E.L., B.K. Coffman, S.J. Fritz & H. Wada 1989. Bacterial production of methane and its influence on groundwater chemistry in east-central Texas aquifers. Geology 17: 495-499. Herczeg A.L. 2000. Geochemical and isotopic tracers of recharge and reclamation of storrn-water in an urban aquifer: Adelaide, South Australia. IAEA Project Res. Agreement NO. AUL-10063. Leaney, F.W.J., A.L. Herczeg & J.C. Dighton, 1994. New developments in the carbosorb method for C-14 determination. Quater. Geochronol. 13: 171-178. Rink-Pfieffer, S., S. Ragusa, P. Sztajnbok & T. Vandevelde 2000. Interrelationships between biological, chemical, and physical processes as an analog to clogging in aquifer storage and recovery (ASR) Wells. Water Res. 34, 7: 21102118. Vanderzalm J., C. Le Gal La Salle, J. L. Hutson, P. Dillon 2000. Aquifer storage and recovery of reclaimed water in the northern Adelaide Plain: initial geochemical changes close to the injection well, Eere conf. 2000, Victor Harbour, SA. Rattray K. 1998.Geochemical reactions induced in carbonate bearing aquifers through artificial recharge. MSc thesis, The Flinders Uni. of S.A., Adelaide.
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Water-Rock interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrogeochemical evolution of karst water system: A case study at Niangziguan Springs, northern China Yilian Li & Yanxin Wang China University of Geosciences, 430074 Wuhan, P.R. China
ABSTRACT: This paper reports the hydrogeochemical evolution of the Niangziguan karst water system in northern China. The temporal-spatial distribution of travertine and chemical and isotopic compositions of karst waters indicate that the evolution of the karst system can be classified into four stages from middle Pleistocene to the present. In addition, hydrochemical signatures, especially the ratios of components such as Ca/S04, Mg/S04, and SO4/HCO3, are indicative of the controlling geochemical processes such as gypsum, calcite and dolomite dissolution and pyrite oxidation along flow path. The flux and concentration of chemical components in groundwater reached peak in the time period when the I1 terrace was formed, and then continued decreasing to present. Mixing between waters from middle Ordovician aquifer and lower Ordovician aquifer causes decrease in TDS in groundwater from recharge area to discharge area.
1 INTRODUCTION The Niangziguan karst water system, located in Yangquan city, northern China, is the most important water source for potable, industrial and agricultural utilization. Its area covers more than 4600 km2. The spring discharge area is on eastern edge of the system, the lowest part relative to whole area, where 12 springs emerge within 50 km2 as predominant natural discharge of regional groundwater (Figure 1). Among the springs, three have become dry, the others are gradually attenuating in discharge. The springs have existed more than 1.8*104 years"'21, as postulated from thermoluminescence dating in travertine deposited from spring water. Study on hydrogeochemical evolution proves to be usehl for analyzing water-rock reaction processes. In this system there are three lithostratigraphic units: Archaen metamorphic rocks, Paleozoic carbonate, sandstone and shale, and Cenozoic alluvium (clay, sand and gravel). The Paleozoic carbonate formation constitutes the main aquifers. It could be divided into four parts. The lowest part is the Cambrian formation that consists of limestone, dolomitic limestone and dolomite, and the thickness is 120 m, but the lower part in this formation is shale and mudstone that hnction as the regional basal aquitard, The middle parts are Ordovician limestone and dolomite which are two main aquifers contributing water to wells and springs. The limestone aquifer Over dolemite is composed of calcite with three interlayers of
1- Topographic Watershed: 2-Moveable Watershed; 3Watertight Boundary; 4-Fault; 5-Well and Watertable; 6-Groundwater Flow Direction; 7-Artificial Discharge Area; %Natural Discharge Area; 9-Contour of Watefiable; 10-River; River: 12-Karst Spring
Fig 1 Division of groundwater dynamics in Niangziguan KUst system
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The Ordovician limestone outcrops over the whole system, especially at northern, gypsum, and the thickness of limestone is up to 300-600 m, and that of dolomite up to 123-203 m. southern and western areas of recharge. There are many karst fracture, karst cave and conduits developing in it. The upper parts are CarboniferousPermian coal-bearing formations that consist of limestone, shale and coal. Around this area, three sides that are northern, western and southern, are higher in topography. The lowest point is located at eastern side in Niangziguan town where twelve springs issue. Groundwater flow from northern, western and southern sides to center area radially, then discharge on the eastern edge as springs (Fig. 1). According to the dynamics of groundwater movement, the flow system could be divided into two hydrodynamic (1) recharge area (A) and (2) discharge area that it could be further divided into artificial discharge area (B) and natural discharge area (C) (Fig.1). There are many wells in artificial discharge area for pumping water. The natural way of discharge is in the form of springs, and almost all springs emerge from Ordovician dolomite. 2 PALEOHYDROGEOCHEMICAL EVOLUTION 2.1 Paleohydrogeofogical evidence: travertine
All of the springs are sited on banks of the Mianhe river within a range of 7 km in length and discharge into river. Four terraces develop on both banks of the Mianhe river, designated as I, 11, I11 and IV terrace respectively away from the banks: the IV terrace is highest and I lowest. No travertine was found on the IV terrace with palaeosol, clay and gravel. Very limited amount of travertine is located at the upper part and the bottom of the 111 terrace that is mainly composed of gravel. Travertine at the upper part of the I11 terrace is found just around the boundary between Middle Ordovician limestone and lower Ordovician dolomite. A huge mass of travertine (more than 40 m thick at some places) develops on the I1 terrace with interlayers of sand and clay. Very limited amount of travertine continues forming on the I terrace consisting of sand and gravel, and most of springs are distributed on this terrace. Except I terrace, there are no springs issuing presently in the other terraces.
vertine to have been preserved. Furthermore, the structure and lithology of IV terrace are similar to I11 terrace except that there is no travertine on the IV terrace. If groundwater discharged as spring at this area, there should have been some travertine preserved. So before I11 terrace was formed, groundwater discharged as subsurface flow to the east of Niangziguan or very few springs discharged very little local groundwater that was undersaturated with respect to calcite in early-middle Pleistocene period. (2) The initial period of the springs discharge. Travertine is found on I11 terrace, implying that springs occurred at the time when I11 terrace was formed. The site of travertine found on the upper part of I11 terrace is just around the division between Middle Ordovician limestone and lower Ordovician dolomite. This strongly suggests that occurrence of the springs resulted because lower Ordovician dolomite functioned as an aquitard relative to Middle Ordovician limestone and retarded regional groundwater flow, when the bed rock was eroded by the river down to the division between limestone and dolomite. The small quantity of travertine formed in this period indicates small discharge and small number of springs, and the spring water was mostly from local flow system. (3) The culmination period of spring development. The huge mass of travertine on I1 terrace reflects the great number of springs. As the bed rock was eroded deeper by river, the springs discharged more waters from intermediate flow system, i.e., groundwaters flowing over a long distance. The change in travertine altitude and mass indicates that springs moved down and their discharges greatly increased, and even more groundwater dissolved much more minerals and had been relatively saturated with calcite. (4) The spring discharge attenuation period. As time went by, river erosion went on, and the locations of the springs continued changing both laterally and vertically (downward), but very little travertine was precipitated. As a result, the regional groundwater table drastically fell. In recent years, largescale abstraction of karst water results in a much faster rate of attenuation of spring discharges, some springs even ceasing outflow. 3 RECENT CHANGES IN HYDROCHEMISTRY
2.2 P a l ~ o h y d r o x ~ o lreconstruction c~l Variation in altitude at which travertine had been deposited implies that springs should have emerged at the sites during geological history. Considering change in height and amount of travertine, four stages springs development could be inferred: (1) No spring or very little spring period. Since no travertine occurs on IV terrace, it is reasonable to postulate that no spring issued at this altitude or very little spring water that did not deposit enough tra-
Over 500 water samples in different hydrodynamic areas have been collected by pumping from wells. Actually, most water samples are mixtures from different aquifers in wells and springs. Although it is difficult to know the change in hydrochemistry in single aquifers, the change in regional hydrochemistry could still be inferred. And variation in component in spring water in Niangziguan could reflect the change in water-rock reaction in this karst water system.
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taining gypsum. If more water migrated from Middle Ordovician aquifer to well, it would be high in sulfate and calcium. High calcium and sulfate concentrations in water in recharge area(A) simply result fiom faster dissolving gypsum in Middle Ordovician aquifer which is widely distributed over most recharge areas. ( 2 ) Low sulfate, calcium and TDS in groundwater were due to intrusion of waters from lower Ordovician aquifer that consists of dolomite to wells. But magnesium is relatively high in groundwater. (3) Pyrite oxidation and acid mine drainage caused extremely high sulfate and low pH in water. Fortunately, it does not seep into groundwater in the whole system, only in very limited local areas. (4) Mixing water flowing from recharge area with that from lower Ordovician aquifer should be the main process controlling the decrease in total dissolved solids from recharge area to discharge area. The amount of water yield from lower Ordovician aquifer is gradually increasing. The proportion for mixing is quite stable in spring water, since the chemical compositions of most springs water is relatively stable in recent several decade.
Table 1. Average chemistry of groundwaters in different hydrodynamic areas (A. B. and C as in Fig.1) of the Niangziguan karst system (in mgL) Region Recharge area(A) Artificial dischargearea(B) Natural dischargearea(C)
Ca Mg Na K
HC03 C1 SO4 TDS
154 32 59 1.58 275 44 311 800 124 29 12 1.93 238 45 225 660
118 33 35 1.84 253 49 213 670
3.1 Regional hydrochemistry The average chemical components calculated from over 500 water samples in different hydrodynamic areas are shown in Table 1. From recharge area to discharge area, the evident feature is that almost all ions decrease in concentration except potassium and chloride. Also the water type changes from SOe * HC03-Ca water to HCO; SOc-Ca(Mg) water. There are so small differences between artificial discharge area and natural discharge area that total dissolved solids(TDS) are almost the same; but the difference between recharge area and discharge area is greater (130 mg/L ). The reasons for such hydrochemical evolution could be related to two geochemical processes: ( 1) oversaturation with respect to solid phases (such as gypsum, calcite and dolomite), resulting in precipitation of some minerals along flow path from recharge to discharge area; and (2) mixing of waters from recharge area with other lower concentration waters. Ratios between the major components are shown in Table 2. Ratios of (Ca+Mg)/SOc and Ca/SOe increase from recharge(A) to discharge(B,C) area, and C a M g and S04/HC03 decrease along flow path. These indicate that it is impossible for gypsum, calcite and dolomite precipitation, but the variation of Ca/Mg shows that dolomite may dissolve or other high magnesium waters may have been added. The high TDS waters are lower in (Ca+Mg)/SOc and CdSOe and higher in SOs/I-ICO3, and it indicates that it is likely for acid mine drainage (oxidation of pyrite in upper coal-bearing formation) seeping into deep water. Hydro eochemical modeling based on PHREEQC m ~ d e l ' ~ ~was ~ ' ~made ~ ] considering the above characteristics of water chemistry. The following ideas are obtained from the modeling results. (1) High TDS in groundwater is the result of gypsum dissolution in Middle Ordovician aquifer con-
3.2 Evolution in spring hydrochemistry Among the springs, some show great change in hydrochemistry with time. The change in sulfate concentration of five springs is given in Table 3. Of the five springs, the highest altitude at which spring emerges is Chengjia spring, second is Shuiliandong spring and Chengxi spring, the lowest is Wulong spring and Weizeguan spring. Judging from Table 3, the higher the altitude of springs is, the greater the change in sulfate. For the Chengjia spring, the concentration of sulfate increased five times from 1978 to 1987 and 12 times to 1990. After that Chengjia spring has been dry. On the other hand, the lower the altitude of spring is, the smaller the change in sulfate concentration. The Wulong spring and Weizeguan spring have stable chemistry, although the flux is slowly decreasing. The second highest spring, Shuiliandong spring, although very little change in sulfate, but ceases outflow from time to time. According to these variation pattern, it could be reasonably inferred that Chengxi Table 3. Variation in sulfate concentration of some springs. Year
Table 2 The ratios between master species in groundwater in Niangziguan karst system
Chengjia 13
24
188
207
202
1987
67
127
207
214
208
135
193
225
223
151
172
208
225
230
1992
164
227
229
241
1993 1994
185 172
219 192
238 222
227 192
1990
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Wulong Shuilian Weize dong guan
1978 1988
Region (Ca+Mg)/SO.l C d S O , CaMg S04/HC03 1.20 2.92 1.43 1.60 A 1.32 2.59 1.20 B 1.83 1.32 2.12 1.07 1.94 C *S04/HC03 ratio in meqL and others in mol/L
Chengxi
spring, which now change rapidly in sulfate and is third highest altitude, will be ceasing outflow as Shuiliandong spring did. The changes in hydrochemical composition and flux of springs indicate that spring at higher altitude mainly consist of water from shallow local flow system and very little from intermediate flow system. In contrast, waters of the lower altitude springs are from regional flow system and little intermediate flow system. The shallow local groundwaters are low in sulfate; the medium-depth groundwaters which comes from Middle Ordovician limestone have highest sulfate concentration; and the deep water that comes from lower Ordovician dolomite is lowest in sulfate.
Plummer, L.N. 1992. Geochemical modeling of water-rock interaction: Past, present, future. Proc. WRI-7, Kharaka & Maest (eds): 23-33. Rotterdam: Balkema. Plummer. L.N. & W. Back, 1980. The mass balance approach: Applications to interpreting the chemical evolution of hydrologic system. Am. J. Sci. 280: 130-142. Plummer. L.N., Busby.J.F & Lee R.W, 1990. Geochemical Modeling of the Madison aquifer in Parts of Montana, Wyoming, and South Dokota. Water Resource Research 9: 198 1-20 14. Plummer, L.N., Parkhurst D.L. & D.C. Thorstenson, 1983. Development of reaction models for ground water systems. Geochim. Cosmochim. Acta 47: 665-685.
4 CONCLUSIONS 1. Change in the amount of travertine shows that mineral dissolution in groundwater had undergone increasing period and came up to highest in the time period when the I1 terrace was formed, then decreasing period until present. 2. From recharge area to discharge area, more and more water migrates from lower Ordovician aquifer into wells and mixes with water from Middle Ordovician aquifer. It causes decrease in TDS along flow path. 3. The total flux of springs decreases continuously, and more and more springs will be drying. But water quality of some springs is turning relatively better. 4. Pyrite oxidation and acid mine drainage is another source for the high sulfate in groundwater.
ACKNOWLEDGEMENTS T h e research work w a s financially supported by the Natural Science Foundation of China (Grant No.49832005) and the Ministry of Science and Technology of China (Grant No.95-pre-39).
REFERENCES Li, Y. 2000. Geochemical Evolution and hydrochemical modeling of the Niangziguan karst water system, Shanxi, China. Ph. D Thesis, China University of Geosciences, China. Li Y., Wang, Y. & A. Deng 2001. Paleoclimate Record and Paleohydrogeological Analysis of Travertine from the Niangziguan Karst Springs, Northern China. Science in China (E) (In Press). Li, Y. Wang, Y. & H. Gao 2000. Hydrogeochemical Implications of Karst Water F1ow:An Case Study at Niangziguan, Northern China. Proceedings of the International Symposium on Hydrogeology and the Enviroment: 207-2 1 1. Parkhurst, D.L 1997. Geochemical mole-balance modeling with uncertain data. Water Resources Research 33(8): 19571970.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Exchange of solutes between primary and secondary porosity in a fractured rock aquifer induced by a change in land-use Andrew. J .Love Department for Water Resources, Adelaide, South Australia, Australia
Andrew. L.Herczeg CSIRO Land and Water,Adelaide, South Australia, Australia
ABSTRACT: The chemical composition of groundwater can be affected by altered land-use with consequent changes in evapotranspiration rates. Groundwater chemistry and environmental tracers in 10 nested piezometers up to 100 m below the water table in a fractured dolomite in a sub humid catchment in South Australia show a continuum of concentrations as a function of depth. Relatively shallow groundwaters (< 40 m) have low salinity (< 1500 mg/L) are relatively young (14C concentrations > % 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 (<2 % MC) at 100 meters below the water table. The chemical and isotope data suggests two end members with a continuum of concentrations
3.2 Origin ofthe end members
Figure 2. (a)-Q Major ion data versus C1, the solid line represents the ion/C1 sea water dilution line: (8) 6% versus 6I8O, MWL= Meteoric Water Line ; (h) I3C versus I4C.
The geochemical, carbon isotopes and stable isotopes of the water molecule provide evidence for mixing between two end members (see above). For C1 concentrations < 13 mmol/L, the NdC1 and SOJCl marine ratio is preserved. The C l B r mole ratios range from 500 to 700 indicating preservation of the marine aerosol ratio (632) with no dissolution of evaporites in the groundwater system (Br data not presented). This indicates that at low chloride concentrations (ie, in the top 40 m) that Na, SO4 and Br behave conservatively relative to their marine aerosol inputs. For the total range in C1 concentrations the WCl ratios are all above the seawater dilution line suggesting addition since recharge. The Mg/CI and CdC1 ratios are above the sea water dilution line, however the Mg/Ca ratio ranges between 3-4 for all concentration ranges close to the marine aerosol input of 5. This may suggest that the input ratio has been preserved via evaporation prior to recharge or alternatively the molar ratio has been quasi preserved despite possible waterhock interactions. For C1 > 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 (<40 m) end member is controlled by concentration of cyclic salts via evapotranspiration. ET as well as dissolution of the dolomite via either de- dolomization or pyrite dissolution paths control elevated solute concentrations for the deep groundwater end member.
The decrease in pH results in waters slightly undersaturated with respect to dolomite, which causes dolomite to dissolve via reaction with CO2 which results in an increase in Mg in solution according to 2: 2C02 + CaMg(CO3)z + 2H20 = = Ca2++ Mg2++ 4H20
REFERENCES Allison, G.B., Cook, P.G., Barnett, S.R., Walker, G.R., Jolly, 1.D & M.W Hughes 1990. Land clearance and river salinisation in the Western Murray Basin, Australia. J. Hydrol. 119:1-20. Hendry, M.J., & L.1 Wassenaar 2000. Controls on the distribution of major ions in a thick surficial aquifer. Water Resour. Res. 36(2):503-513. Love, A.J., Cook, P.G., Herczeg, A.L., & C.T. Simmons 1999. Environmental tracers as indicators of groundwater flow and evolution in a fractured rock aquifer, Clare Valley, South Australia. In: Isotope Techniques in Water Resources Management: IAEA, Vienna 10-14
(2)
For reaction 2 to proceed there needs to be an additional source of C02, which could be provided in this anoxic environment via microbial oxidation of organic matter. If this scenario is correct we would expect that the equivalent ratio of (Ca + Mg)/HC03 would approach 2, as is observed in Figure 3b. Under this scenario the 613C of the DIC would trend towards equilibrium with the dolomite. As an alternative, scenario 2 invokes dissolution of dolomite via reaction with protons as a result of the dissolution of reduced sulfur minerals (e.g. pyrite) according to reaction 3:
-
FeS2 + 14Fe3++ 8H2O = = 15Fe2+ + 2S04 + 16H'
(31
Figure 3. (a) Ca versus SO4 (b) HC03 versus Ca+Mg
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Water-Rock Interaction 2001, Cidu (ed.), 0 2001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Redox chemistry of a river-recharged aquifer in the “Oderbruch” region in eastern Germany G .Massmann & A .Pekdeger Department of Geosciences, Free University of Berlin, Germany
C .Merz Centre for Agricultural Landscape and Land Use Research, Miincheberg, Germany
M .-Th.Schafmeister Institute for Geological Sciences, Ernst-Moritz-Arndt-University Greifswald, Germany
ABSTRACT: The “Oderbruch”, Germany’s largest polder region, has been artificially drained during the past 250 years. The steep hydraulic gradient between the river Oder and the aquifer results in the lateral infiltration of river-water into the shallow aquifer. A number of redox-controlled, chemical changes occur when oxic riverwater enters into the anoxic environment where the water is progressively reduced. The aim of the study is to develop a detailed understanding of both hydraulic and geochemical processes taking place during bankfiltration. The paper focuses on the processes that control migration, dissolution and precipitation of the dominating iron and manganese compounds. 1 INTRODUCTION Redox processes play an important role in natural aquatic systems. They control the distribution and migration of redox sensitive species such as 0 2 , Fe2+, N03, H2S, Mn2’ or C&. It is generally recognized that redox processes proceed fi-om the highest energy yield downwards, resulting in a characteristic sequence of reduction processes. A common scenario is the infiltration of oxygenated river water into an organic rich aquifer, a situation which is also present in the Oderbruch Region. The objective of the still ongoing study is to describe the complex infiltration process in the vicinity of the dike at an exemplary field site (Bahnbrucke) in the Oderbruch Region. The aim is to identlfl redox processes along individual flow-paths. Modeling of transport and reaction of the affected species is currently in process. Former studies in the region concentrated on the shallow, near surface groundwater only (e.g. Kofod et a1.1997) and suggested that the deeper groundwater has a quantitatively substantial influence on the overall turnover rates. Therefore, the main focus of the current study is on the deeper infiltration paths of the bank-filtrate.
polder. It is located in eastern Brandenburg at the border to Poland, about 50 km north-east of Berlin. The region has been artificially drained during the past 250 years. Melioration measures included the redirection and diking of the river Oder as well as the installation of an extensive drainage ditch and pumping station system. Today, as a result of the drainage activities, a major part of the region lies below the river water table and only the dike prevents the agriculturally intensively used area fiom flooding. The steep hydraulic gradient between the water table of the Oder and the aquifer results in permanent lateral infiltration of river water into the shallow, confined aquifer. The Oderbruch has been chosen by the German Research Foundation (DFG) for the study of redox pro-
2 SITE DESCRIPTION 2.1 The Oderbruch Polder Region With an area of approximately 800 kmz, the Oderbruch represents Germany’s largest enclosed river
Figure 1. Location of the Oderbruch river polder including the field-site “Bahnbrucke”.
561
ter sampling. Water samples were analysed for Ca, Mg, K, Na, C1, NO3, Br, SO4, NO2 by ion chromatography (Dionex DX 500), for Mn and Fe by ICP (Jobinyvon), for NH4, PO4, S (Uvicon 931 Kontron) and SiOZ (Dr 2000 Hach) by photometry and for DOC by a DOC-Analyser (Shimadzu).
4 RESULTS 4.1 Hydrogeology
Figure 2. Schematic “Bahnbrucke”.
map
of
the
piezometer
The 20-22 m deep aquifer consists of glacio-fluvial sands and gravel. It is sandwiched by a boulder mar1 at the base and an alluvial loam at the top. The upper parts of the aquifer are mainly fine to medium sized sands with hydraulic conductivities ranging between 5.5”lO-j to 4.0*104 d s , getting larger towards the deeper parts (hydraulic conductivities of 2.9-8.9* 1O4 d s ) . Due to the strong current, the river base is highly permeable and hydraulic contact between river and groundwater is unrestrained (hydraulic conductivities between 9*104 and 1.5*10” d s ) . Mean flow velocities vary between 0.5 - 1.5 &day close to the river declining to 0.2 - 0.3 d d a y further inland. Figure 3 shows the schematic flow behavior along a projected line from the piezometer transect near the river to piezometers 956019561 at 620 m distance. For the construction of the cross-section water levels at low-water on Sept. 28, 1999 were used. The main geological features are included in the picture. Even at low river water level, a major part of the percolating water is discharged by the main drainage ditch running parallel to the dike, resulting in a complicated flow pattern with a strong vertical component. The following descriptions refer to the deeper groundwater paths only (Fig. 3). These are more representative for the polder as a whole since they reach further beyond the drainage ditch.
field
cesses in natural aquatic systems within the priority program “geochemical processes with long-term consequences in anthropogenically affected seepage water and groundwater”. 2.2 Field-Site Bahnbriicke A 1 kmz exemplary field-site “Bahnbrucke” (Fig. 2), located adjacent to the river Oder in the northern Oderbruch, was chosen for detailed investigations. 24 conventional piezometers at varying depths plus 3 multilevel wells were constructed at a distance of 3 to 620 m from the river. The area is intensively surveyed and monitored for relevant hydraulic and hydrochemical parameters. Major hydraulic features in the area are shown in Figure 2: The river Oder with the adjacent dike and the major drainage ditch both running parallel to the dike.
4.2 Water chemistry During bank-filtration of oxic Oder river water a sequence of redox reactions can be observed with in-
3 METHODS Surface- and groundwater samples were collected every 2 months over a period of 1 year. Measurements of Eh, pH. 0 2 , temperature and conductivity were carried out in the field in a flow cell. Filtration with 0.45 pm membrane filters was done immediately after sample retrieval to prevent iron and manganese precipitation. Samples for cation analysis were preserved with nitric acid. Alkalinity was titrated aRer lab return (Titrolein 96). Full analysis was generally performed one day af-
Figure 3. Cross-section through the aquifer, Sept. 28, 1999.
562
0 - P
Figure 5. Spatial concentration of Mn2+in piezometers with filter levels between 12 and 18 m below sea level. Interpolation performed with Surfer (Golden Software, Inc. 1997). Data points used for kriging shown as black dots.
150 1
mg/l reflects the river water concentration. S2- was not detected. According to the current state of the observation campaign the described behaviour of all redox components is stationary. Calculations were performed with PHREEQC for Windows 1.0 (Parkhurst & Appelo 2000) for all sampling points. A selection of saturation indices (SZ for relevant minerals plus the corresponding Mnl
5 0 & 1 , ' , I I , , I , , 1 i 0 100 200 250 300 350 400 450 5W 550 600 650 dike drain
distance from river Oder [m]
Figure 4. Concentrations of major redox species with increasing distance of the river. Only piezometers with filter-levels between 12 and 18 m below sealevel are shown.
creasing travel and reaction time of the infiltrating water. Selected data of one representative sampling campaign (January 2000) is shown in Figure 4. Of the electron acceptors present in the river water, 0 2 and N03- are consumed within the first few meters or even decimeters of underground passage and cannot be found in any of the observation wells. As a result of the commencing Mn-(hydr)oxide reduction, the Mn2' content of the water constantly increases up to a distance of 150 m from the Oder where it rises to a distinct maximum of 5-6 mg/l Mn2'. After reaching the peak concentration, the Mn2" content decreases down to concentrations below 1 mg/l at a distance of 620 m inland. Alkalinity and pH both show a similar maximum at 150 m river distance. The Fe2' concentration of the groundwater increases just after Mn2' does and continues to do so until 620 m distance fEom the river where it reaches the highest observed concentrations of 2-5 mg/l Fe2', which is still below the peak concentrations of Mn2'. The redox environment can be classified as postoxic after Berner (1981). The water does not pass the ferrous zone, and S042-reduction does not commence. The S042-content of the groundwater of about 100
563
Figure 6. Spatial concentration of Fe2':Interpolation performed with Surfer (Golden Software, Inc. 1997). Data points used for kriging shown as black dots.
pass the postoxic, ferrous zone. Sulfate reductior was not observed. REFERENCES
Figure 7. Saturation indices of Mn-carbonates and -hydroxides in 40, 140, 350,620 and 5000 m fiom the river.
content of the sample is shown in Figure 7. In contrast to foregoing Figures an additional groundwater sample in 5 km river distance is included. Various Mn-(hydr)oxides are potential sources for Mn?' in solution since they are subsaturated without exception (Fig. 7). The actual concentration of dissolved Mn2' is, however, controlled by the saturation state of Mn-carbonates (crystalline or amorphous). Rhodochrosite is considered to be the most common reduced Mn(I1)-mineral in nature (Berner 1980). It precipitates in aquatic system and thus removes Mn" fiom the groundwater according to the following reaction:
Berner, R. A. 1980. Early diagenesis - A Theoretical approach. Princeton: Princeton University Press Berner, R.A. 1981. A new geochemical classification oj sedimentary environments. J Sed. Petrol. 5 1,2: 359365 Golden Software, Inc. 1997. Surfer Version 6.04. A Surface Mapping System Kofod, M., J. Schuring, J., C. Merz, A. Winkler, T. Liedholz, I. Siekmann & M. Isenbeck-Schroter 1997. Der geochemische EinfluD des Sickenvassers aus landwirtschaftlich genutzten Flachen auf das Grundwasser im Oderbruch Matsunaga, T., G. Karametaxas, H.R. von Gunten & P.C. Kichner 1993. Redox chemistry of iron and manganese minerals in river-recharged aquifers: A model interpretation of a column experiment. Geochim. Cosmochim. Acta 57: 1691- 1704 Parkhurst, D.L. & C.A.J. Appelo 2000. User's Guide to PHREEQC (Version 2)--A Computer Program for Speciation, Batch-Reaction, One-Dimensional Transport, and Inverse Geochemical Calculations. US Geol. Surv. Water Resour.
~ 2+ HCO~' =M ~ c + o H~'. ~ It appears that the precipitation of rhodochrosite is encouraged when the dissolution of Mn- and Fe(hydr)oxides produces a more alkaline and thus C03" richer solution. Similar observations have been made by Matsunaga et al. (1993). Although carbonates are supersaturated up to 350 m distance inland they do not precipitate until a gradient emerges and SI approaches 0. Further inland in the central Oderbruch, Rhodochrosite is subsaturated and Mn2' concentration of the groundwater therefore rises again. This is due to the lower pH of approximately 7 in this area in comparison to about 8 near the Oder. 5 CONCLUSIONS A sequence of redox reactions can be observed during infiltration of oxygenated river water into a confined aquifer in the northern Oderbruch, Germany. The river water is freed of its 0 2 and N03' content before reaching the fist observation well. Mn2' is released by means of Mn-(hydr)oxide reduction. The actual concentration of Mn2' is controlled by the saturation state of Mn-carbonate (Rhodochrosite). Fe2' content of the groundwater rises with increasing distance throughout the observed area. Up to a distance of 620 m from the Oder, the water does not 564
Water-Rock lnteracfion 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The origin of Na-HC03 type groundwater in an eastern section of the Lower Namoi River catchment, New South Wales, Australia W.McLean & J.Jankowski UNSW Groundwater Centre, School of Geology, University of New South Wales, Sydney, N.S. W., Australia
N .Lavitt Centre for Water and Waste Technology, University of New South Wales, Sydney, N.S. W., Australia
ABSTRACT: Environmental stable isotopes and hydrogeochemical data were used to study the heterogeneous alluvial aquifer system of the Lower Namoi River catchment. This study has identified carbon sources in the groundwater system and highlighted mixing relationships between groundwaters from bedrock, Great Artesian Basin (GAB) and alluvial and surface water recharge. In the southeast of the catchment, data indicates that Na-HC03 groundwaters f7om alluvial aquifers are of meteoric origin. However, isotopically heavy 613C signatures for groundwaters from the middle alluvial aquifer are similar to GAB waters and indicate that mixing is occurring between GAB and alluvial groundwaters. A 613C signature of -8.9%0 for the GAB groundwaters may be attributed to an external source of mantle COz. Na-HC03 and mixed cation-HC03 groundwaters from alluvial aquifers have lighter 613C values of -15 to -1 1%0and are consistent with chemical evidence of genesis via carbonate mineral dissolution by groundwater charged with CO;! generated in the soil zone via a photosynthetic pathway,
1 INTRODUCTION
2 GEOLOGY
The Lower Namoi River catchment is a regional groundwater system covering 5 100 km2. These groundwaters evolve from Na-HC03 to Na-HC03-CI and Na-CI chemical type water along potential flow paths. The system has been heavily pumped over the last 30 years for irrigation related to cotton farming. To maintain the sustainability of this system it is vital to: 1) understand the hydraulic connection between the three alluvial aquifer systems and the geochemical and mixing processes occurring; and 2) identify and understand recharge processes such as the impact of groundwater leaking upward leaking groundwaters from the GAB. Numerous hydrogeochemical and isotopic investigations have been carried out in the GAB (Herczeg et al. 1991, Calf & Habermehl 1984). Na-HC03 groundwaters from the Jurassic units of the GAB in the Southeastern part of the Lower Namoi River catchment (the principal recharge zone of the GAB) were found to have 613C values of -7%o which are heavier than other 613C values in the GAB (Calf & Habermehl 1984, Calf 1978). This paper will discuss the origin of Na-HC03 groundwaters in the southeastern portion of the alluvial aquifer system and use stable isotopes to examine the role of mixing between surface (river) water, alluvial and GAB groundwaters.
In the Lower Namoi River catchment Miocene to modern alluvial sediments lie unconformably on the Jurassic to Cretaceous bedrock units of the GAB. These sediments consist of terrestrial sandstones, siltstones and mudstones of shallow marine origin. The Jurassic Pilliga Sandstone outcrops to the east of the study area and forms the main recharge zone for the GAB (Calf & Habermehl 1984). During the Oligocene-Miocene (Wellman & McDougall 1974) volcanic activity, related to the northerly movement of the Australian continent over one or several hotspots, occurred in the area. This formed the Nandewar Igneous Complex and its associated lava flows. Alluviation associated with the Namoi River system commenced in the late Miocene forming three alluvial units that comprise the main aquifer systems in the study area based on age and compositional differences (Williams 1993). The lowermost unit, the Cubbaroo Formation, is restricted to the main Namoi palaeochannel that runs north-northwest of the present river course. It consists of quartz sand and gravel supported by a clay matrix with minor carbonaceous stringers. The middle Plio-Pleistocene Gunnedah Formation subcrops over the entire catchment and in some areas reaches a thickness of up to 80 m. It consists of moderately well sorted sand and gravel interbedded with pre565
Figure 1. Typical cross section (Gurleigh section) showing the three alluvial formations. 1 : Narrabri Formation, 2: Gunnedah Formation and 3: Cubbaroo Formation. Figure 2. Bivariate plot of Na + HC03 (% meq/l) v EC.
dominantly brown to yellow clay. The uppermost, 10 to 40 m thick, Pleisto-Holocene Narrabri Formation conformably overlies the Gunnedah Formation and forms the ground surface in the majority of the catchment. The formation consists predominantly of brown clay with minor sand and gravel with clays dominated by montmorillonite and sporadic calcareous and ferruginous nodules. Figure 1 represents the geology of a typical cross section (the Gurleigh Section) in the southeast of the catchment and the focus area for this study. The three alluvial formations are clearly identified along with the main Namoi palaeochannel and bedrock high separating the modern river channel from the palaeochannel. 3 HYDROGEOLOGY The Namoi River catchment may be considered as; three aquifers, coincident with the geological forma- . tions described above. The main productive aquifer: ; are associated with the middle Gunnedah and lowei Cubbaroo Formations. Originally these aquifer5 ; were pressurized, however, due to heavy abstrac- . tions over the previous 30 years these aquifers have: become dewatered in certain areas of the catchment. Consequently, this heavy pumping has resulted ir i significant changes to lateral and vertical hydraulic : gradients. Potential hydraulic gradients indicate flow in s 1 northwesterly to westerly direction away from thc ; Namoi River. The groundwater system is mainly r e charged in the eastern part of the catchment from tht : Namoi River and its tributaries with diffuse rechargt : believed to occur during major flooding events. Other contributions come from diffuse infiltration o f rainfall, leakage from GAB, inflow from the neigh, bouring catchments to the southeast and northwes t and on-farm water losses.
4 GROUNDWATER CHEMISTRY Two main hydrogeochemical groups characterise the alluvial aquifers in the Gurleigh section and are rep-
resentative of groundwaters in the eastern portion of the catchment. Na-HC03 groundwaters are dominant in the Cubbaroo aquifer and those parts of the Gunnedah and Narrabri aquifers not under direct influence of river recharge (Fig. 2). The majority of other alluvial groundwaters are dominated by HC03 but have a mixed cation composition due to the effect of recharge from the Namoi River. In some samples from the Narrabri Formation C1 increases in the shallow groundwater system. McLean et a1 (2000) discuss the hydrogeochemical characterisation of the entire catchment. GAB groundwaters sampled in the study area are Na-HC03 in type, and are mainly 90% pure NaHC03 (Fig. 2). They commonly have pH values ranging from 7.1 to 8 and are strongly reducing (Eh values ranging from -200 to -320 mV). EC values range from 1050 to 1250 pS/cm. In a study of the evolution of groundwater for the entire GAB. Habermehl (1983), concluded that the chemical evolution of Na-HC03 groundwaters of the Lower Cretaceous and Jurassic sequences of the GAB in this eastern recharge zone probably involved ion exchange. Herczeg et a1 (1991) concluded that the NaHC03 groundwaters in the eastern recharge zone of the GAB to the north of this study area are generated by the dissolution of carbonates and silicates by CO2 generated in the soil zone, followed by the exchange of Ca2+ and Mg2+ for Na+ bound to clay minerals, thus introducing Na' ions into solution. These processes are also likely to occur in the study area. In the alluvial aquifers, similar reactions are involved in the generation of Na-HC03 groundwater, where groundwater charged with either soil or biogenic CO2 dissolves carbonate minerals present in the soil zone, and upper and middle aquifers releasing Ca and Mg into solution. In addition, weathering of silicates releases K, Na and HC03 into solution at a slower rate than carbonate reactions. Albite is the dominant feldspar in the system (Williams 1993)'and Na is contributed by its incongruent dissolution to kaolinite or montmorillonite. The dissolution of albite releases Na+ and HCO3- into solution
566
nate groundwater rich in C02. GAB groundwater has a ratio of 1:1, showing that both total alkalinity and DIC dominated by HCO3-. Groundwaters mixing with GAB groundwater have ratios slightly greater than 1. All other groundwaters have ratios much greater than 1. In some shallow groundwaters there is an increase in DIC without an increase in total alkalinity. This is due to the fact that these are modern groundwaters that have had minimal contact with carbonate minerals. Figure 3. Ca+Mg vs. S04+HC03 indicating ion exchange.
5 STABLE ISOTOPES
in an approximate ratio of 1:l. Ratios of NdHC03 commonly exceed 1 and are as high as 28, thus indicating that either Na' is being added to the system or HCO3 is removed. Given that saturation indices are undersaturated with respect to calcite and dolomite it is more likely that Na' is being added to the system via an ion exchange mechanism. Figure 3 indicates that ion exchange is a process occurring in the system. Groundwaters plotting below the 1:1 dissolution line are undergoing ion exchange with Ca+Mg depleted in comparison to S04+HC03. These processes are also occurring in HCO3- groundwater with mixed cation composition, however, mixing with river water is reducing the dominance of Na+. Another process occurring in the middle aquifer south of the palaeochannel is mixing with Na-HC03 groundwaters that are recharging the alluvial system directly from the GAB. Where this occurs, HC03 concentrations are observed to increase with depth (up to 1000mg/L), both dissolved inorganic carbon (DIC) and total alkalinity increase, and pH values rise above 7. Figure 4 shows the Deffeyes style relationship between total alkalinity and DIC. Since in the pH range of the samples the concentration of non carbon species are negligible it can be assumed that total alkalinity is equal to carbonate alkalinity. Additionally, the pH of groundwater is less than 8.3 so groundwater would have negligible C032-thus the comparison of carbonate alkalinity (rnHC03- + 2rnC0327 against DIC (rnH2C03 + rnHCO3- + can be used to clearly discrimi-
The stable isotope 613C is used to identify sources of carbon in the groundwater system and most importantly mixing between bedrock and alluvial groundwaters. The geochemical processes controlling carbon species is shown by the plot of 613C versus DIC-' (Fig. 5). River waters have an avera e 613C of -9.5%0. This is to be expected since the 6 3C value of atmospheric CO2 is -7%o and these waters are open to C02. The majority of alluvial groundwaters from both upper and middle aquifers have a range of 613C values between -15 and -1 1%0that reflect the intermixing of DIC from several biogenic and inorganic carbon sources including: soil zone C02, carbonate dissolution and CO2 resulting from the oxidation of organic matter. Salomons and Mook (1986) have shown that 613C signatures of -15 to -12%0 are characteristic of the dissolution of nonmarine carbonates such as soil zone carbonate (613C concentrations of 0 to -6%0) by CO2 developed in the soil zone via respiration from vegetation that photosynthesised under a C3 type pathway (6l3C%o -23%0). Na-HC03 groundwaters from the middle alluvial aquifer with pH > 7, high DIC and total alkalinity plot in a region with GAB groundwater indicating that this groundwater either has a similar source of carbon or is a mixture between GAB and alluvial groundwater. The alluvial Na-HC03 groundwater from the middle aquifer have an average 613C value of -9.1%0 and offset the main grouping of samples with a negative slope trend of decreasing I/DIC (increasing DIC) with increasingly enriched 613C %O values. The GAB groundwaters have an average 613C %O value of -8.9%0. This value is anomalous for groundwaters in the southeastern part of the GAB (Calf & Habermehl 1984) and Jankowski & McLean (2001) attributed this heavy 613C to an inorganic source of carbon: ingassing of enriched 613C~1cgeogenic CO2 combined with terrestrial carbonate dissolution. Schofield (1998) calculated an average 613C%0value for mantle derived CO2 of about -5.5%0 for the Dubbo region, NSW Lavitt (1999) similarly found groundwaters with a maximum value of -2%0 in the Lower Mooki River catchment NSW resulting
5
Figure 4. Total alkalinity (meq/l) v DIC (mmol/L).
567
River catchment can be explained by carbonate and silicate weathering reactions driven by groundwater charged with biogenic CO;?.Ca and Mg generated by these reactions exchanges with Na in clays. Carbon13 values that range from -15 to -1 l%o indicate that DIC is derived from the interaction of carbonates in the soil zone with biogenic CO2 respired by plants that utilize the C3 photosysthetic pathway. Heavy 613C values for alluvial Na-HC03 groundwater with an elevated DIC, alkalinity and pH values greater than 7 can be attributed to the upward flux of groundwater with heavier l3CDIcproduced by the interaction of CO2 with carbonate cements of the GAB bedrock units mixing with the alluvium from Jurassic GAB aquifers. 6l80 and 62H values confirm that groundwaters located close to the Namoi River in the upper and middle aquifers are receiving modern river recharge, thus explaining their heavier atmospheric CO2 equilibrated 613C values. ACKNOWLEDGEMENTS The authors thank the National Heritage Trust, Department of Land and Water Conservation and Cooperative Research Centre for Landscape Evolution and Mineral Exploration for fknding.
Figure 6. Plot of 6'*0 vs. 62H.
from inhsing geogenic CO;! and enhanced silicate weathering reactions. Although it is possible that direct ingassing of magmatidmantle CO2 is the source of heavier 13CDIc in the middle aquifer, low PC02 values indicate that this is unlikely. Rather, the enriched 613C %O values are with the result of groundwater with heavier l3CDIC entering the alluvium from the Jurassic Pilliga Sandstone unit of the GAB in this area. 6l80 and tj2H values are plotted in Figure 6 . Groundwaters from the GAB cluster left of the GMWL and have an average 6l80 value of -6.5%0 and ?i2H value of -41.2%0. Alluvial groundwaters recharged by GAB water plot closely to this group. The majority of alluvial groundwaters plot along a trendline together with surface waters that lie to the right of the GWML. This trend clearly identifies groundwater from the shallow and middle aquifers that receive modern river recharge. River water has an average 6l80 value of -2.6%0 and 62H value of 17.5%0,and falls to the right of the GMWL because the climate is semiarid and the effects of surface evaporation would enrich 6"O compared to the global precipitation average. Additionally, river water is affected by evaporation prior to recharge of the groundwater system. CONCLUSIONS The origin of Na-HC03 groundwaters in the alluvial aquifers in the southeastern part of the Lower Namoi 568
REFERENCES W, G. E. 1978. An investigation of recharge to the Namoi Valley aquiferj using environmental isotopes.Aust. J Soil Res. 16: 197-207. Calf, G.E. & M.A. Habermehl 1984. Isotope hydrology and hydrogeochemistry of the Great Artesian Basin, Australia. In: Proc. Int. Symp. isotope Hydrology in Water Resources Development, IAEA, Vienna, Austria, 12-16 Sept 1983: 3974 13. International Atomic Energy Agency, Vienna, 1984. Herczeg, A.L., T. Torgersen, A.R. Chivas & M.A. Habermehl 199I. Geochemistry of ground waters from the Great Artesian Basin, Australia. J. Hydrol. 126: 225-245. Jankowski. J. & W. McLean 2001. Origin of sodiumbicarbonate waters in southeastern part of the Great Artesian Basin: Influx of biogenic or magmatic C02. In R. Cidu (ed), Proc. 10th int. Symp. Water-Rock Int. Villasimius, Italy, 10-1.5June, 2001. Rotterdam: Balkema (this issue). Lavitt, N. 1999. Integrated approach to geology, hydrogeology and hydrogeochemistry in the Lower Mooki River catchment. Ph.D. Thesis, UNSW, Sydney. McLean, W., J. Jankowski & N. Lavitt 2000. Groundwater quality and sustainability in an alluvial aquifer, Australia. In 0. Sililo et al (eds), Proc. Xrcy Cong. IAH, Cape Town, South Africa, 26 Nov-1 Dec 2000: 567-573. Rotterdam: Balkema. Salomons, W & W.G. Mook 1980. Isotope geochemistry of carbonates in the weathering zone. In P. Fritz & J.C. Fontes (eds), Handbook of Environmental Geochemistry: Vol2 The Terrestrial Environment: 239-269. Amsterdam: Elsevier. Schofield, S. 1998. The Geology, Hydrogeology and Hydrogeochemistry of the Ballimore Region, central New South Wales. PhD Thesis, UNSW, Sydney. Wellman, P. & I. McDougall 1974. Potassium-argon on the Camozoic volcanic rocks ofNew South Wales. J Geol Soc. Aust. 2 1:247-272. Williams, RM. 1993. The Cainomic geology, hydrogeology and hydrochemstry of the unconsolidated sedimentsassociated with the Namoi River in the Lower Namoi Valley. Tech Rep. CentreN d Resmc.
Water-Rock Interaction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Mineralised waters and deep circulations in the French-Italian Alps J .P.Novel & G.M .Zuppi Dipartimento di Scienze Ambientali, UniversitZl Ca' Foscari di Venezia, Venezia, Italia
M .Dray Centre de Recherches Ge'odynamiques, Thonon-les-Bains, Universite'P. et M . Curie, Paris, France
S .Fudral & G.Nicoud Laboratoire de Ge'ologie et Hydrogbologie des Aqu$&es de Montagne, Universite' de Savoie, France
P.Lacombe Alpetunnel GEIE, Chambe'ry, France
ABSTRACT: The presence of several water points characterised by a high sodium-chloride content, depleted in oxygen-18, tritium free and with a very low carbon-14 activity, is very well known in springs in the French-Italian Alps. Nowadays, the geological, hydrological, and geochemical studies, in progress, related to the tunnels for the new high speed railway crossing the French-Italian Alps shows, in the boreholes drilled for this purposes at a mean elevation of 700 msl, the presence of brines (mainly NaC1, with an electric conductivity up to 220 mS.cm-'). Brines appear to be the base of the geological, structural and hydrodynamic systems, and replenish the most important tectonic contacts. These brines are mobilised in depth by fresh waters coming from the surface. These types of waters reappear on the surface as thermal springs. They are also intercepted by prospection boreholes at different depths and during tunnelling.
1INTRODUCTION
2 NEW DATA
The presence of water characterized by an high sodium-chloride content, depleted in oxygen-18, varying in tritium content and carbon-14 activity, is very well known in springs in the French-Italian Alps (Vuataz 1982, Cruchet 1983 & 1985, Dazy et al. 1987). Some of them, such as Salins, Brides, Uriage, Saint Vincent, Ceresole Reale, Vinadio, Valdieri, and Bex are located on the verge of external crystalline massifs (Mercantour-Argentera, Belledonne-Pelvoux, Mont Blanc), whereas others are situated along the main tectonic lines (Accidents de Fond de France, de Grand Maison, de Belle Etoile). Numerous water inflows, rich in sodium, poor in tritium, have been recognized in alpine tunnels, at a mean elevation of 1300 msl (Nagra 1988, Markha1 1999). The origin of these waters has been the concern of several publications leading to two main hypotheses. The first one is the geochemical origin from a brine similar to that from the Canadian shield already described by Frape & Fritz (1982), Fritz & Frape (1987), and Bottomley et al. (1999); the other one hypothesizes a geological origin called theory of "halocin6se" (Debelmas & Kerkhove 1980, Mascle et al. 1986), based on the presence of deep evaporitic deposits .
Nowadays, the present geological, hydrological, and geochemical studies in progress, for the new high speed railway tunnels crossing the French-Italian Alps (Fig. 1) show the presence of brines, in the drilled boreholes at a mean elevation of 700 msl. They exhibit mainly NaCl character (Fig. 2), with an electric conductivity up to 220 mS.cm-'. Brines appear to be the base of the geological, structural and hydrodynamic systems, and replenish the most important faults and the main tectonic contacts. These brines squeezed by tectonic activity, are
Figure 1. Localization of the new railway tunnels through the French-Italian Alps.
569
Figure 2. C1 vs Na Fluid chemistry indicates halite leaching; Na enrichment is linked to the interactions with silicate minerals.
Figure 3. Deuterium versus Oxygen-18. The points are aligned along the WMWL; the isotopic depletion is linked to changes in climatical conditions on the recharge areas.
mobilised in depth by fresh waters from the surface. These phenomena suggested by Mascle (Mascle et al. 1986) in the Alps and named “halocin&se”have been studied in detail in the Pyrenean chain (McCaig et al. 2000) and in the South-east of France (Reynaud et al. 2000). Waters with similar chemical content reach the ground surface as thermal waters.
Brides-les-Bains and -9.0 %o VS V-SMOW in Dignes. The whole set of waters are aligned along the global meteoric line (Fig. 3) without indicating any fractionation coming from thermal waters (Dazy et al. 1987, Vuataz 1982). Waters from confined aquifers intercepted by deep (from 600 to 1500 m) boreholes show NaCl concentration up to 10 g per litre. These waters show more depleted values in “ 0 (from -15.0 %O vs VSMOW to -16.5 %o VS V-SMOW), which do not match the present precipitation (Dray et al. 1998). Deuterium values and “ 0 values are aligned along the World Meteoric Water Line (Novel et al. 1995, Novel & Zuppi 2000). This isotopic depletion is not linked to an altitude effect but to a temperature effect associated to colder climatical conditions existing in the recharge area probably during the deglaciation following the LGM, about 18000 years BP. The larger water quantity available at that time period allowed a high hydraulic pressure (Blavoux et al. 1993) and, consequently, a per descensum circulation, even at important depths. As quaternary waters mobilise halite deposits they try to move upwards with their high salt content but confined systems prevent them to do so.
3 DISCUSSION AND CONCLUSIONS The presence of brines under more than one thousand meters lithological cover has been kept only at sites where the fracturation is poor and the quantity of fresh water capable of leaching the salts stocked is low The high solubility of halite does not allow its identification in the outcropping formations or in the cuttings of boreholes (anhydrite is mentioned in the majority of the geological logs even if the water contains higher quantities of chlorides). Nevertheless, fluid chemistry indicates the chloride leaching to be prevalent with respect to the sulphate one. Anyhow, data exibit a Na enrichment linked to the interactions with silicate minerals, which form the matrix of the geological systems. This reinforces the hypothesis of the structural confinement of the evaporites. In addition to a high sodium chloride, the spring show variable dissolved salt contents with high tritium content and high 14C activity. But salted waters found at different depths during borehole drilling and/or tunnelling reveal no tritium and low I4C activity. This highlights the fact that the percolating water mobilizes brines, trapped within the tectonic structures. The isotopes of the water molecule reflect the mean altitude elevation of the recharge area where precipitation infiltrates to dissolve evaporites. In fact, these springs exhibit variable ”0 content between -14.5 %O vs V-SMOW in
REFERENCES Blavoux, B., Dray, M., Ferhi, A., Olive, P., Groning, M., Sonntag, C., Hauquin, J.-P., Pelissier, G. & P. Pouchan 1993. Palaeoclimatic and hydrodynamic approach to the Aquitaine basin deep aquifer (France) by means of environmental isotopes and noble gases. Isotope techniques in the study of past and current environmental changes in the hydrosphere and the atmosphere. Proceedings of a symposium, I A E A Vienna 19-23 April 1993, 296-306. Bottomley, D.J., Katz, A., Chan, L.H., Starinsky, A., Douglas, M., Clark, I.D. & K.G. Raven 1999. The origin and evolution of Canadian shield brines : evaporation or
570
freezing of seawater ? New lithium isotope and geochemical evidence from the Slave craton. Chem. Geol. 155: 295-320. Cruchet, M. 1983. Relations entre l’hydrogtologie, le thermalisme et les circulations d’eaux uranifkres dans les roches fissurtes. Les massifs cristallins externes de basse Maurienne (Savoie). Thbse de 3bme cycle, Univ. de Grenoble, 235 p. Cruchet, M. 1985. Influence de la dtcompression sur le comportement hydrogtologique des massifs cristallins en Basse Maurienne (Savoie, France). Ge‘ol. Alp, 61: 65-73. Dazy, J., Dray, M., Jusserand, C., Pasqualotto, M. & G.M. Zuppi 1987. Caracttrisation isotopique des eaux thermornintrales des Alpes du Nord Franco-italiennes. In “Isotope techniques in water resources development ”, DIEA, Vienna, 3-24. Debelmas, J. & C. Kerkhove 1980. Les Alpes FrancoItaliennes. Ge‘ol. Alp. 56/57: 21 -58. Dray, M., Jusserand, C., Novel, J.P. & G.M. Zuppi 1998. Air mass circulation and the isotopic “shadow effect” in precipitation in the French and Italian Alps. Proceedings of an International. Symposium on isotope techniques in the study of past and current environmental changes in the hydrosphere and the atmosphere, I.A.E.A. proceedings, Vienna, 107 - 11 7. Frape, S.K. & P. Fritz 1982. The chemistry and isotopic composition of saline groundwaters from the Sudbury Basin, Ontario. Canadian Journal of Earth Sciences 19, 4: 645-661. Fritz, P. & S.K. Frape (eds.) 1987. Saline water and gases in crystalline rocks. GAC Special Paper 33, Geological Association of Canada, 259. Mc Caig, A. M., Tritlla, J. & D.A. Banks 2000. Fluid mixing and recycling during Pyrenean thrusting: evidence from fluid inclusion halogen ratios. Geochimica et Cosmochimica Acta 64: 3395-3412. Martchal, J.C. 1999. Observation des massifs cristallins alpins au travers des ouvrages souterrains. Hydrogkologie 23: 2142. Mascle, G., Amaud, H., Dardeau, G., Debelmas, J., Dubois, P., Gidon, M., Graciansky, (de) P.C., Kerckhove, C. & M. Lemoine 1986. Halocinkse prtcoce sur la rnarge ttthysienne alpine : vers une rtinterprttation des zones de gypse des Alpes. C.R. Acad. Sc. Paris 302, Se‘rie II, 15: 963-968. Nagra 1988. Berichterstattung uber die Untersuchungen der Phase I am potentiallen Standort Piz Pian Grand (Gemeinden Mesocco und Rossa, GR). Technischer Bericht, 89-129. Novel, J.P., Ravello, M., Dray, M., Pollicini, F. & G.M. Zuppi 1995. Contribution isotopique (’H, l80, 3H) i la comprthension des mtcanismes d’tcoulement des eaux de surface et des eaux souterraines en Vallte d’Aoste (Italie). Geogr. Fis. Dinam. Quat. (Italia) 18: 315-319. Novel, J.P. & G.M. Zuppi 2000. Hydrogtologie et gtochimie isotopique des circulations souterraines profondes des tunnels de base et de Bussoleno (future liaison ferroviaire Lyon - Turin). Rapport final de la phase I. Doe. Alpetunnel GEIE, 95 p. Reynaud, A., Guglielmi, Y., Mudry, J. & A. Emily 2000. Apport de l’hydrochimie la reconstitution paltogtographique du Trias du Sud-Est de la France (Alpes du Sud et Provence). Eclogae Geol. Helv. (in press). Vuataz, F.D. 1982. Hydrogtologie, gtochimie et gtothermie des eaux thermales de Suisse et des rtgions alpines limitrophes. Mate‘riaux pour la ge‘ologie de la Suisse Hydrologie n”29, 174 p.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Chemical and isotopic signatures of interstitial water in the French Chalk aquifer and water-rock interactions C .Plain, L .Dever, C .Marlin & E .Gibert Laboratoire d'Hydrologie et de Gebchimie Isotopique, Univ. Paris-Sud, 91405 Orsay Cedex France
ABSTRACT: A 700-m core has been drilled in the Chalk of the Paris basin (France) with the aim to know the evolution of the interaction between interstitial water and the matrix and to recognize the chemical stratification of the aquifer. The variation of the " 0 (from -4 to O%O vs SMOW) and 2H (from -50 to O%O vs SMOW) contents point out a mixing between meteoric water and a marine solution. The total salinity also increases with depth from 5 to up to 30 mS.cm-' at the bottom of the Chalk series. The chemistry of the deep salty water is Ca2+/S0;- and Ca2'/HC03- dominant and could be derived from connate water. The solid phase, through the isotope variations (13C, "O), show an increase of recrystallization of the grain matrix with depth.
1 INTRODUCTION
following the evolution and stabilisation of the electrical conductivity of the solution. The solid consisting mainly of Chalk, the leached solution will depend on the CO2 partial pressure controlling the solubility of carbonates, i.e. Ca2+,Mg2+and HC03 .
In order to recognize the origin of the interstitial water in the Chalk and the modification of this host rock caused by the interaction with water, an undisturbed 700-m core has been drilled in the South East of Paris (France) through the Tertiary formation and the Chalk. In this area, the Chalk reaches the maximum of thickness of the whole Paris-Basin, i.e. 600 m. This core goes through all the stages of the Upper Cretaceous. The Chalk sampled is white and dedolomitized in the upper part of the core, i.e. 100 to 168 m below the soil surface (bss.), and overlays a thin layer of massive dolomite (168 to 184 m b.s.s.). Below the dolomite layer, the rock consists of a slighty dolomitic Chalk (Blanc & Gily 2000).
2.1 Chalk porosity Since the Chalk is saturated with water along the core, the total porosity (p) can be calculated as (Equation 1):
were Mw, weighted water content, pw, water density, MT, total weight of the Chalk sample, and Pb, the bulk density. The porosity strongly decreases with depth, from 50% at 100 m b.s.s. to 5 % at 620 m b.s.s. (Fig. 1). The dolomite at 180-200 m b.s.s. level is marked by a low porosity (mean porosity of 20% ; Plain et al. 2000a). The global evolution of porosity along the core could be explained either by the increase of the lithostatic pressure due to compaction of materials and/or by the increasing precipitation of secondary carbonates in the microporosity with depth (Plain et al. 2000b).
2 THE AQUEOUS PHASE The interstitial water has been analysed for its chemical composition (Electrical Conductivity and major elements) and isotopic content ("0, 2H). Water for 'H and " 0 measurements has been extracted from the Chalk microporosity by vacuum distillation on samples packaged in watertight buckets to avoid isotopic fractionation. The chemical composition of interstitial water has been reached through the leaching of dried sediments (1 05OC) under controlled conditions in laboratory. The leaching experiment has been monitored by
2.2 Stables isotopes The 62H and 6"O profiles show increasing values with depth, from -50.7 to 1.3 %O and from -7.00 to 573
Figure 1. Evolution of porosity with depth.
Figure 2. variations of 6'H and 6 ' * 0 contents with depth. The accuracies of a2H and 6I80 are i 2 %O and k0.2 %O respectively.
4.15 %O vs SMOW respectively (Fig. 2). However, this trend is not regular downward, as displayed by the isotopic variations above and below the dolomite layer (Fig. 2). High heavy isotopic contents are observed above the dolomite level, while low 6values sign the core base.
The slight increase of C1- downward cannot directly confirm the hypothesis of a mixing between a connate water (marine solution) and a meteoric water. However, the above mentioned mixing could have occurred by diffusion of C1- from the bottom to the top of the core. The excess of Na' with respect to chloride can be
On Figure 3, all the samples are plotted along a line with a slope of 3.1, and thus located below the Global World Meteoric Line (GWML). The analytical points fall into the domain of evaporated water and that of the mixing between meteoric water and an evaporated seawater, as described by Sofer & Gat (1975). 2.3 Chemistry
The electrical conductivity of interstitial water highly increases with depth from 5 mS.cm-' at 100 m b.s.s. to 30 mS.cm-' at 680 m b.s.s. (Fi . 4). The chemical facies of interstitial water is Ca -HC03-S042' dominated. The conductivity evolution can be explained either by a mixing between a salty isotopically enriched water and a meteoric water (Edmunds et al. 1992), or by evaporation of the interstitial solution. The chemical contents of the leached water show an increase of all the major elements with depth. Explaining the evolution and origin of the interstitial water requires the study of the C1- concentration.
F+
Figure 3. Zi2H vs 6"O diagram. The accuracies of 62H and 6"O are +2 %O and i0.2 %O respectively. The dot line represents the evaporation of seawater (Sofer & Gat 1975).
574
Figure 4. Evolution of electrical conductivity with depth (interstitial water).
explained by cation exchange with clay minerals (Edmunds et al. 1992). The chemical evolution with depth could be explained by evaporation. This hypothesis is supported by the increase of sulphate concentration, SO$ becoming the dominant ion in the deeper part of the Chalk. However, pyrite oxidation could also justify for an increase in sulphate. Additional measurements are thus needed to discuss the origin of interstitial water. In this objective, the determination of Br- and 34Scontents are in progress.
3 SOLIDPHASE A serie of carbonate samples were analysed for I3C and " 0 contents along the borehole in order to know whether recrystallization has occurred. The " 0 contents roughly decrease from -1 to -4 %O vs PDB while 613C-valuesshow an increase from 0 to 3.0 %O vs PDB with depth (Fig. 5). These trends suggest an increase of secondary calcite proportion with depth as shown by : 1) the decrease with depth of the " 0 contents of recrystallised carbonates which suggests a secondary precipitation either from more depleted water and/or at a lower temperature; 2) the decrease of strontium concentration in the matrix as shown by Le Callonec et al. (2000), and which could be explained by
Figure 5. 6"O and the 6°C evolution of carbonates with depth. The accuracies of 6"O and 6I3C are 10.2 %O and f0.15 %O respectively
dissolution-reprecipitation processes, producing a step-by-step purification of the secondary calcite; 3) the increase of I4C activity of samples with depth. The dolomite exhibits an enrichment of +2.8 %O vs PDB in "0 and +0.2 %O vs PDB in I3C compared to surrounding carbonates. According to Fritz & Smith (1970) and Fontes et al. (1970), this enrichment can only be explained by a dolomitization of the Chalk since a dolomite precipitating in the same environment / at the same time than calcite will present an "0-enrichment of +5 to +6 %o. Moreover, the hypothesis of a secondary dolomite is in agreement with the results of Blanc & Gely (2000) who studied the dolomite of the borehole with a scanning electronic microscope.
4 CONCLUSION The interstitial water of the chalk is strongly varying with depth, becoming more enriched in heavy isotopes and more salty downward. This evolution could be the result of a mixing between an evolved connate marine-originating water and a meteoric water. Moreover, vertical distribution of interstitial water chemical contents is mainly due to an ionic diffhsion process. Cation exchanges take place at least between Ca2' and Na+, explaining the excess of 575
this latter element with respect to Cl-. At present, further studies are required to understand both the origin and the evolution of this interstitial water. Nevertheless, the decrease of porosity and of " 0 contents of carbonates can be explained by recrystallization of secondary calcite presenting depleted &values due to reprecipitation in more depleted water and at lower temperature. ACKNOWLEDGMENTS This work is integrated into the French project CRAIE700.
REFERENCES Blanc, P. & J-P. Gely 2000. Etudes petrographique et mineralogique de la diagenese carbonatee de la Craie du Cretace superieur des forages profonds 701 (Poigny) et 702 (Sainte Colombe) (Region de Provins, Seine et Marne). Bulletin d'lnformation des Ge'ologues du Bassin de Paris 37(2) : 87-100 Edmunds, W.M. Darling, W.G. Kinniburgh, D.G. Dever, L. & P. Vachier. 1992. Chalk groundwater in England and France : hydrogeochemistry and water quality. Research Rpt SD/92/2 British Geological Survey, Keyworth . Fontes, J.C. Fritz, P. & R. Letolle 1970. Composition isotopique, mineralogique et genese des dolomies du Bassin de Paris. Geochimica et CosrnochimicaActa 34 : 279-294. Fritz, P. & D.G.W. Smith 1970. The isotopic composition of secondary dolomites. Geochimica et Cosmochimica Acta 34 : 1161-1173. Le Callonec, L. Renard, M. Pomerol, B. Janodet, C. & E. Caspard 2000. Donnees geochimiques prkliminaires sur la serie Cenomanienne-Campanienne des forages 70 1 et 702 du programme Craie 700. Bulletin d'lnformation des Gkologues du Bassin de Paris 37(2) : 1 12-1 19 Plain, C. Dever, L. Marlin, C. & E. Gibert. 2000a. Approche chimique et isotopique des eaux interstitielles de la Craie et des phenomenes &interactions liquide-solide (diagenese precoce etlou tardive). Bulletin d'Information des Gkologues du Bassin de Paris 37(2) : 132-136 Plain, C. Dever, L. Gibert, E. & C. Marlin 2000b. Caracteristiques chimiques et isotopiques des eaux interstitielles de la Craie du Bassin de Paris. RST 2000, Paris. C o r n . Oral, abstract 1 p. Sofer, Z. & J.R. Gat 1975. The isotope composition of evaporating brines : effect of the isotopic activity ratio in saline solutions. Earth and Planetary Science Letters 26 : 179-186.
576
Water-Rock lnferacfion 2001, Cidu (ed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Water-rock interaction processes in the main thermal springs of Sardini (Italy) Monica Proto & Costanzo Panichi Istituto Internazionale per le Ricerche Geotenniche, Pisa, Italy
Paola Zuddas & Francesca Podda Dipartimeizto di Scienze della Terra, Universita di Cagliari, Italy
ABSTRACT: For many years the thermal systems of Sardinia have been studied with great interest, with the aim of evaluating the potential use of geothermal energy. Chemical studies based on main components and trace elements have shown that groundwater circuits in the main hydrothermal systems of Sardinia develop prevalently in the Paleozoic basement both in the presence and absence of deep COZ. The chemical compositions of thermal springs have been ‘reconstructed’by w/r interaction processes between cold shallow waters and different sets of minerals in the main reservoirs at different thermal conditions. The results obtained so far show: 1) that temperature plays a small role in defining the solute concentrations in the studied systems, because the reservoir temperatures are all around 100°C, while pH variations (6.7 to 9.0) and total salinity are controlled by the CO2 pressure in the solutions; and 2) that the mineral assemblages are very similar to each other in spite of existing differences in the geological setting of the thermal areas.
1 INTRODUCTION Natural waters acquire their chemical characteristics by dissolution and chemical reaction with the solids, liquids, and gases contacted during the various stages of the hydrological cycle (Stumm & Morgan 1981). The thermal waters of Sardinia have been chosen as an experimental field of this application. This work is an attempt at reconstructing the chemical composition of these specific thermal waters, and underlines the contribution of several factors, such as temperature, COZ,ionic strength, and the mineral phases involved in the water-rock interaction. In Sardinia there are two large geologic domains: the Paleozoic basement and the Alpine complex. The Paleozoic basement is made up of metamorphic rocks of varying grade, and an extended Carboniferous intrusive complex. Small outcrops of Mesozoic marine series are also present in few areas. The Alpine complex is present mainly in the west of island. It is made up of the Oligo-Miocene calcalkaline volcanic series and Tertiary sediments. This complex fills two main tectonic structures: the Oligo-Miocene Sardinian rift and the more recent Plio-Quaternary Campidano graben. The main thermal systems are located along regional faults, generally on the borders of the Sardinian rift, where the Paleozoic basement is in tectonic contact with the Oligo-Miocene Sardinian
rift and the more recent Plio-Quaternary sediments of the Campidano graben. The chemical composition of the thermal waters (Frau 1993, 1994) is shown in Table 1.
Figure 1. Simplified geological map of Sardinia. 1: Paleozoic basement; 2: Mesozoic carbonate rocks; 3 : Oligo-Miocene andesitic and ignimbritic rocks and Tertiary sediments; 4: PlioQuaternary sediments; 5: Main faults; 6: Thermal waters.
577
Table 1 Chemical composition of the thermal waters of Sardinia (from Frau 1993, 1994). Ion concentrations are expressed in mgA.
Na
K
SiOz
Alk
SO4
C1
<0.1
156
2.8
41
38
40
208
<0.1
203
1.9
47
45
44
295
34
13.0
517
21
54
908
50
322
27
6.7
870
41.5
48
1748
77
512
8.6
631
0.4
1012
7.2
48
17
50
2540
7.3
675
2.3
1090
63
32
287
2865
pH
Ca
Mg
42
9.0
13
55
8.6
15
Acquacotta (AQ)
45
6.7
Sardara (SA)
55
7.2
s u campu (SC)
43
Springs
T
Benetutti (BN) Fordongianus (FD)
Casteldoria (CD)
75
49
2 APPLIED METHODOLOGY
3 RESULTS
The Phreeqe computer program (Parkhurst et al. 1980) was used to simulate the various phases of waterhock interaction. The initial solution is assumed to come from a meteoric water, and to be immature because of its poor circulation rates and low temperatures (on average lower than 25°C). As it sinks, the temperatures get higher reaching a maximum in the main thermal reservoir. It is also assumed that the water rock interaction increases on increasing the circulation time. The very low Tritium contents in the thermal waters of Sardinia suggest a circulation longer than 30 years without mixing with recent shallow waters (Proto 1999). Chlorine is taken as a reference parameter of the circulation time, both because it is a mobile element and because it is considered to come from the breaking and leaching of fluid inclusions commonly present in the rocks. As regards water-rock interaction, the set of mineral phases that contribute major ions to the geocheinical characterisation of the solution must be defined. The choice of the effective phases was based on lithological evidence in the various specific areas. The saturation index (SI) was taken to control the solubility reactions of the effective mineral phases and state the relationships among the dissolved solutes. Three different stages have been used to describe the water-rock interaction processes at different temperatures, ionic strengths (IS), and CO2 conditions. Stage 1 - infiltrating water: meteoric water interacting with a specific set of minerals at surface conditions (atmospheric pC02, T=25"C). Stage 2 - deep reservoir water: water interacting with a specific set of minerals at reservoir conditions. Stage 3 - emerging water: a solution rising from the reservoir at the point of emergence of the thermal springs.
The results obtained in each stage for the Benetutti and Sardara thermal systems are illustrated below as examples. These systems have different geological settings and are characterised by different chemical compositions. 3.1 Benetutti thermal system The Benetutti thermal system is located in the Tirso Valley, at the tectonic contact between the Paleozoic basement and the Oligo-Miocene Sardinian rift (Fig. 1). Simulations of these waters show that the interaction between infiltration water and silicate minerals, which are the main constituents of granite lithotypes, is most important in the acquisition of solutes. Stage 1 - The initial solution is defined by very low salinity and a NaCl concentration as low as 6*1O"M, as in the shallow waters circulating in granite lithotypes (Frau 1993). In this stage the solution is at equilibrium with respect to calcite (Calc) and strongly undersaturated with respect to anhydrite (Anhy). The silica phase considered is chalcedony (Chalc). The specific mineral phases and related saturation indices are the following: Minerals SI
Anhv -2.3
Calc 0
Chalc -0.5
The chemical composition (expressed in mg/l) of the computed solution at the end of this stage is: pH 8.18
Ca Mg 33.1 -
Na 22.9
K
-
SiOz SO4 HC03 C1 5.33 43.6 53.8 35.5
Stage 2 - The low-salinity water percolates, and at deeper levels reaches a temperature of 93°C and higher salinity. According to data reported by Ghezzo & Orsini (1982) for granite rocks in the area, the mineral 578
paragenesis is principally made up of quartz (Qtz), albite (Ab), microcline (Micr), chlorite, and kaolinite (Kaol). Calcite is also considered because it is typical of rock fracturing, though no carbonate formations are present in the area. The specific saturation indices of this stage are the following: Minerals ST
Ab -1 2
Micr 0.2
Kaol 0.6
Qtz 0
Stage 2 - The solution reaches a maximum temperature of 107°C and interacts with the mineral phases of the granite metamorphic basement: calcite, dolomite, quartz, albite microcline, and kaolinite. In this step p(COz)=10d.8atm is chosen. Saturation indices are as follows: Minerals LogpCOz SI -0.8
Calc
83
Ca 13
Mg
Na 155
K 3
Si02 SO4 HC03 40 40 60
BNc BNa
Ca Mg Na K Si02 SO4 HC03 C1 12.7 <0.1 155 2.76 42.1 40.8 38.4 213 12.8 <0.1 156 2.74 40.9 40.3 37.8 208
3.2 Surduru thermul system The Sardara spring waters are brackish, neutral, and of the HC03-type. The Sardara thermal system is located in the Campidano area (Fig. l), at the contact between the Paleozoic basement and the more recent sediments of the Campidano graben. Stage 1. The infiltrating water is characterised by very low salinity and a NaCl content as low as in the shallow cold waters circulating in the basement rocks of the study area (Frau 1993). The specific saturation indices of this stage are: LogpCOz -3.5
SI
Calc
Do1
Chalc
0
-0.5
-0.5
DH 7.54
Ca 41
Mg 6
Na 14
K
-
Si02 5
Calc
0
0
SO4 HC03 88 46
Ca 88
ME 5
Na 887
K 43
Si07 SO, 55 77
HCO? 1590
C1 496
pH Ca Mg Na K SiOz SO4 HC03 C1 SAC 7.2 28 6.5 857 41.4 53.1 76.8 1560 496 SAa 7.2 27 6.8 869 41.1 54.1 76.8 1750 511
4 DISCUSSION AND CONCLUSIONS Table 2 summarises the main calculated chemicalphysical parameters of the Sardinian thermal waters, mainly referred to the conditions assumed for stage 2 of the simulation of each thermal system. This stage is considered the most important in the geochemical evolution of groundwaters. Table 2. Main chemical-physical parameters of the Sardinian thermal waters referred to stage 2. BN=Benetutti, FD=Fordongianus, SA=Sardara, AQ=Acquacotta, SC=Su Campu, CD=Casteldoria. Area Springs pCOz (bar) TDS (mg/l) PH T,(OC) Minerals Calcite Dolomite Quartz Albite Microcline Chlorite Kaolinite IS
~~
pH 8.14
Qtz
Stage 3 - During the rising stage, the solution cools down to 55"C, and conditions of partial reequilibrium with kaolinite and calcite occur. A solution (SAC) that is very similar to the composition of the Sardara waters (SAa) is obtained.
The chemical composition of the computed solution at surface conditions (T = 25"C, atmospheric pCO;!) is: ~
Kaol 0.7
An external source of CO;! generally facilitates hydrolysis with a consequent increase in salinity. At reservoir conditions the obtained solution is neutral and of the HC03- type:
C1 213
The obtained solution reflects the composition of the Benetutti thermal springs, except for pH and alkalinity. The differences in pH are mainly due to the temperature at which hydrolysis occurs. As a matter of fact, in alkaline waters as the temperature decreases, increases in pH values are also observed (Fouillac 1983). Stage 3 - During the rising stage, the solution cools and reaches re-equilibrium with respect to kaolinite and chalcedony. The comparison between the computed solution (BNc) and the sampled one (BNa) at 45°C expressed in mg/l is: pH 9.0 9.0
Micr 0.8
0
The chemical composition of the computed solution at reservoir conditions (T=93"C) defined at the end of stage 2 is: pH
Ab -1.0
C1 21
579
Tirso Valley BN-FD 1O".' 500-700 z8.0 z1OO SI 0
0 <-1.o 0-0.5 0 0-0.5 0.08
Campidano SA-AQ g10-l 1000-3000 z7.0 90-110
SI 0 0.5 0 -1.0 >0.5
0.5-0.8 0.02-0.05
SC
Anglona
3000 z7.5 ~100 SI 0
CD 103.5 4000 <7.0 ~120 SI 0
0 <-1.0 0-0.5 0 0.5 0.08
0 <-1.5 0-0.5 0 0.5 0.1
10-~.~
examples of the thermal systems of Benetutti and Sardara.
REFERENCES
Figure 2. Simplified scheme of the three stages of the geochemical model for the thermal water of Benetutti and Sardara. 1 : Paleozoic schisdgranite; 2: Cixerri Formation (Eocene); 3 : Andesitic lavas and breccias (Oligocene-Lower Miocene); 4: Ignimbrite deposits; 5: Miocene tuffs and sediments; 6: Quaternary deposits.
Foulliac, C. 1983. Chemical geothermometry in C02-rich thermal waters. Example of the French Massif Central. Geothermics 12 (213): 149-160. Frau, F. 1993. Selected trace elements in groundwaters from the main hydrothermal areas of Sardinia (Italy) as a tool in reconstructing water-rock interaction. Miner. Petrogr. Acta 36: 281-296. Frau, F. 1994. A new hydrothermal manifestation in the Campidano graben, Italy: the Su Campu borehole (Monastir). Miner. Pefrogr. Acia 37: 155-162. Ghezzo, C. & J. B. Orsini 1982. Lineamenti strutturali e composizionali del Batolite ercinico sardo-corso in Sardegna. Guidn alla Geolopa del Paleozoico sardo. Glide Geologiche Regioriali. Soc. Geol. It. : 165-18 1. Parkhurst, D.L., Thorstenson, D.C. & L. N. Plummer 1980. PHREEQE - A computer program for geochemical calculations. US. Geological Siirvey Richmorid, VA. Proto, M. 1999. L’importanza dell’interazione acqua roccia nella formazione delle acque dei principali sistemi termali della Sardegna. Universita di Pisa. Bachelor Degree Thesis. Stumm, W. & J.J. Morgan 1981. Acquatic Chemistry. New York: Wiley & Sons.
The differences in the temperature of the deep reservoir (Tres)are not significant, since they are all about 100°C. The natural solutions do not always reach equilibrium conditions with respect to some interacting minerals, even in the deep stages characterised by high temperatures. In the various minerals and their disequilibrium rates, no large differences among sets of waters are observed. All solutions were undersaturated with respect to albite, though at a slightly different rate for higher-salinity waters. Conditions of slight oversaturation occur in all the waters with respect to kaolinite and microcline, and saturation conditions prevail with respect to calcite and quartz. All things considered, Table 2 indicates that the specific set of ((effective))minerals for the different waters is similar enough, though the individual thermal areas have a different geological setting. CO2 pressure plays an important role in controlling the pH and TDS values. Low CO2 waters are alkaline and have a low TDS, while high CO2 waters are neutral, slightly acid, and characterised by medium-high salinity. Hence, the different ion contents in solution, expressed as different values of ionic strength, as well as the greater influence of CO2 dissolution on the reactivity of minerals, is the main differentiation factor. Figure 2 shows the three stages defined for the 580
Wafer-Rock lnferaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Arsenic and other redox-sensitive elements in groundwater from the Huhhot Basin, Inner Mongolia P.L.Srnedley British Geological Survey, Wallingford, Oxfordshire, UK
M.Zhang, G.Zhang & Z.Luo Huhhot Antiepidernic and Sanitation Station, Huhhot, Inner Mongolia, China
ABSTRACT: Arsenic concentrations ranging between <1 pg 1'' and 1484 pg 1-' (n=73) have been detected in groundwaters from Holocene lacustrine and alluvial aquifers in the Huhhot Basin of Inner Mongolia. Populations from the affected areas present with a number of As-related health problems. The high As concentrations occur in anaerobic groundwaters from the low-lying central parts of the basin and are associated with moderately high Fe as well as high Mn, NH4, dissolved organic carbon (DOC) and P concentrations. As(II1) constitutes around 60-90% of the total dissolved As. The highest As concentrations are also found in groundwaters with low SO4 concentrations and indicate that As mobilisation occurs under strongly reducing conditions, where SO4 reduction has occurred. High concentrations of fluoride (up to 6.8 mg I-') also occur in some of the groundwaters where Na-HC03 ions dominate. 1 INTRODUCTION Arsenic has been recognised in groundwater from a number of major aquifers in recent years and has received much scientific and media attention as a result of the element's health impacts. Long-term use of high-arsenic drinking water has been linked to a number of health problems, including skin disorders, cardiovascular problems and cancer (e.g. Smith et al. 1998). The WHO guideline value for arsenic in drinking water is 10 pg 1-* but national standards vary and many countries retain the pre-1993 WHO guideline value of 50 pg 1.' as their standard value. Among the affected regions are parts of northern China, including areas of Xinjiang and Shanxi Provinces and Inner Mongolia (Luo et al. 1997, Lin and Tang 1999). Many of these areas are poorly documented and the extent and causes of the contamination not well understood. One area with recognised arsenic-related health problems is the Huhhot Basin of Inner Mongolia. The basin, with an area of around 4800 km2, lies on the margins of the Yellow River Plain and incorporates the city of Huhhot, the capital of Inner Mongolia Autonomous Region (Fig. l). Symptoms of chronic arsenic poisoning were first recognised in rural populations in the Huhhot Basin in 1990. The distribution of symptoms recognised to date is largely restricted to a few villages in the central parts of the basin. The most prevalent of these symptoms include skin keratosis and skin pigmentation disorders, although increased incidences of skin, lung, bladder and liver cancer have also been recog-
nised (Luo et al. 1997). Of the internal cancers, lung cancer is the most prevalent. Feng and Lamm (1999) identified a distinct dose-response relationship with observed symptoms in the exposed populations. Since 1990, around 200 menic-affected patients have been identified in the area. In this study, groundwater samples have been collected from 73 groundwater sources (70 boreholes and 4 dug wells) in the region to assess the regional distributions of arsenic and related constituents in the groundwaters and to identify the main geochemical processes involved. 2 REGIONAL SETTING The Huhhot Basin lies on the edge of the Gobi Desert and so experiences an arid climate with an average annual precipitation of around 440 mm (mainly during July to September) and average monthly temperatures ranging from - 13 to +22"C. The basin is drained by the Black River and other minor river systems and is bounded to the north by the Da Qing Mountains (typical altitude 2200 m) and to the south and east by the Man Han Mountains (Fig. 1). The crystalline rocks forming the mountains are highly folded and fractured and the basin is faultbounded (Zhang and Shao 1991). The alluvial plain has a gentle west-south-west slope (altitude around 1100-980 m) and is composed of a thick sequence (around 1500 m) of sands, silt and clay, largely of Quaternary age. The sediments are the product of a former lake which occupied the basin until around 581
in the deep aquifer, artesian boreholes. These have largely been installed over the last few decades. Traditional water sources in the region were usually shallow hand-dug wells abstracting from the topmost part (around 10 m) of the aquifer. 3 HYDROGEOCHEMISTRY 3.1 Major constituents
Fig. 1. Map of the Huhhot Basin showing dissolved oxygen in groundwaters from the shallow and deep aquifers.
8000 BP (Li, pers. commun. 1998), together with alluvial-fan deposits along the mountain foothills. Holocene sediments outcropping at surface are notably coarser-grained in the basin margin areas than in the low-lying parts of the basin. The Black River is currently an ephemeral river which in times of flood, flows into the Yellow River beyond the limit of the basin. Drainage to the Yellow River ceases at other times due to seasonal evaporation and irrigation losses. The low slope of the basin and poor drainage to the Yellow River catchment has resulted in sluggish groundwater flow and the generation of an essentially closed basin. Groundwater is an important resource and its abundance was one of the key reasons for the siting and development of Huhhot City, where the resource is abstracted for industrial and domestic supply. The population of Huhhot is currently around 3 million people, while that in the rural parts of the basin is estimated to be around 300,000. Abstraction for irrigation and domestic use is also important in the rural areas, although additional water for irrigation is taken from the Black River. Wells in the aquifers abstract from either shallow (mostly <30 m) or deep (>loo m) aquifers. The piezometric surface is usually shallow (typically 1-3 m below ground level) and the deep aquifer is artesian along the northern margin of the basin. Groundwater abstraction for potable use in the rural areas is from a variety of hand-pumped and electric tubewells and,
Groundwaters vary in chemical composition both laterally along the flow gradient and with depth. Along the basin margins, shallow groundwaters are generally of Ca-HCO? type but become more saline with Na-HCOs-dominant or mixed-ion types in the central parts of the basin and the deep groundwaters. Highest salinities occur in the shallow groundwaters in the south-central part of the basin, largely as a result of evaporation, exacerbated by irrigation. Concentrations of C1 in the shallow aquifer reach up to 919 mg I-* and Na up to 835 mg I-'. Boron reaches up to 2 mg I' (Smedley et al. 2000a). Groundwater pH is near-neutral to alkaline (7.00-8.56). The groundwaters display a strong redox gradient between the basin margins and the lower-lying central parts of the basin. This reflects at least in part, the changing lithology of shallow sediments from relatively coarse marginal alluvial deposits to finer lacustrine sediments in the central part of the basin. The change leads to the development of increasingly confined aquifer conditions downgradient. Both shallow and deeper groundwaters are aerobic (dissolved oxygen up to 9.8 mg I-') along the basin margins but become anaerobic downgradient (Fig. 1). Redox potentials similarly evolve from high values (up to 426 mV) to a minimum of -74 mV. The aerobic groundwaters contain nitrate (N03-N up to 34 mg I-'), so4 (up to 1007 m I-') and small concentrations of Se (up to 5 pg 1-f:). These diminish to below detection limits in the anaerobic sections of the aquifers. Redox-sensitive elements in the anaerobic groundwaters include in particular, Fe (concentrations up to 4.5 mg I-'), Mn (up to 1.3 mg I-') and NH4-N (up to 18 mg 1-'; Smedley et al. 2000a). Concentrations of P and HC03 are also high in many of the groundwaters, particularly in central parts of the basin. Maximum observed concentrations were 3.1 mg 1.' and 1150 mg I-' respectively. One feature of particular note in the groundwaters is the relatively high concentrations of DOC. Concentrations of organic matter are especially high in groundwaters from the deep aquifer (up to 30.6 mg I-' DOC), which are often discoloured as a result of the presence of humic acids. Concentrations of DOC reach up to 14.9 mg I-' in sampled groundwaters from the shallow aquifer. Organic matter in the sediments is not unusually enriched (average concentration 0.26 wt % in 12 analysed sediment samples ranging 582
Fig. 3. Relationship between total As (AsT) and SO4 in groundwaters from the Huhhot Basin. The WHO guideline value (10 pg I-') and Chinese national standard (50 pg 1-') for As in drinking water are also given.
Fig. 2. Regional distribution of total As (AsT) in groundwaters from the shallow and deep aquifers of the Huhhot Basin.
from sands to clay). However, the high concentration of dissolved organic matter in many of the groundwaters is believed to have been an important control on the generation of reducing conditions in the low-lying parts of the aquifers. 3.2 Arsenic Concentrations of As in the sampled groundwaters range between <1 and 1484 pg I-' (median 2.9 pg I-'; n=59) in the shallow aquifer (100 I m) and between <1 and 308 pg 1-' (median 128 pg I-'; n=14) in the deep aquifer (>loo m). The regional distributions of As in the groundwaters are shown in Fig. 2. Concentrations in the aerobic groundwaters from the basin margins are universally low. High concentrations, greater than the WHO guideline value (10 pg I-') and the Chinese standard (50 pg I-') occur in the low-lying part of the basin where groundwaters are anaerobic. Measured concentrations of As(II1) range between <0.9 and 1290 pg 1-' (median 2.7 pg I-') in groundwaters from the shallow aquifer and between <0.9 and 199 pg I-' (median 100 pg I-') in the deep aquifer. Arsenic(II1) typically constitutes some 60-90% of the total As present (Smedley et al. 2000a). Other As species in solution were not measured directly but are believed to be dominantly As(V), although minor amounts of organic As species may also be present, as has been determined elsewhere in Inner Mongolia (Lin and Tang 1999). Fig. 3 shows that a general negative correlation
exists between total As and SO4 in the Huhhot groundwaters. Many of the high As concentrations exist where SO4 is 5 mg 1.' or less. This suggests that As has been released into solution under strongly reducing conditions, coincident with reduction of SO4. Also worthy of note is the broad correlation between dissolved As and DOC (Fig. 4). It is not clear whether the As concentrations are enhanced by complexation with DOC, or whether the correlation of the two is simply a coincidence of both constituents being found prevalently under reducing conditions. High concentrations of dissolved organic matter have been found in some other reducing highAs groundwaters (e.g. Varsanyi et al. 1991), though they are not a characteristic feature of all such groundwaters (Smedley et al. 2000b). Only a limited number of samples of dug-well water were collected in this study, but of those investigated, two had concentrations in excess of 50 pg I-' (i.e. 556 yg I-', and 204 pg 1.'; Table 1). These high-As wells are also from the low-lying central part of the basin. The presence of high concentrations of As in dug-well waters contrasts with the situation in other reducing groundwater environments with As problems such as Taiwan (Guo et al. 1994) and the Bengal Basin. In Bangladesh, observed As concentrations in dug-well waters are around 10 pg 1.' or less (Kinniburgh and Srnedley
Fig. 4. Relationship between AsT and DOC in groundwaters from the Huhhot Basin. The WHO guideline value and Chinese national standard for As are also given.
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value of 21 pg 1-' exceeded the WHO guideline value for Ni (20 pg I-').
Table 1. Chemical data for dug-well waters from the Huhhot Basin. Sample Well Water Eh depth level m m mV HB2 3.5 1.5 87 6 2.0 HB18 HB58 8 4.0 247 HB4 9 2.0 326
HC03
DOC
SO4
AsT
mg 1.' 740
mg 1.' 9.3
424 901
2.5 11.4
mg IF' 144 34 254 168
yg 1.' 556 48.9 <1 204
-
4 SUMMARY High concentrations of As in groundwaters from the Huhhot Basin are associated with strongly reducing conditions, where dissolved Fe, Mn, NH4-N and DOC are relatively concentrated and where SO4 concentrations are relatively low, suggesting that some sulphate reduction has occurred. High concentrations of HC03 and P are also a notable feature of many of the affected groundwaters. The hydrogeochemical features described have many similarities with other reducing high-arsenic groundwater provinces such as Bangladesh, Taiwan and Hungary. However, the arid conditions and importance of evaporative processes results in groundwaters of higher salinity than in these other areas. As a result, fluoride is present in some groundwaters at concentrations up to 6.8 mg 1.' and poses an additional health problem.
2000). The redox conditions of the Huhhot dug wells are apparently variable, but the sample with the highest As concentration has a low Eh (87 mV; Table l), suggesting that the groundwater in the well is reducing, despite being open to the atmosphere. The presence of relatively high concentrations of organic matter (DOC up to 11.4 mg 1-I) may be an important reason for the maintained reducing conditions. However, high concentrations of so4 (up to 254 mg l-'; Table 1) suggest that some oxidation of sulphide minerals in the aquifer may have taken place. 3.3 Other trace elements
Of the other inorganic constituents in the Huhhot groundwaters, that with the greatest health significance is fluoride. Concentrations in the deep aquifer range between 0.13 and 2.35 mg 1'' but reach higher values in the shallow aquifer (0.14-6.8 mg I-'; Fig. 5) where groundwaters are Na-HC03-dominated (low-Ca). Fluoride shows no overall correlation with As. Of the analysed samples, 26% exceeded the WHO guideline value for F of 1.5 mg 1-'. Some of the groundwaters also contain concentrations of U in excess of the WHO (1998) guideline value of 2 pg 1-I. Concentrations range between <0.01 and 53 pg 1.' in the shallow aquifer and between <0.01 and 12.3 pg 1.' in the deep aquifer. 46% of analysed samples exceeded 2 pg I-'. The health impacts of chronic ingestion of U at these concentrations are poorly known, but the concentrations observed should be cause for some concern. Of the remaining trace elements measured, 7% exceeded the WHO guideline value for Mn in drinking water of 0.5 mg I-'. None exceeded the values for Pb, Cr, MO,B or Hg. Only one sample with a
REFERENCES Feng, X.L. & Lamm, S.H. 2000. IMCAP Report, Consultants in Epidemiology & Occupational Health, Washington DC. Guo, H.R., Chen, C.J. & Greene, H.L. 1994. In: Arsenic Exposure and Health, eds: Chappell, W. R., Abernathy, C. 0. and Cothern, C.R. 129-138. Kinniburgh, D.G. & Smedley, P.L. 2000 (eds.) Arsenic in groundwater in Bangladesh. BGS Technical Report, wc/oo/19. Lin, N. & Tang, J. 1999. Tne study on environmental characteristics in arseniasis areas in China. Scient. Geog. Sinica, 19, 135-139. Luo, Z., Zhang, Y., Ma, L., Zhang, G., He, X., Wilson, R., Byrd, D.M., Griffiths, J.G., h i , S., He, L., Grumski, K. & Lamm, S.H. 1997. Chronic arsenicism and cancer in Inner Mongolia. In: Arsenic Exposure and Health Effects, eds: Abernathy, C. 0. et al. 55-68. Smedley, P.L., Nicolli, H.B. & Luo, Z. 200021. Arsenic in groundwaters from major aquifers: sources, effects and potential mitigation. BGS Technical Report, WC/99/38. Smedley, P.L., Kmniburgh, D.G., Huq, I., Luo, Z. & Nicolli, H.B. 2000b. International perspective on arsenic in groundwater. In: Arsenic Exposure and Health Effects, eds: Chapell, W., et al. Chapman & Hall. Smith, A. , Goycolea, M., Haque, R. & Biggs, M.L. 1998. Marked increase in bladder and lung cancer mortality in a region of northern Chile due to arsenic in drinking water. Environ. Health Persp., 97, 259-267. Varshnyi, I., Fodr6, Z. & Bartha, A. 1991. Arsenic in drinking water and mortality in the southern Great Plain, Hungary. Eiwiron. Geochem. Health, 13, 14-22. WHO 1998. Guidelines for Drinking- Water Quality. Addendum. WHO, Geneva. Zhang, Z. & Shao. S. 1991. The Quaternary of China. China Ocean Press.
Fig. 5. Relationship between F and well depth in groundwaters from the Huhhot Basin.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Water-Rock reactions in a deep barite-fluorite underground mine, Black Forest, Germany Ingrid Stober Geological Survey, Albertstr. 5,D-79104 Freiburg, Germany
Yinian Zhu & Kurt Bucher Institute of Mineralogy, Albertstr. 23b, 0-79104 Freiburg, Germany
ABSTRACT: Mineralization in the more than 700 m deep underground barite and fluorite mine Clara (Black Forest, Germany) occurs in 3 major vein systems within the gneisses. They are associated with quartz, carbonate, various sulfides, natural alloys and about 250 species of secondary minerals. Water transport is strongly focused within the mineral veins owing to their high fracture and cavity related permeability. Water from 25 different sampling points in and around the mine were collected and analyzed monthly during one year in order to elucidate the interaction of surface water with the various rock types. On the basis of the relative proportions of the major ions 3 types of water could be distinguished: Ca-HC03-, Ca-SO4-HC03- and Ca-S04-C1-water. The source of the solutes is related to fluid-rock reaction, especially feldspar weathering. All waters within the mine are very close to quartz saturation. They are saturated or a bit undersaturated with respect to calcite after relatively short flow distances through the fracture pore space. All waters are undersaturated regarding dolomite or gypsum. All waters fall in the stability field of kaolinite which is an observed major alteration product in veins and rocks.
1 GEOLOGICAL SETTINGS
AND
2 WATER SAMPLES
HYDROLOGICAL
Water samples from >20 different localities (exploration drillholes, open fractures etc.) in and around the barite and fluorite mine Clara were collected monthly between in 1997 and 1998. We collected snow and rain water as well. Water temperature, pH, electrical conductivity, Eh and 02-concentration were measured in-situ. Major cations (Li', Na+, K+, Ca2+,Mg2+)and major anions (Cl-, SO:-, NO;, F-, Br-) were analysed by IC, trace elements by Flame AAS and by hollow graphite furnace AAS. Analyses with E.N. better than 5% were used in the study.
The underground barite and fluorite Clara mine is situated in the central gneiss complex of the Black Forest, SW-Germany. The gneiss contains locally small granite veins. The rocks in the mine area are predominantly banded, coarse grained biotiteplagioclase-gneisses, locally with garnet. Subordinate garnet-amphibolite and garnet-rich gneiss is present as well. The gneiss consists mainly of plagioclase ( A ~ ~ o - A ~orthoclase, o), biotite, quartz, amphibole and accessory minerals. Small fractures are filled with calcite, gypsum and pyrite etc.. The crystalline basement is partly covered by triassic sandstone. The mineralized veins (barite, fluorite) are also present in the (Fig. 1). The barite and fluorite mineralization occurs in 3 major vein systems in the Black Forest gneisses (Huck 1986). In the crystalline basement water flows along specific geological discontinuities (Stober 1995). Water transport in the Clara mine is strongly focused in the mineral veins owing to their high fracture and cavity related permeability.
3 RESULTS The in situ measured pH values vary between 4.75 (snowmelt) and 8.30. The major components of selected waters are shown on the Schoellerdiagramm (Fig. 2). Each water type has distinct patterns. The compositional variation is larger for the anions than for the cations. 585
Figure 1. Vertical section through the Clara mine with barite and fluorite.
In all waters Ca+Mg > 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 pH<2 and refrigerated. Analysis was by Inductively-Coupled Plasma Optical Emission Spectroscopy (ICP-OES). Samples for anion analysis were collected (unfiltered & unacidified) and refrigerated for analysis by ion-exchan e chromatography. Samples for stable isotopes (6’80 and Z2H) were collected in glass vials fitted with sealed lids, and analysed at Monash University using a Finnigan MAT 252 mass spectrometer. Stable isotope data were normalised following Coplen (1988) and are expressed relative to V-SMOW.
Figure 2 . Simplified hydrostratigraphic section for the Murray Basin, modified after Brown (1989).
recharged through surface lakes and small depressions perched above the water table in the central Wimmera palaeo-sand ridges, between Edenhope and Nhill (Fig. 1). Groundwater flow in the Parilla Sand and Renmark Group aquifers east of the Douglas Depression is to the north. West of the depression the flow is mainly towards the northwest. Groundwater flow in the Murray Group Limestone is predominantly west-northwest. Typical hydraulic gradients in the Parilla Sand and Renmark aquifers are approximately 0.0025 m/m. Horizontal hydraulic conductivity is -0.5 to 5 &day in the Parilla Sand and -1 to 7 m/day in the Renmark (Lawrence 1975, McAuley 1992). The Douglas Depression is a topographically low ‘trench’ 5 - 10 km wide and containing over 80 salt lakes and playas along its 120 km length. The Wimmera River within the depression flows north and acts as a groundwater discharge zone for the Parilla Sand aquifer (Brownbill 1995), particularly under low-flow conditions where deep pools in the river become stratified with saline water at depth and fresh water at the surface (Anderson & Morison 1989). During wetter periods, much of the salt is flushed from the river, either by natural flow into Lake Hindmarsh, or recharged back into the underlying Parilla Sand aquifer. However, isotopic and chemical data from recent groundwater sampling suggests that in areas within the Douglas Depression highly saline groundwater from the Parilla Sand aquifer penetrates the underlying Bookpumong Beds (upper confining layer) and Geera Clay (semiconfining unit) and recharges the deeper Renmark Group aquifer (Fig. 2). Although part of a major recharge area for the southern Murray Basin, the Parilla Sand (and in some areas the Renmark Group) shows significant sensitivity to local topographic controls, such as the central Wimmera palaeo-sand ridges and the Douglas Depression. This suggests that local recharge and discharge areas are critical controls on the hydrodynamics of this region.
3 RESULTS AND DISCUSSION Total dissolved solids (TDS) contents in groundwater in the palaeo-sand ridges (at Jungkum, Fig.1) of both the Parilla Sand and Renmark Group aquifers are low: 650 and 600 mg/L respectively. On either side of this central ridge, TDS contents are significantly higher (Fig. 3) despite the palaeo-ridges being on average only 40-50m higher than the surrounding area (70m higher than the Douglas Depression). Groundwater quality in the Douglas Depression, -27 km east of Jungkum, contrasts greatly to that in the palaeo-sand ridges. As groundwater flows through numerous salt lakes in the depression, TDS in the Parilla Sand aquifer increases from south to north along the depression: Jilpanger (4400 mg/L), Quantong ( I 6,000 mg/L) and Duchembegarra (105,000 mg/L). Natimuk, which is located midway, but slightly east of the depression, has a TDS 590
100000
10000 I000
10
- Gerang Gerung ........... -.-.
Duchernbegarra Quantong
---
-.._
Natirnuk Jilpanger Jungkum
Figure 3. Major ion concentrations in groundwater sampled from wells within the study area.
in the Parilla Sand similar to that of the southern depression eg. Jilpanger (4400 mg/L). Gerang Gerung, located at the northern end but slightly west of the depression, has a TDS in the Parilla Sand of 950 mg/L. This indicates that highly saline water in the Parilla Sand within the depression is having little effect on the regional chemistry of the Parilla Sand aquifer outside the depression. This observation is supported by groundwater elevation data collected previously (McAuley et al. 1992, Brownbill 1995), and also in this study. The chemical and head data suggest that groundwater flow in the Parilla Sand aquifer enters the Douglas Depression in the south and is funnelled northwards along a narrow flowpath. Within the basal Renmark Group aquifer, TDS contents are relatively low in wells outside the depression: eg. Jungkum (600 mg/L) and Natimuk (1200 mg/L). Midway along the depression at Quantong and Duchembegarra, TDS contents of 9000 mg/L and 17,000 mg/L respectively were recorded throughout 1999. At Jilpanger (south), where the Renmark Group is absent, high TDS contents (1 1,500 mg/L) were recorded in groundwater obtained from the Geera Clay overlying basement. Major ion concentrations of groundwater from the Parilla Sand and Renmark Group aquifers are presented for these wells in Figure 3. The Mg/Ca ratio for groundwater in the Parilla Sand aquifer indicates that the molar ratio is <1 in Jungkum and Gerang Gerung. At Jilpanger & Natimuk the ratio is also low, although groundwater from these locations is more saline. At Quantong and, in particular, Duchembegarra, the Mg/Ca ratio increases significantly, as TDS increases. The Mg concentration in groundwater at Duchembegarra is significantly higher than waters from the same aquifers. The high variation in Mg concentrations over such short distances cannot be explained simply by water-rock interaction, as there are no major changes in aquifer mineralogy. The likely source of the Mg is from the
salt lakes within the Douglas Depression that have Mg concentrations up to 23,400 mg/L (Brownbill 1995). The Mg concentration within the salt lakes generally increases from south to north along the depression (RWC 1990). The increase in Mg proportionally to Ca in these salt lakes may be due in part to Ca loss by precipitation of gypsum, and also because of a potentially increased mobility of Mg due to its smaller ionic radius. Major ion concentrations for waters within the Renmark aquifer (Fig. 3) indicate lower TDS contents outside the depression and upgradient of the salt lakes (Jungkum & Natimuk), compared to those wells within the depression and downgradient of the salt lakes (Quantong, Duchembegarra & Gerang Gerung). When Mg/Ca ratios and TDS contents from the Parilla Sand aquifer are compared, it is clear that water quality in the Douglas Depression is compromised throughout the entire system. This contrasts greatly with higher water quality in the Renmark Group aquifer in areas outside and upgradient of the depression. This drastic variation in water quality indicates that reflux of brines from salt lakes into the underlying aquifers occurs in this area, and that localised flow paths exist at Duchembegarra between the Parilla Sand and Renmark Group aquifers which transport brines from the near-surface to depths of over 180m. The Bookpurnong Beds and the lowpermeability layer that includes the Geera Clay, separate the Parilla Sand from the Renmark Group throughout most of the Douglas Depression, including Duchembegarra (Fig. 2). The vertical hydraulic conductivity (K,) of these units ranges from 4.5 x 10-5to 1 x 10-4m/day in the Bookpurnong Beds and approximately 2 x 1O-’ d d a y in the Geera Clay (Lawrence 1975). Although the K,is low, both these units contain significant quantities of sand and silt in places (Brown & Stephenson 1991). These more permeable s a d s i l t zones may enable more rapid fluid transmission between aquifers, particularly where in the absence of obvious fractures within the Bookpurnong Beds and Geera Clay, vertical solute transport appears to be driven by density flow (eg. Wooding et al. 1997). This has implications for the transport of brines from the surface and near-surface deeper into the basin. Such a scenario is supported outside the depression at Gerang Gerung (north) where TDS contents within the Renmark Group aquifer of 10,500 mg/L were recorded in 1999. At this site the TDS contents of the overlying Parilla Sand (950 mg/L) and Murray Group Limestone (1400 mg/L) are significantly lower than the Renmark Group. It appears that solutes in groundwater from the Renmark Group aquifer at Gerang Gerung have not come fiom the overlying units (despite a downward vertical flow) but rather have come from upgradient, most likely
591
the basin, thereby reducing the water quality of the deeper aquifers. Evidence suggests that salt-induced degradation of the groundwater and surface environments within the Douglas Depression is not due to land-clearing practices, but rather a result of long-term hydrogeological processes. REFERENCES
Figure 4. Stable isotopes of *H/"0 for the Parilla Sand and for Renmark Group aquifers within the Douglas three sampling periods during 1999. Note that Gerang Gerung has been sampled only Once (November), and is therefore not included.
from salts entering this aquifer as reflux brines in the Douglas Depression. Stable *H & "0 data from across this area shows changes occurring between formations at individual locations and over time at the same locations (Fig. 4). This eliminates the possibility of leakage down boreholes, which could have accounted for the brines detected at depth in the central depression. Wells within the central depression (Quantong, Duchembegarra and Natimuk) show a clear evaporative signature in the Parilla Sand aquifer in January & June 1999. By November 1999, Quantong and Natimuk values had moved back towards the meteoric water line (MWL), whereas Duchembegarra showed no movement. All wells maintained consistent groundwater chemistry throughout the year, suggesting that isotopic variation is also not a result of leakage down boreholes, or recent changes in groundwater chemistry.
Anderson, J.R. & A.K. Morison 1989. Environmental consequences of saline groundwater intrusion into the Wimmera River, Victoria. BMR Journal of Australian Geology & Geophysics. 11: 233-252. Brown, C.M. 1989. Structural and stratigraphic framework of groundwater occurrence and surface discharge in the Murray Basin, southeastern Australia. BMR Journal of Geology & Geophysics. l l : 127-146. Brown, C.M. & A.E. Stephenson 1991. Geology of the Murray Basin, Southeastern Australia. Canberra, Australia, Bureau of Mineral Resources, Geology & Geophysics. Brownbill, R.J. 1995. The hydrology and chemistry of shallow groundwater and associated lake systems in the southern Wimmera area, Victoria, M.Sc. (unpublished), School of Earth sciences, university of Melbourne, Australia, Coplen, T.B. 1988. Normalization of oxygen and hydrogen isotope data. Chem. Geol.: Isotope Geoscience Section. 72: 293-297 Lawrence, C.R. 1975. Geology, Hydrodynamics and Hydrochemistry of the Southern Murray Basin, Geological Survey of Victoria. McAuley, C. et al, Rural Water Corporation, Victoria (1992). Horsham Hydrogeological Map. (1 :250,000 scale), Australian Geological Survey Organisation, Canberra, Australia. Rural Water Corporation (RWC), Victoria. 1990. Mallee/Wimmera Regional Groundwater Investigation. Unpublished. Wooding, R. A., Tyler, S.W. & I. White 1997. Convection in groundwater below an evaporating salt lake: 1. Onset of instability. Water Resources Research 33: 1199-1217. Wooding, R. A., Tyler, S.W., White, I. & P.A. Anderson 1997. Convection in groundwater below an evaporating salt lake: 2. Evolution of fingers or plumes. Water Resources Research 33: 1219-1228.
4 CONCLUSIONS Although situated in a regional recharge area for the Murray Basin, the hydrodynamics in the Wimmera Region are clearly influenced by local topographic controls. Highly saline water (brine) is able to travel from the surface or near surface to depth (-180 m) relatively rapidly (by density flow) and over short distances (<30 km) due to brine reflux. As the Wimmera Region is located on the margin of the Murray Basin, this has implications for the transport of large quantities of salts further (and deeper) into 592
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Sulfide-free and sulfide-bearing waters in the Northern Apennines, Italy L.Toscani & G.Venturelli Dipartiinento di Scienze della Terra, Universita di Parma, Italy
ABSTRACT: The investigated waters include Ca-carbonate, Ca-sulphate, Na-carbonate and Na-chloride types. Ca-carbonate waters, which circulate in fractured calcite-bearing quartz-feldspathic sandstones outcropping along the Apennine watershed, are dominated by Ca-carbonate dissolution; Ca-sulphate waters are always related to gypsum-bearing lithotypes; the Na-carbonate, strongly reduced group is involved in intensive interaction with clay minerals dominated by Na-Ca exchange and silicate dissolution, and by 0 2 and CO2 progressive consumption. The Ca-sulphate waters from Tabiano baths have unusual composition among the investigated S(-11)-bearing waters: low pH and very high S(-11) content. In these waters, there is bacteriamediated reduction of sulphate, which is derived from Ca-sulphate minerals.
1 INTRODUCTION The investigated springs and wells are located in the Parma and Reggio Emilia Provinces (Northern Italy) along the NE side of the Apennine belt. The springs and wells issue f r o d o r are located in several geological formations, which, on the basis of their petrographic characters, may be grouped as follows: (1) quite permeable quartz-feldspathic sandstones with interbedded pelites (Eocene-Miocene), which mostly outcrop along or close to the Apennine watershed; (2) poorly permeable to strongly fractured carbonate, marly to silty-pelitic formations usually strongly folded and faulted (Cretaceous to Pliocene); (3) strongly permeable (karst processes) evaporites (Trias). 2 HYDROGEN AND OXYGEN ISOTOPES The lj2H and 6l80 values of the investigated waters cover a narrow range (-67 to -58%0, and -9.7 to-8.l%0 V-SMOW, respectively) although they may vary significantly at the different places. These values are similar to those found in springs of the Northern side of the Piemonte-Ligurian Alps, NW of Genoa (mostly in the range of -69/-58%0 and -9.W-8.2%0 respectively; Marini et al. 2000a) and close to the values reported in a preliminary isotopic map of the precipitation in Italy (Longinelli et al. 2000). Only a
water from a well (St. Andrea Bagni) exhibits unusual values (F2H = -47.5%0, 6"0%0 = -4.9) and this will be discussed later. The only station of isotope monitoring close to the investigated area is San Pellegrino in Alpe, which is located on the opposite (WSW) side of the Apennine chain (1520 m above sea level). For this station, the meteoric water line (SPL in Fig. lA, IAEA data) is 62H = (8.28k0.24) 6"O + (15.2rt1.8) (1993- 1995) and, on average, 62H = -47, and 6 ' * 0 = -7.6. In Fig. lA, this line is reported together with the data obtained on waters and two snow samples from the investigated area. With only evident exception of a well located at the St. Andrea Bagni, the investigated waters fall along SPL. Thus we can conclude that the data obtained on the studied springs and wells agree with a dominant contribution of recent meteoric waters. S(-11)-free springs sampled along the Apennine watershed exhibit positive correlation for 62H and 6l80 vs elevation of the springs. The obtained gradient (A&%, / Aelevation (hm)) is about -1.1 for tj2H and -0.18 for 6l80. These gradients found for the springs suggest that altitude effect on rain may be expected in the Northern side of the Apennine chain. On the contrary, the correlation has not been found for the S(-11)-bearing waters, that is in agreement with a more complicated water recharge and circulation at depth.
593
3 CHEMISTRY OF THE WATERS AND WATERROCK INTERACTION
Table 1. Average compositions of the waters (mg/L). Groups S(-II)-jree (1) Ca-carb
The investigated waters exhibit wide chemical variations. Average values for the main types of waters are reported in Table 1.
(2) Ca-carb (3) Ca-sulph
3.1 S(-II)-Ji.eewaters
S(-II)-bearing (4) Na-carb
Group (1) (Table 1). These Ca-carbonate waters interact with fractured quartz-feldspathic sandstones (quartz, plagioclase, K-feldspar, muscovite, biotite, chlorite, calcite, clays) with interbedded pelites (illite, chlorite, f montmorillonite). They are mostly undersaturated with respect to calcite (saturation index, SIcalcite = - 1.9 to 0.13) and exhibit moderate calculated Pco2 (43 to 272 pascal). C1 and Na are comparable with those found in rain waters. The Cdalkalinity ratio (equivalents) around unity (1.02) indicates that calcite is practically the only source of calcium and alkalinity for these waters which circulate rapidly. Group (2). It also consists of Ca-carbonate waters coming from calcarenites, calcilutites, marly limestones with abundant pelitic component (illite, illite/smectite, chlorite). The C1 values resemble those of group (l), but Na is higher; moreover Sr may be very high (up to 2.5 mgk). Saturation to oversaturation in calcite is the rule (SIcalcite up to 0.99). The average Cdalkalinity ratio is lower (0.91) than in group (1) suggesting incipient ion exchange in presence of calcite and/or silicate dissolution (see below for detail); in particular, four samples of these groups coming from a narrow area, exhibit negative correlation between Ca and alkalinity and positive between Na and alkalinity, in agreement with the first hypothesis. With respect to group (l), group (2) indicates stronger water-rock interaction which may be caused both by slower circulation and to lower average grain size of the rocks involved. Group (3). The Ca-sulphate waters interact with Triassic evaporites containing gypsum (634S = 15.5%0, V-CDT), anhydrite, thin dolomitic levels and limestones; the rocks may be very permeable because of developed karst processes. The waters are rich in Sr (up to 10.8 mg/L), saturated in calcite and in gypsum and have higher Mg than the previous groups. C1 and SO4 are roughly correlated suggesting that chlorides are still present at depth in the Triassic evaporites.
(5) Ca-sulph (6) Ca-carb (7) Na-chl
HRD
pH
74 39 192 57 1140 724
7.65 0.32 7.82 0.28 7.78 0.16
alk HC03 74 38 214 70 92 26
2.6 0.8 3.2 1.5 5.6 1.7
11.2 9.4 43 39 1067 723
50 (64) 2005 141 248 148 429 (724)
8.97 0.56 6.33 0.16 7.36 0.17 8.53 0.84
515 134 612 72 391 38 406 308
49 (84) 382 (399) 9.5 8.0 3733 (4990)
23 23 1824 389 48 31 74 (84)
Groups S(-II)-free (1) Ca-carb
H2S
Ca
Mg
Na
K
nd
(2) Ca-carb
nd
(3) Ca-sulph
nd
25 12 62 17 372 230
2.9 2.2 9.0 5.2 52 41
2.6 2.0 17 (23) 6.9 4.7
0.6 0.3 0.9 0.6 1.2 0.7
15 (21) 627 27 84 19 93 (161)
3.2 3. I 107 49 26 14 44
207 98 437 (418) 34 24 2472 (2894)
2.6 (3.7) 19 10 2.1 1.6 21 (25)
S(-II)-bearing (4) Na-carb (5) Ca-sulph (6) Ca-carb (7) Na-chl
2.0 (2.2) 162 59 0.5 0.4 3.3 (3.8)
(81)
For the definition of the groups, see text. In Italics, standard deviation; values in bracket are only indicative of wide dispersion. HRD = hardness as CaC03; nd = not detected.
3.2 S(-II)-bearing waters
Group (4). The Na- carbonate waters issue from several lithotypes such as calcarenites, calcilutites, marly limestones (all with abundant illite, illitehmectite, chlorite) and quartz-feldspathic sandstones with abundant pelites. They may exhibit very low Ca, Mg and calculated Pco2 values (down to 1.2, 1 mg/L and 3-5 pascal respectively) and high Na, B, F, alkalinity (419, 4.9, 3.1 and 723 m g k respectively) and pH (9.4). In spite of the frequently low Ca content, these waters are saturated to oversaturated in calcite (SIcalcite 0.19 to 1.09). The Na character may be acquired both by Na-
594
Ca(+Mg?) ionic exchange, involving clay minerals, and dissolution of Na-bearing silicates. If only ion exchange occurs in the system, we have:
decrease and make possible sulphate reduction mediated by bacteria. Reduction of gypsum by CH4 may be described by the general overall reaction:
(i) d alkalinity = dNa + dCa = 0 (equivalents), where dCa is negative. If the solution is saturated in calcite and, as ionic exchange proceeds, calcite is dissolved to maintain equilibrium, we may write: (ii) d alkalinity = dCai, = dNa + dCa > 0
(iii) CaS04 2H20 + CH4 + CaC03 + H2S + 3H20
It is noteworthy that in the case (ii) is dCai, < -dCa since introduction of Ca in amount minor than that exchanged is sufficient to maintain the calcite equilibrium. Thus, the ionic exchange progress would lead to alkalinity increase and Ca decrease in the solution. Actually, the distribution of the Nacarbonate and Ca-carbonate H2S-bearing waters in Fig. 1B do not follow a well defined trend; however, although they are scattered, the upper limit of their distribution has a negative slope in agreement with different parental waters and ionic exchange. Ion exchange is also supported by the positive correlation between B and NdCa; actually, B may be desorbed during interaction between clays and waters. Dissolution of Na-bearing silicates may also increase the the Na/Ca ratio and decrease the Ca content (calcite precipitation) as modelled, for instance, by Marini et al. (2000b) for waters of the Bisagno valley, Genoa (Italy). It is noteworthy, however, that our waters exhibit very low Ca and unusually high NdCa atomic ratio (about 460) which were never reached in the model. Summing up, it may be concluded that the genesis of the Nacarbonate waters may be related both to ion exchange and silicate dissolution during rather long water-rock interactions.
Group (5). The Ca-sulphate waters of Tabiano baths (Toscani et al., 2001) are present in clastic, conglomeratic to marly-pelitic sediments of Messinian age. These rocks contain some gypsum, sulphur and organic matter, and are characterised by low to moderate permeability. The waters are rich in Li and Na (on average 0.77 and 437 mg/L respectively) and S(-11) (162 mg/L as H2S) and have low pH and high calculated Pco2 (13-32 kpascal). Moreover, they are saturated in gypsum and calcite (SIgypsum E 0.0, SIcalcite E -0.2 to 0.0). It is noteworthy that Tabiano is located in a belt where surface emissions of methane and other hydrocarbons are common. Also very small amounts of dissolved CH4 may drastically contribute to Eh
which would proceed to the right also at very low PCH, . The Tabiano sediments contain organic matter (about 1% in similar sediments: Mattavelli et al. 1983), which is very probably responsible for the high N(-111) content in the waters (up to 9.6 mg/L as N h ) . The organic matter may be the substratum for microbial-mediated reduction of sulphate according to the overall reaction: (iv) CaS04 2H20 + 3H20
+ 2CH20 + CaC03 + H2S + CO2
where CH2O is indicative of organic matter. The 634Svalue in the Messinian sulphate drilled in the Mediterranean sea is on average 21.5%0,in the Messinian Ca-sulphate from Tuscany (Dinelli et al., 1999) 20.0%0f2.8%0(standard deviation) and in the Messinian sea water about 22.2+1%0(Claypool et al. 1980, Pierre & Fontes 1978, Ricchiuto & McKenzie 1978). On the basis of these data and assuming that the total sulphur dissolved in the waters from Tabiano originates only by dissolution of sulphates, the expected mean value of 634S in the Tabiano sulphate phases (634S) can be easily calculated by mass balance assuming rock + water as a closed system for sulphur:
634sG (15.01x634ss(+vI)+ 2.99x634Ss(-II))/18.00= 19.9%0 where 6 3 4 S ~ ( (= + ~27.3%0, ~) V-CDT) and 6 3 4 S ~ ( (= -~~) -17.1%0) are the values for the dissolved sulphate and sulphide in the analysed water, and 18.00, 15.01 and 2.99 are total sulphur, S(+VI) and S(-11) concentrations (mmol/L), respectively. The calculated value is very close to the average obtained by Dinelli et al. (1999), apparently suggesting reduction of sulphates from Messinian sediments.
Group (6). The Ca-carbonate waters also issue from different lithotypes including calcarenites, calcilutites, marly limestones (all with abundant illite, illite/smectite, chlorite) as well as quartzfeldspathic sandstones with abundant pelitic component. This group has higher Sr in respect to the Na-carbonate type (1.4 against 0.44 mg/L). Waters of this group interacting with sandstones chlorite/vermiculite have high Mg (34-39 mg/L). 595
ACKNOWLEDGEMENT Financial support was from CNR, grants 98.00242.CT.05 and 98.03337.ST74 and from the University of Parma.
REFERENCES
Figure 1. A: cross = snow, diamond = S(-11)-bearing waters, square = S(-11)-free waters. B: filled symbols = S(-11)-bearing waters; open symbols = S(-11)-free waters. Circle = Cacarbonate, triangle = Ca-sulphate, square = Na-carbonate, diamond = Na-chloride (Eq = equivalent)
Group (7). The rare Na-chloride waters form a very heterogeneous group coming from sandstones of variable mineralogical composition. These waters are always saturated to slightly oversaturated in calcite and exhibit variable PCO2. An investigated water, which comes from a well drilled in Middle Miocene gadoil-bearing conglomeratic to pelitic sediments (St. Andrea Bagni) has peculiar features. It exhibits high salinity (x18 g/L TDS), Cl- (10.8 g/L), Na (6.4 g/L), B (46.4 mg/L), Li (1.9 mg/L), Sr (27.5 mg/L), NdC1 atomic ratio (0.91) very similar to that of the sea water, and high 62H (-47.5%0)and 6’*0(-4.9%0)(V-SMOW). Mixing between meteoric and “formation waters”, possibly related to evaporite environment, is a reliable hypothesis for the genesis of this samples. Saline waters are common in the area and used for curative purposes since the Roman age (e.g. the Salsomaggiore baths).
Claypool, G.E., Holser, W.T, Kaplan, I.R., Sakai, H. & I. Zak 1980. The age curves of sulphur and oxygen isotopes in marine sulphate and their mutual interpretation. Chemical Geology 28: 199-260 Dinelli, E., Testa, G., Cortecci, G. & M. Barbieri 1999. Stratigraphic and petrographic constraints to trace element and isotope geochemistry of Messinian sulphates of Tuscany. Mem. Soc. Geol. It. 54: 61-74. Longinelli, A, Selmo, E. & 0. Flora 2000. Isotopic composition and tritium activity of atmospheric precipitation in Northern Italy. 5* Intern. Isotope Workshop, Cracovia , July 2000. Marini, L., Bonaria, V., Guidi, M., Hunziker, J.C., Ottonello, G. & M. Vetuschi Zuccolini, 2000a. Fluid geochemistry of the Acqui Terme-Visone geotermal area (Piemonte, Italy). Appl. Geochem. 15: 917-935 Marini, L., Ottonello, G., Canepa, M. & F. Cipolli, 2000b. Water-rock interaction in the Bisagno valley (Genoa, Italy): application of an inverse approach to model spring water chemistry. Geochim. Cosmochim. Acta 64: 26 17-2635. Mattavelli, L., Ricchiuto, T., Grignani, D. & M. Schoell 1983. Geochemistry and habitat of natural gases in PO basin, Northern Italy. Am. Ass. Pet. Geologists Bull. 67: 2239-2254 Pierre, C. & J.C.Fontes, 1978. Isotope composition of Messinian sediments from the Mediterranean sea as indicators of paleoenvironments and diagenesis. In Initial Report of the DSDP 42.1, pp. 635-650. Ricchiuto, T.E. & J.A. McKenzie, 1978. Stable isotopic investigation of the Messinian sulphate samples from DSP Leg 42A, Eastern Mediterranean sea. In Initial Report of the DSDP 42.1, pp. 657-660. Toscani, L., Venturelli, G. & E. Savini, 2001. Geochemical features of H2S-bearing waters of the Tabiano baths, Parma province, Italy. Quaderni di Geologia Applicata 8: 1-8.
596
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Elevation, landuse and water-rock interaction effects on groundwater quality S .Tweed, T.R .Weaver & G.P.Masur Hydrogeology and Environnzent Research Group, School of Earth Sciences, University of Melbourne, Victoria, 301 0,Australia.
I .Cartwright Hydrogeology and Environment Research Group, Department of Earth Sciences, Monash University, Clayton, Victoria, 3800, Australia
ABSTRACT: Determining processes controlling groundwater chemistry is essential for effective water resource management and protection. The Yarra Catchment is an important water resource in SE Victoria, Australia. This catchment is characterised by fractured rock aquifers; a high baseflow component to rivers; high, localised recharge; and diverse land use. Here, we identify processes controlling groundwater chemistry through stable isotope (0 and H) and major ion geochemistry, Landsat imagery, and geochemical modelling of groundwater in the Dandenong Ranges sub-catchment. Various stages of groundwater flow are identified. At high elevations, Na-Cl molar ratios (1.0) and low TDS values (-50-100 mg/L) indicate direct recharge of precipitation into sedimentary or granitic aquifers. Where precipitation directly recharges younger basalt aquifers TDS values increase (-100-150 mg/L). At lower elevations, Na-Cl ratios decrease, reflecting an increase in C1 and a decrease in Na. In these areas, higher major ion concentrations (TDS - 350-450 mg/L) and pH values reflect cumulative effects of water-rock interaction and agricultural land use lower in the catchment. ment. Both natural and human-induced processes affecting groundwater chemistry in the catchment are investigated with the intention of understanding areas where groundwater, and inevitably surface water, are susceptible to contamination. Groundwater flow paths from areas of recharge to discharge
1 INTRODUCTION The Yarra Catchment (4045 km2) is a significant area for water supply in Melbourne, SE Australia (Port Phillip Catchment and Land Protection 1999). Within the Yarra Catchment, most of Melbourne’s water supply is collected in reservoirs (Fig. 1). Groundwater use in the catchment includes domestic and stock, irrigation, dairy and industry. The geology of the Yarra Catchment forms an open and interconnected multilayered fractured rock aquifer (Shugg 1996): major aquifers include Silurian to Devonian siltstones and sandstones, Devonian acidic volcanics, and Tertiary Older Volcanic basalts. Silurian to Devonian marine sediments underlie most of the Melbourne region and outcrop mainly to the north and east. During the Devonian, emplacement of acid intrusive and volcanic rocks occurred throughout the region. These volcanics now form the highest erosional surfaces in the Yarra Catchment. During the Tertiary, extrusive Older Volcanic basalts infilled ancient streams of dissected Silurian to Devonian bedrock. These basalts presently outcrop in elevated areas to the east, where basalt capped ridges occur. Groundwater vulnerability is considered in this study because of the estimated baseflow of over 50% to surface waters (Shugg & O’Rourke 1995); a lack of investigation into groundwater quality; and intensive and diverse landuse in the Yarra Catch-
Figure 1. Location of Dandenong Ranges sub-catchment, reservoirs and surface water of the Yarra Catchment (after Port Phillip Catchment and Land Protection 1999).
597
Table 1. Groundwater Chemistry of the Dandenong Ranges, winter 1999.
1. Recharge, granite, sedimentary
3.Intermediate, 4.Intermediate, 5. Discharge, 6.Irrigation,
2.Recharge, basalt
mix
Groundwater 231 elevation (m AHD) 225 235 241 243 Temperature ("C) 1 1.7 13.9 12.1 13.7 16.8 PH 5.82 5.18 5.55 6.29 6.54 EC (p/cm) 112 124 136 176 222 60 147 Eh (mV)* 234 285 239 Ca ( m g U 3.5 2.6 2.0 9.5 11.7 7.8 10.2 Mg ( m g m 2.5 1.7 2.9 15.2 19.1 Na ( m g m 16.0 19.5 20.0 2.8 2.9 K (mgk) 2.0 0.5 1.2 24.7 30.4 Si (mg/L) 9.8 4.2 5.3 1.7 1.8 2.1 so4 (mgk) 1 .o 3.5 c1( m g U 21.8 32 20.8 35 32 HC03 (mg/L) 13 6 31 10 39 SI SiOz (a) -1.03 -1.40 -1.30 -0.63 -0.54 SI Albite -5.51 -9.30 -9.31 -2.69 -1.58 $51 CO2 (8) -1.72 -1.18 -1.77 -1.82 -1.97 SI = saturation index calculated using PHREEQCI (Charlton et al. *Eh measured relative to Ag/AgCl electrode.
225 170 161 15.0 13.8 13.9 8.20 6.23 6.68 304 317 407 40 196 -13 18.6 5.8 25.1 6.1 9.3 8.4 38.1 45.0 44.4 2.0 1.7 4.3 7.0 42.0 17.3 4.6 1.6 1.6 40 59 55 52 37 70 -1.19-0.40-0.79 -2.37-2.07-1.49 -3.52-1.69-1.87 1997).
in the Dandenong Ranges sub-catchment (Fig. 1) are investigated to aid assessment of primary controls on the groundwater chemistry on both local and regional scales.
basalt
220 16.5 6.41 437 116 22 27.2 28.8 2.6 27.7 0.9 57.3 82 -0.58 -4.18 -1.53
sedimentary
basalt
168 13.0 7.05 741 35 25.1 34.2 54.4 2.2 25.8 0.1 98 168 -0.61 -3.32 -1.90
226 13.5 7.92 700 -66 39.0 34.4 59.2 3.8 23.6 5.4 106 131 -0.66 -0.61 -2.86
Investigation of spatial and temporal landuse changes of the Yarra Catchment involved the use of Landsat datasets flown in the summers of 1980, 1988 and 2000. Datasets were processed and classified using ERDAS Imagine 8.4. Preliminary results indicate that Landsat imagery is useful in locating the more fertile and geochemically reactive Older Volcanics basalt soil compared with the less-fertile and more aerially extensive outcrop of the Silurian and Devonian sandstone and siltstone soil.
2 FIELD INVESTIGATIONS Results of field investigations from the Dandenong Ranges for winter 1999 are shown in Table 1. Results are presented in groundwater flow path groups defined by elevation, geology and groundwater chemistry (Fig. 2). Temperature, pH, electrical conductivity (EC), redox (Eh), and alkalinity (as HC03-) were measured in the field. Samples for cation analyses were filtered and acidified in the field and analysed by ICP-OES at the University of Melbourne. Anions were analysed by ion chromatography at the Wairakei Research Centre, NZ, and stable isotopes were analysed using a Finnigan MAT 252 mass spectrometer at Monash University.
3.2 Origin of groundwater Groundwater samples were analysed for 6l80 and 62H to determine the origin of groundwater and processes occurring during groundwater flow. 6l80 and 62H values (-6.2 to -5.5 %o, and -38.1 to -32.2 %O respectively) lie on the Melbourne Meteoric Water
3 DISCUSSION 3.1 EJjcects of elevation and landuse on groundwater chemistry Changes in surface elevations throughout the Yarra Catchment result in spatial variations in groundwater chemistry. Spatial heterogeneity on a local scale in the Dandenong Ranges is the result of localised recharge events in areas of high surface elevations, and possibly changes in land use affecting local recharge. Groundwater in these areas is fresher compared to groundwater in lower elevations that are further from the recharge source. This relationship is also observed throughout the Yarra Catchment.
Figure 2.Flow path groups for groundwater in the Dandenong Ranges: 1. 0 Recharge sedimentary, 2. Recharge basalt, Intermediate mix, 4.A Intermediate basalt, 5.0Discharge sedimentary, 6.A Irrigation basalt.
3.e
598
Line (MWL) (Fig. 3). This indicates that groundwater has a meteoric origin, and has been subjected to neither evaporation nor high-temperature waterrock interaction. Stable isotope ratios in groundwater are slightly lower than those in nearby surface water. Suggesting that either the groundwater has long residence times reflecting recharge under different climatic conditions, or that seasonal rechar e to the groundwater system may occur. 6l80 and 6#H values from the IAEA database (IAEANirMO 1999) show a large range of precipitation values for the Melbourne region during the years 1995 to 1997 (Fig. 3). These values indicate that isotopically depleted rains are not seasonally dependent. However, seasonal biases to recharge are likely to exist because of Melbourne’s temporal climate, where most rainfall occurs during late winter and spring.
3.3 Groundwater flow paths The 6 flow path groups identified (in Fig. 2) do not
Figure 3. Stable isotopes of groundwater, surface water and precipitation in the Dandenong Ranges, illustrating a range of groundwater values, higher surface water values and a meteoric origin for groundwater: 0 groundwater, 0 surface water, Dandenong Ranges precipitation, +IAEA summer precipitation, A IAEA autumn precipitation, 0 IAEA winter precipitation, and IAEA spring precipitation.
-
aquifers have weathered zones extending 20 to 50 m deep. In areas where recharge is via basalt ridges, infiltration of rainfall is rapid and weathered basalt zones extend to depths of 30 to 80 m. Kaolinite is the dominant clay for both units in the upper weathered profile. Smectite-montmorillonite dominates the base of the weathered basalt profile and illite dominates the base of the weathered sedimentary profile (Shugg 1996). Due to the relatively high cation exchange capacities of illite (200peq/g) and smectite (1000 peq/g) (Deer et al. 1992),cation exchange is considered a potential process influencing the groundwater chemistry along all flow paths. Influences of cation exchange are observed at the intermediate stages of the groundwater flow paths in the basalt aquifer. Increased residence times (group 4) result in a decrease in Na, Ca and HC03 concentrations, whilst Mg and Si concentrations increase. A significant decrease in Na:Mg ratios from groups 3 to 4 (3:l to 1:2) indicates that cation exchange is most likely occurring. Dilution will also play a role in groundwater chemistry of group 4 due to mixing with groundwater that has been recharged via the sedimentary aquifer. Mixing of groundwaters exposed to different water-rock interactions are also observed in local discharge zones. Located in topographic lows, discharge zones are usually in sedimentary aquifers down-gradient from basalt aquifers. At the major local discharge area in the Dandenong Ranges there are two nested flowing artesian bores (group 5). One represents local discharging groundwater (45 m deep), the other represents regional discharging groundwater (300 m deep). Regional groundwater predominantly interacts with the sedimentary bedrock (TDS = 390 to 460 mg/L), whereas local groundwater flow paths travel mainly via the basalt aquifer (TDS = 360 to 420 m a ) . As a result of the
represent the entire evolutionary trend for groundwater in the Dandenong Ranges. However, they do represent key stages along groundwater flow paths. Major cations proceed from sodium dominated to increased concentrations of Mg and Ca (Na-Mg-Ca). The major anions proceed from C1 to HC03 - C1 dominated groundwater. 1. Recharge via sedimentary or granitic aquifer: short residence times, little water-rock interaction Na=Cl, low TDS (- 50 to 100 mg/L), and intermediate surface elevations. 2. Recharge via basalt aquifer: short residence times, little water-rock interaction, Na=Cl, low TDS (- 100 to 150 mg/L), and high surface elevations. 3. Intermediate residence times for groundwater in basalt aquifers and groundwater travelling through sedimentary aquifers via basalt: Na>Cl, moderate TDS (-150 to 250 mg/L) and surface elevations. 4. Increased residence times for groundwater in basalt aquifer: CbNa, high TDS (-200 to 250 mg/L) and intermediate surface elevations. 5. Local and regional discharging groundwater via sedimentary aquifer: CbNa, high TDS (-350to 450 mg/L), and low surface elevations. 6. Irrigation area, basalt aquifer: CbNa, high TDS (-400 mg/L), and intermediate surface elevations.
3.4 Water-rock interaction along groundwaterjow paths
Investigation of the primary controls, which determine changes in groundwater chemistry along the flow paths, involves the examination of water-rock interactions and groundwater mixing. Within the Dandenong Ranges, recharge areas of sedimentary 599
different flow paths the local discharging groundwater has greater concentrations of Mg and Ca than the regional groundwater. Effects of water-rock interactions at different stages along the groundwater flow paths are also evident from the variable silica concentrations (3.3 to 42.0 mg/L as Si). Contributing to this variation are groundwaters from basalt ridges in local recharge areas, which have deeply weathered profiles (80 m). Weathering of the fresh basalt may result in high Si concentrations because of water-rock interaction between slightly acidic recharging precipitation and olivine ((Mg,Fe)zSiO& pyroxene (e.g. MgSi03), and amphiboles (e.g. Mg7(Si401&(OH)2). Interpretation of potential mineral dissolution or precipitation and ion exchange reactions occurring in the groundwater of the Dandenong Ranges was also investigated through groundwater modelling (PHREEQC (Parkhurst 1995)). PHREEQC output data (Table 1) indicates that all groundwater samples are just undersaturated with respect to amorphous Si02 (Si02(,)). SiOz(,) is considered instead of quartz because in low temperature environments SiOq,) precipitates and dissolves relatively rapidly. SiO2(,) therefore represents the upper limit of dissolved aqueous silica in natural waters (Drever 1997). Accession of ions along the different groundwater flow paths is reflected in some of the saturation indices approaching zero as major ion concentrations increase (e.g. albite). The SI results for CO*(,) (Table 1) indicate that groundwater may be in equilibrium with the highly vegetated soil of the Dandenong Ranges, e.g. soil Pco2 = 10-3to 10" atm (Clarke & Fritz 1997).
for the protection of water quality within this Catchment. REFERENCES Clark, I. & P.Fritz 1997. Environmental Isotopes in Hydrogeology. New York: Lewis. Charlton, S.R., Macklin, C.L. & D.L.Parkhurst 1997. PHREEQCI - A graphical user interface for the geochemical computer program PHREEQC. U.S. Geological Survey Water-Resources Investigations Report 97-4222. Deer, W.A., Howie, R.A. & J.Zussman 1992. An introduction to the rock-forming minerals, second edition. Longman Scientific and Technical and John Wiley & Sons, Inc. Drever, J.I. 1997. The Geochemistry of Natural Waters, Surface and Groundwater Environments, third edition. R.A. McConnin (ed.). Prentice-Hall, Inc. IAEAIWMO. 1999. Global Network for isotopes in precipitation. The GNIP Database. Release 3, October 1999. Laboratory code: 9486800. htm URL: httpllw ww.iaea.org/programs/ri/gnip/gnipmain. Port Phillip Catchment and Land Protection. 1999. Yarra Catchment Action Plan website: http://www.nre.vic.gov.au/catchment/portphillip/yarra/inde x.htm Parkhurst, D.L. 1995. User's Guide to PHREEQC - A Computer Program for Speciation, Reaction-Path, AdvectiveTransport, and Inverse Geochemical Calculations. U.S. Geological Survey Water-Resources Investigation. Shugg, A. 1996. Hydrogeology of the Dandenong Ranges Fractured Rock Aquifers and the comparison with similar aquifers in Victoria. MSc thesis. Sydney: University of Technology (unpubl.). Shugg, A. & M.O'Rourke. 1995. Groundwater in the Yarra Basin. Technical Document. YarraCare, Victoria.
4 CONCLUSIONS Recharge plays a key role in this groundwater system. The regional groundwater flow regime of the Yarra Catchment is superimposed by high, localised recharge in areas of high surface elevations. Stable isotope results indicate a meteoric origin for groundwater. These results also suggest that groundwater was either recharged during different climatic conditions (long residence time), or there has been preferential recharge of cooler rains entering the groundwater system. Groundwater flow path groups show clear distinctions between groundwaters from areas of recharge through to discharge, and reflect differences in surface elevations and waterrock interactions. A high base flow component to streams, coupled with the expansion of intensive land use, contributes to the vulnerability of water supply and quality in the Yarra Catchment. Therefore, identification of local recharge and discharge areas, and the controls on groundwater chemistry of the different flow path groups are important factors 600
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrochemical patterns of the Gavarres hydrological system and its surrounding aquifers (NE Spain) E.Vilanova & J.Mas-Pla Gabinet d’Andisi Territorial i Anzbiental. Unitat d’Hidrogeologia, Dept. de Geologia, Universitat Autdnoma de Barcelona, 08193 Bellaterra
ABSTRACT: The results of hydrochemical analysis in a crystalline massif and its surrounding sedimentary basins show that the dominant chemical process is silicate, primary involving feldspars, and carbonate dissolution enhanced by CO2 atmospheric input, with variations related to local rock lithology. Observed chemical compositions reflect different degrees of water-rock interactions, pointing out the effect of preferential flow through fracture zones and the influence of the recharge area lithologies.
1 INTRODUCTION Hydrochemical studies are of special interest in identifying recharge areas in mountain systems; especially in those where potentiometric data do not provide relevant information about flow-paths. The objective of this paper is to investigate the origin and pathways of groundwater within the Gavarres mountain with the use of geochemical evidence. For this purpose, 36 groundwater samples from wells and 3 from springs are analysed, and major elements used to determine the main hydrochemical processes. Chemical patterns will then be used to identify the recharge areas in this massif.
CO2 rich waters placed in granitic (Sl, S2 and S3) and metamorphic environments (G4 and G5).
3 RESULTS AND DISCUSSION The dominant water type is calcium bicarbonate, except two CO2 springs that are sodium bicarbonate and two deep granitic wells that are calcium chloride-bicarbonate (Fig. 2). Despite their different geological origin, all groundwater samples have similar compositions, mainly due to the reaction of CO2 charged water with silicate or carbonate rocks. However, some differences can be noted based on specific relationships. For instance, silica content is
2 GEOLOGICAL SETTING The Gavarres mountain range is located in northeastern Spain (Fig. 1). It consists of an Hercinian massif with metasedimentary and acid vulcanosedimentary rocks. A large granitic calc-alkaline intrusion appears in its southern part composed basically by quartz, orthoclase, plagioclase (albite and oligoclase) and biotite (RoquC 1993). Sedimentary materials fill the surrounding areas. In the northern part, Paleogene materials (limestones, mark and sandstones) overlay the Palaeozoic succession. As a consequence of the graben structure formed during Neogene, the surrounding basins were filled up with detrital nonconsolidated sediments (being in the southwestern basin of arkosic nature). Quaternary deposits, mainly of alluvial origin, are related to the present drainage system. The regional fractures that limit the Gavarres massif, allow the occurrence of
Figure 1 . Geological map of the studied area.
601
Table 1. Results from chemical analysis and Saturation Indices calculations. Aquifer types are designed by sample numbers: 1-5 C0,rich waters, 6-9 arkosic, 10-16 granitic, 17-20 metamorphic, 21-28 limestones, 29-30 sandstones, 3 1-38 alluvial.
s1 s2 s3 G4 G5 S6 s7 S8 E9 s10 s11 s12 S13 S 14 S15 G16 G17 Gl8 E19 E20 E2 1 E22 E23 E24 E25 E26 E27 E28 E29 E30 E3 1 E32 E33 E34 E35 E36 E37
T
pH
C
CaL+
Mg"+
Na+
K+
Sr"
Si
"C 14.4 14.1 15.4 16.1 11.6 16.7 16.5 15.3 19.2 17.4 15.7 16.7 15.9 17.1 17.5 14.9 15.9 16.8 17.8 17.8 16.2 16.5 16.5 20.7 25.6 17.4 17.6 20.0 16.7 19 20.1 20.7 18.8 17.7 18.8 17.2 15.4 16.7
pS/crn 2150 4750 3230 2060 15 14 91 1 82 1 850 753 1114 846 493 91 1 1672 1385 626 804 820 900 1017 784 849 1110 923 613 806 1403 1444 816 840 1055 804 1099 1265 925 1226 690 1150
mg/l 350.5 303.5 164.3 176.8 248.6 86.7 97 89.6 101.9 104.4 88.2 45.6 9.5.7 120.3 109.1 82.2 85.3 70.9 105.6 82.1 79.9 112.2 165.7 124.8 78 111 155.5 203.1 122.5 118.6 144.1 93.1 137 147 125.6 176.4 97.2 164
mg/l 52.1 133.1 49.5 90.5 27.3 14.2 12.7 9.4 14.0 17.5 20.6 9.3 16.0 42.5 30.0 11.5 24.6 13.6 20.5 26.0 31.0 26.4 14.9 12.4 2.6 11.3 29.4 39.4 11.9 14.0 16.7 20.0 20.2 27.5 16.4 23.7 10.3 41.3
mg/l 53.6 591.4 426.8 137.8 26.8 53.3 54.2 57.6 36.1 85.9 58 47.1 51 71.4 83.8 34.1 41.2 69.2 40 79 30.6 19.3 40.4 45.6 26 33.3 86 44.9 23.4 33.3 54.2 44.8 38.3 59.7 42.4 26 28.7 47.5
mg/l 2.6 5.7 7.2 17.9 1.5 2.6 1.7 1.2 1.o 1.8 3.1 0.8 1.1 2.7 2.2 3.3 1.2 14.2 1.49 1.4 3.O 2.1 1.o 2.6 <2.5 <2.5 8.2 1.9 0.6 1.6 2.0 3.3 1.1 2.9 4.2 1.o 3.7 <2.5
mgll 0.6 10.6 3.6 1.5 0.5 0.4 0.3 0.4 0.4 0.5 0.5 0.1 0.4 0.3 0.6 0.2 0.4 0.6 0.7 0.3 5 .O 0.4 0.5 0.4
mg/l
6.1 6.5 6.3 6.2 6.3 7.1 6.9 7.2 7.3 7.1 7.4 7.5 7.0 7.1 7.2 7.3 6.9 7.4 7.1 7.3 6.9 7.4 7.1 7.3 6.9 7.4 7.1 7.3 7.6 7.4 7.1 7.1 7.5 7.0 7.4 7.2 7.2 7.4
a good indicator of aquifer lithology, showing its maximum concentrations in granitic and arkosic environments. pH is in the range 7-8, except for CO2 rich waters with values from 6 to 6.5 (Table 1). Caland 10-'.7atm. culated pCOz ranges between 10-0.5 Molar proportions of major dissolved ions in granitic areas follow the sequence HC03--Ca2+>Na+>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
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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
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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.
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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
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Y
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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
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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 (<10,000 years old) and they discharge through the sediment with an approximate rate of 0.2 c d y e a r (Eldefield et al. 1999). Furthermore, the sediments at Site 1031 have different physical properties (e.g. high permeability) compared to the other sites that are related to lithology or fluid upwelling (Giambaldo et al. 2000). Smectite and zeolite found in the basal sediments are typically low-temperature hydrothermal minerals compatible with the temperature of the fluid circulating on the flank of the ridge axis (Davis et al. 1997). Stable and radiogenic isotope analyses on the clay fractions are in progress and should provide further constraints on the conditions of the alteration. 4 CONCLUSIONS Smectite and zeolite formation and calcareous fossil dissolution is occurring in the basal sediments at Site 1031. These reactions consume Mg, Fe and Na from hydrothermal fluids and release Ca to the solution. Mass balance calculations show elemental behaviors similar to those calculated from the low-temperature alteration of the basaltic crust (Alts and Teagle 1999). It is difficult to quantify the possible influence of reactions occurring in sediments on the composition of the upwelling fluid. Howerver, the hydrothermal paragenesis described here suggests that the low-temperature phases are compatible with the temperature of the fluid circulating in basement of the ridge axis flank.
REFERENCES Alt J. & D. A. H. Teagle 1999 The uptake of carbon during alteration of ocean crust. Geochimica et Cosmochimica Acta, 63 : 1527-1535. Buatier M., Monnin C., Friih-Green G. and A. M. Karpoff 2001. Fluid-sediment interactions related to hydrothermal circulation in the Eastern Flank of the Juan de Fuca Ridge. Chemical Geol., in press
618
DaesslC, L. W., Cronan, D. .S., Marchig, V. and M. Wiedicke 2000. Hydrothermal sedimentation adjacent to the propagating Valu Fa Ridge, Lau Basin, S . W. Pacific. Marine Geology, 162: 479-500. Davis, E.E., Fischer, A.T., Firth, J.V. et al., 1997. Proc. ODP, Init. Repts, 168: Ocean Drilling Program, College Station, TX. Eldefield, H., Wheat, C.G., Mottl, M.J., Monnin, C. and B Spiro 1999. Fluid and geochemical transoprt through oceanic crust: a transect across the eastern flanc of the Juan de fuca ridge, Earth Planet. Sci. Let.. 172: 151-165. Giambalvo E. R., Fisher, A., Martine, J. T., Darty, L. and R. P. Lowell 2000. Origin of elevated sediment permeability in a hydrothermal seepage zone, eastern flank of the Juan deFuca Ridge, and implications for transport of fluid and heat. Journal of Geophysical Research, 105: 9 13-928. Honnorez, J. 1981. The aging of the oceanic crust. In The Oceanic Lithosphere (C. Emiliani Ed). The Sea, 7: 525-597. Mottl, M. J. & C.G Wheat 1994. Hydrothermal circulation through mid-ocean ridge flanks: fluxes of heat and magnesium. Geochim. Cosmochim. Acta, 58: 2225-2237,
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hyperkarstic phenomena in the Iglesiente mining district (SW-Sardinia) J.De Wade Dipartimento di Scienze della Terra, Via Trentino 51,09127 Cagliari
P.Forti Istituto Italian0 di Speleologia, Via Zamboni 67, 40127 Bologna
G. Pema Studio GEOAMBIENTE, Salita dei Molini 49, 38050 Villauano-Trento
ABSTRACT: The Cambrian carbonatic rocks of the Iglesiente mining district contain important mineral deposits that have led to intense mining activity over the centuries. The same rocks have been affected by polycyclic karst phenomena since the Ordovician. The exploitation voids go down to -250 m, as a result of a pumping system, and have revealed many natural caves. The study of these caves has clearly demonstrated that "normal" karstification is less significant than hyperkarstic dissolution determined by difFerent processes. The caves are filled with sediments, chemical deposits such as calcite speleothems, enrichments of economic minerals (Ag-rich galena, barite, smithsonite, etc.) and crystals (calcite, barite, quartz, etc.). The natural caves of the mines of Iglesiente and their morphologies and fillings related to hyperkarstic phenomena, represent a unique and inestimable scientific resource, that should be adequately protected and valued.
1 INTRODUCTION
Formation (mainly dolomitised rocks) and the San Giovanni Formation (intensely karstified limestones). The Iglesias Group (Middle Cambrian-Lower Ordovician) is divided into two Formations: the Campo Pisano Formation, composed of nodular limestones, followed by a thick succession of shales of the Cabitza Formation. After a long period of continentality and an important tectonic phase (Fase Sarda) the Sea returned to occupy this area with the deposition of the Ordovician conglomerates (Puddinga), followed by Silurian and Devonian sediments. After the Hercynian orogenesis a long continental period started in the region (Carboniferous-Middle Trias), only shortly interrupted in Middle-Triassic and in Paleocene-Eocene times.
In Iglesiente (SW Sardinia) the Cambrian carbonatic rocks contain deposits of galena, sphalerite, pyrite and barite of sedimentary genesis. These mineralisations have been subjected to re-circulation phenomena, hydrothermalism, alteration and reconcentration during the complex geological history that has affected this area. The same carbonatic rocks have been involved in several karstic cycles since the Ordovician. Each cycle can be independent or involve earlier ones.
2 GEOLOGY OF IGLESIENTE The Palaeozoic rocks are described by Bechstadt and Boni (1996). From a stratigraphic point of view the Cambrian succession in Southwest Sardinia is divided in three major Groups: Nebida Group, Gonnesa Group and Iglesias Group (Fig. 1). The Nebida Group (Lower Cambrian) is composed of delta and coastal sediments and is divided into two Formations: the Matoppa Formation (sandstones and shales) and the Punta Manna Formation (oolithic limestones and calcareous sandstones followed by sandstones with carbonatic fossiliferous lenses and strata). The Gonnesa Group (Lower-Middle Cambrian) is characterised by typically carbonatic deep-sea sediments and is divided into two Formations according to the trilobite content: the Santa Barbara
3 HYDROGEOLOGY
The exploitation of the metal deposits started in prehistoric times, escalating during Pisan domination and in the last 150 years. The most important mines are fi-om S to N: Barega, Seddas Moddizzis, S. Giovanni, Campo Pisano, Monteponi, S. Marco, Masua, Acquaresi (Fig. 1). In 1870 the exploitation that started at the outcrops and deepened fast, reached the aquifer at level +74 m. To be able to proceed with the mining a 7 km long tunnel was constructed that, starting at the 619
Karst aims
'
5km
~
rn
Baregu
Figure 1. Location map and generalised stratigraphic column of the Cambro-Ordovician sequence.
one that hosts the metal deposits that represents a more superficial aquifer. The connection between these two karstic aquifers is established along a NS fault in the Monteponi area (Perna 1995, 1998).
swamp of Sa Masa close to the Sea, reached Monteponi at level + 13 m. This tunnel, excavated in impermeable rocks, reached the carbonatic rocks and cut into a large and important karst structure (Grande Sorgente) from which a great amount of water drained, lowering the phreatic water level in the whole mining district. Later on, once the mine exploitation reached this new water level, a pumping station was placed at - 1 5 followed ~ by installations at -50m, -100 m and -200 m at Monteponi and -250 m at S. Giovanni, with an increasing (up to 2000 Ys) amount of water evacuated (Civita et al. 1983). The lowering of the karstic phreatic level reached the Acquaresi mine situated about 9 km to the NW, and the Barega barite mine distant 6 km to the SE. The initially fresh water, with the progression and the deepening of the mine, became more and more salty due to sea water intrusion, with a complex chemical composition and the increased concentration of heavy metals (Table 1). The most recent hydrogeological studies have led to the hypothesis of a sea water intrusion through a very deep Messinian karstic system, seated in a Cambrian carbonatic seqhence separated fiom the
4 CHEMISTRY OF HYPERJSARST From a chemical point of view the hyperkarstic phenomena that characterise the carbonat ic sequences of Iglesiente are extremely variable and complex, due also to the great geochemical variability of the different fluids that circulate or have circulated in the past; schematically these fluids can be subdivided into: - Thermal fluids of medium-high temperature (rich in metallic ions); - Thermal fluids of low temperature (rich in CO2 and/or HZS); - sea water (high salinity and rich in organic matt er) ; - Meteoric water of rapid infiltration (rich in oxygen and in organic matter). Separately, these fluids have been or are still responsible for formation of caves and speleothems. For example the fluids of mediumhigh temperature (Forti 1996) have been important and active in the depositional process and the following remobilization of orebodies inside preexisting karstic conduits in more or less distant geological times (Ludwig et al. 1989). The meteoric fluids, rich in oxygen, have enabled and still enhance the oxidation of sulphide deposits with the consequent activation of a complex series of chemical reactions related to the sulphur cycle (Forti & Perna 1985, Hill & Forti 1997). These last fluids lead to the evolution of a whole series of accelerated corrosion morphologies and, in the mean time, cause the deposition of a large number of secondary mineralisations, which nature
Table 1. Chemical analysis of the discharged mine water at the end of the tunnel near Sa Masa (Fontanamare, Iglesias, 1996).
620
INFILTRATION OF METEORIC WATER
'''
INFILTRATION OF METEORIC WATER
i4+444
"+
Solutionpans
Travertines
HYPERKARST BY STRONG ACIDS
,--heric annstomrsed
Percolation Zone Epiphreatic (Oscillation: Zone
Saturated
(phreatic)
HYPEFXARST BY MIXLNG
Zone
I
Figure 2. Theoretical section of the Iglesiente karstic massif with the main observed karstic and hyperkarstic mechanisms and their relative corrosion andor depositional morphologies.
and type essentially depend on the particular local chemistry of these fluids and on the composition of the bedrock which these are in contact with. More often though the speleogenetic and depositional hyperkarstic effects are caused by the mixing of two or more different fluids. In fact, from time to time, the mixing of meteoric waters with thermal or marine fluids, or of thermal fluids with marine waters or of all three of these fluids can respectively lead to the following hyperkarstic reactions: simple undersaturation and consequent karstification, simple oversaturation followed by deposition of speleothems, and concomitant corrosion and deposition when the phenomenon of incongruent dissolution is primed (Forti 1996). To make the situation even more complicated one must keep in mind that the hyperkarstic reactions that effectively can take place don't only depend on the chemistry of the fluids and of the rocks, but also on the area of the aquifer in which these processes have to take place. In fact, many of the hyperkarstic phenomena are specific of a well determined hydrogeological situation and, therefore, can manifest only in some parts of the karst aquifer (percolation, epiphreatic (oscillation) and phreatic (saturation) zone). It is not possible in this work, to describe in detail all the hyperkarstic active processes of the Iglesiente rnining district. These can be schematically
subdivided into four groups on the basis of the dominant processes: - Hyperkarst by strong acids (biochemical reactions of the sulphur cycle, active in all the areas of the aquifer, but prevailing in the percolation (aerate) zone); - Thermal hyperkarst (processes due to the rising of thermal fluids, almost exclusively active in the saturated (phreatic) zone and on the interface with the epiphreatic (oscillation) zone); - Hyperkarst by mixing (reaction connected to the mixing of sea water with meteoric infiltration water, essentially on the interface between phreatic and epiphreatic zone); - Hyperkarst by ionic diffusion (redox reactions andor double exchange exclusively in the phreatic zone). The principal corrosion and/or depositional morphologies observed in the Iglesiente district and correlated to the different types of hyperkarstic reactions have been assembled and schematised in Figure 2.
5 EXAMPLES OF HYPERKARST PHENOMENA In the mined areas of Iglesiente many natural karstic caves have been encountered and the main underground springs are related to these (Forti & Perna 1982). All are the product of different karstic
62 1
cycles dating from Ordovician to Recent times. In the Monteponi mine, the water inlets in the higher part of the deposit are correlated to still active karstic phases with traditional morphologies and speleothems, while in the deeper areas there is the rising of thermal waters, rich in salts, that remobilise and enrich the metallic ores. Also in the Nebida and Masua mines to the north traditional karsts are present in the upper levels, while in depth these cavities are related to hyperkarstic reactions. In these deep karst structures deposition of calcite macro crystals and intense corrosion phenomena through percolation of acid fluids deriving from the alteration of sulphide deposits can be observed (Phaff caves at Masua) (De Vivo et al. 1987). One of the main characteristics of this area is the presence of large cavities filled with residual clayey sediments (Bini et al. 1988). At Buggerru, in the extreme north of the area, an enormous deposit of oxide ores has been mined, being the result of the alteration of primary sulphide bodies, the transport of metals and their deposition in sinkhole-like cavities. In the south, where the Triassic transgression has occurred, residual and chemical deposition have taken place in karst cavities, with the formation of barite (Barega) (Cortecci et al. 1989) and Ag-rich galena (S. Giovanni). In this last mine a natural void (Santa Barbara cave) has been encountered (Forti & Perna 1984) with traditional speleothems. During the Trias, tabular crystals of barite grew upon these speleothems and, on top of this barite, calcite and aragonite are actually depositing. In the long geological history of Iglesiente the circulation of superficial water, of chemically complex fluids, deriving from the alteration of metal deposits, of hydrothermal fluids and of Sea water have determined important phenomena of dissolution and alteration of rocks, with redeposition of metallic minerals, barite, calcite, SiOl making this mining region a study area of exceptional interest.
Iglesiente-Sulcis mining district, Southwestern Sardinia, Italy. Mineralium Deposita 24( 1): 34-42. De Vivo, B., A. Maiorani, G. Perna & B. Turi 1987. Fluid inclusion and Stable Isotope Studies of Calcite, Quartz and Barite from karstic caves in the Masua mine, Southwestern Sardinia, Italy. Chemic. Erde 46: 259-273. Forti, P. 1996. Thermal Karst Systems. Acta Carsologica 25: 99- 1 17. Forti, P. & G. Perna (eds) 1982. Le cavita naturali dell'Iglesiente. Mem. Ist. It. Speleologia 2( 1): 1-229. Forti, P. & G. Perna 1984. La turistizzazione della Grotta di Santa Barbara nella Miniera di S.Giovanni e la creazione di un muse0 minerario ad essa connesso (Iglesias-Sardegna Sud Occidentale). Atti Conv. Intern. Grotte Turistiche 181187. Borgio VarezziCigna. Forti, P. & G. Perna 1985. L'ipercarsismo con particolare riguardo all'Iglesiente (Sardegna Sud Occidentale). Natura AIpina 36 (2-3): 85-100. Hill, C. & P. Forti 1997 Cave minerals of the World Nat. Spel. SOC.:1-464. Ludwig, K.R., G. Perna, R. Vollmer, B. Turi & K.R. Simmons 1989. Isotopic constraints on the genesis of base-metal ores in Southern and Central Sardinia. Eur. J. Mineral. 1: 657666. Perna, G. 1995. I1 carsismo profondo nel Sulcis-Iglesiente (Sardegna Sud Occidentale) e nel Trentino-Veneto (Alpi Sud Orientali Italiane). "Carsismo messiniano": esempi di carsismo profondo correlato con il livello del Mediterraneo nel Messiniano. Atti M u . Civ. Rovereto, Sez. Archeol., Storia, Sci. Nut. 10: 327-378. Perna, G. 1998. Deep Messinian karst in Mediterranean area. Proc. 1.2'~ Intern. Congress ofspeleology I : 397-399. La Chaux-de-Fonds.
REFERENCES Bechstadt, T. & M. Boni 1996. Sedimentological, stratigraphical and ore deposits field guide of the autochtonous Cambro-Ordovician of South western Sardinia. Mem. Descr. Carta Geol. Ital. 48: 1-390. Bini, M., M. Cremaschi, P. Forti & G. Perna 1988. Paleokarstic fills in Iglesiente (Sardinia, Italy): sedimentary processes and age. Ann. Soc. Gkol. Belg. 1 1 1: 149-161. Civita, M., T. Cocozza, P. Forti, G. Perna & B. Turi 1983. Idrogeologia del Bacino Minerario dell'Iglesiente. Mem. Ist. It. SpeIeoIogia 2(2): 1-137. Cortecci, G., J.-C. Fontes, A. Maiorani, G. Perna, E. Pintus & B. Turi 1989. Oxygen, sulfbr and strontium isotope and fluid inclusion studies of the barite deposits fiom the
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Diagenetic zeolite and clay minerals in Miocene Great Bahama Bank carbonate sediments (ODP Leg 166, Site 1007) A .M Karpoff EOST, CNRS-Centre de Gkochinzie de la Surface, Strasbourg, France
S .M .Bernasconi Geology Institute, ETH-Zentrurn, Zurich, Switzerland
C .Destrigneville & P.Stille EOST, CNRS-Centre de Gkochimie de la Surface, Strasbourg, France
ABSTRACT: To study the fluid-flow processes and effects of sea-level fluctuations in the margin of a carbonate platform, seven sites were drilled during ODP Leg 166 on the western flank of the Great Bahama Bank, a prograding sequence active since late Oligocene. At the distal Site 1007, the early Miocene section is distinguished from the overlying units by a specific authigenic paragenesis made of Mg-smectite and abundant K-clinoptilolite filling foraminifer and pore spaces. The intense diagenetic transformation, recorded by sediment and interstitial fluid geochemical trends, implies by lack of any volcanic input the remobilization of Si-, Al-and Mg-rich fluids and appears promoted by in situ water-rock reactions. Thermodynamic modelling and isotopic data specify the diagenetic processes. The implications on the structure of the GBB platform is discussed.
1 SETTLING AND OBJECTIVES Fluid flow and water-rock interactions within massive sedimentary deposits provide a mechanism of chemical transfers from lithosphere to hydrosphere. The objectives of the Leg 166 of the Ocean Drilling Program (ODP), by drilling 7 sites along a transect on the western flank of the Great Bahama Bank (GBB), were to define the causes and effects of eustatic sealevel fluctuations, and to examine fluid flow processes through a carbonate platform (Eberli et al. 1997). The Bahamas Transect represents a prograding sequence, from late Oligocene to Holocene. Recovered sediments, with carbonate contents varying between 50 and go%, are alternations of wakestones or limestones and clay-rich intervals which are interpreted to reflect enhanced erosion during sealevel lowstands and incipient sea level rise. (Eberli et al. 1997). The present study is focused on the noncarbonate minerals, from the clay-rich levels, in order to investigate, besides the variations of detrital input throughout time, the potential effect on the characteristics of the deposits of the diagenetic alteration, which documents evidence of fluid flow. 2 SITE, SAMPLES AND METHODS Samples are from sequences recovered at the distal Site 1007, located on the toe-of-slope of the western GBB in 647 m of water depth in the Straits of Florida
(Fig. 1). The drilled section consists of 1235.4 m of a nearly continuous sequence Pleistocene to upper Oligocene in age divided into eight lithostratigraphic units. Facies are: foraminifer wackestone, foraminifer nannofossil chalk and limestone, bioclastic packstone to wackstone, mudstone, and thin clay- and organicrich layers. Increasing compaction-dissolution features are observed below 1120 m below sea floor (mbsf), just below the depth location of the seismic sequence boundary P2 (Eberli et al. 1997). Data are obtained on bulk sediment (clay-rich layers) and extracted carbonate-fkee clay fractions (<2pm). The standard X-ray diffraction analyses (XRD) are made on non-oriented powders of sediments and on 4 types of oriented aggrcgatcs of the <2 pm fractions. Samples were observed under scanning electron microscope (SEM) and fine fractions under transmission electron microscopes (TEM,STEM)with an energy dispersive X-ray analysis system (EDS).Major and trace element concentrations in bulk samples are measured by ICP-AES and ICPMS. The isotopic ratios on bulk sediments are determined by conventional mass spectrometric analysis of CO, gas and reported as &values in %O relative to the VPDB and vSMOW standards. Shipboard interstitial water analyses were performed on 5 to 15 cm-long whole core sections (Eberli et al. 1997). Thermodynamic modeling uses the Tardy & Garrel's method (1 974) for mineral constants calculation and the option of the KINDIS computer code (Made et al. 1990).
623
wsw Straits of Florida
Great Bahama Bank Clino
3
ENE
Unda
0-
a,
--E
-
z05 x
sb 1.0
-
1006
Pieistocene ~
-
t1.5
Fig. 1: Geometry of the GBB margin from high-resolution seismic line and position of the drilling sites; (a,...) seismic sequences and (A,...) boundaries (from Eberli et al. 1997). 3 SILICATE MINERALS ASSEMBLAGES The minerals identified (XRD) are: aragonite, calcite, Mg-calcite, scarce dolomite, quartz, plagioclase, pyrite, clinoptilolite,and phyllosilicates. The clay minerals in the <2pm fraction are: chlorite, kaolinite, illite, mixedlayers illite-smectite (IS), smectite. In few samples, palygorskite and celestine occur (TEM, SEM). The mineral assemblages discriminate between the sedimentary sequences of the GGB. 3.1 Plio-Pleistocene detrital clay assemblage In Plio-Pleistocene clayey interbeds, which were more studied at both distal Sites 1006 and 1007, comprise together with carbonates, quartz, plagioclase, a clay assemblage made of chlorite, illite, smectite, illitesmectite mixed-layers (IS), kaolinite and palygorskite. The detrital assemblages, with greater smectite input within late Pliocene units than in Pleistocene oozes, relate either varying source areas or palmclimatic change in weathering intensity for the current- or wind-transported clays (Karpoff et al. 2001). 3.2 Early-late Miocene clay and zeolite paragenesis The non-carbonate paragenesis of the Miocene section relates to a different origin than that of the overlying sequences. The detrital input is reduced, quartz-dominated, and during the late Miocene period includes feldspar, illite, and scarce palygorskite. Accessory minerals are pyrite, celestite, opal-CT or cristobalite. Smectite is the prevalent clay mineral. Abundant zeolite characterizes the early Miocene sequence below the sample at 1094 mbsf. In the mid-Miocene deposits, the homogeneous smectite population forms interlaced elongate flakes with curled edges (TEM),a morphology ascribable to authigenic clay. From STEM data the Mg-clay formula is: [Si,] [All,, Mg,,,] 0,, (OH),. In early-Miocene zeolite-rich section, cogenetic smectite exhibits smoothly crumpled curled-edged flakes, another
autugenic morphology, with an average formula: Mg0.571 [ k . l 9 ] 1' 0 (OH), . [si3.9! A10.9] LA1I,'26 Clinoptilolite as euhedral crystals less than 15 pm long fills up foraminiferal chambers, pore spaces, and in place epigenizes coccoliths. The EDX analyses indicate that zeolite is K-rich (SUAlz4.5) within tests and more Ca-rich (SYAlz4.3) in the matrix. The diagenetic front correlates with increasing compaction-dissolution features and major calcite recrystallization. Such a clay-zeolite paragenesis is unusual in carbonate-rich sequences lacking a volcaniclastic input or far above basaltic basement, such as the Site 1007. 4. CONDITIONS OF FORMATION OF THE DIAGENETIC PARAGENESIS The precipitation of celestite (SrSO,) often described in pelagic sediments (e.g. Swart & Guzokowski 1988; Richter 1996) can remove major quantities of Sr2+ from pore-fluids. Cristobalite and opal-CT are common diagenetic phases in pelagic deposits (e.g. Kastner et al. 1977; Jakobsen et al. 2000). In carbonate deposits the secondary silica phases could be related to either an enhanced biogenic input or the in-situ alteration of silicates, or the intrusion of Si-rich fluids. Authigenic and diagenetic clays are widespread in oceanic realm (e.g. Kastner 1981; Karpoff 1992). Smectite from GGB Miocene samples shows all features of authigenic clay. In the uppermost intervals, smectite could be developed at expense of IS mixedlayers; but in deeper deposits, its Si-rich composition, and cogenesis with zeolite, argue better for a direct precipitation from pore-fluids. The formation of clinoptilolite, which temperature of stability ranges from 0°C to 150-2OO0C,requires a high Si/Al ratio of the reactive phases and a high silica activity. Almost all descriptions of massive zeolite occurrences in oceanic sediments, together with smectites, involve volcanic material as precursor (e.g. 624
Kastner 1981; Aoki & Kohyama 1998). In few case however, zeolite is precipitated into almost pure calcareous sediments fi-om solutions from lowtemperature alteration of the underlying basalt (Bernoulli et al. 1978; McKenzie et al. 1980). De Ros et al. (1 997) attributed the clinoptilolite occurrence in Miocene deposits of the African margin to increased Si activity of the pore waters linked to dissolution of opaline bioclasts. In the GBB Miocene sediments, no volcanic or glass shards were found or described in adjacent sites. A biogenic silica input is evoked, probably having formed the late Oligocene chert-rich layer in bottom hole. Then, the more than likely process for the smectite-clinoptilolite and associated celestite formation seems to be the interaction under alkaline conditions of a Si-enriched fluid with the carbonates and primary silicates. The latter should have provided Si, Al, K elements, and the carbonates the Ca, Mg, Sr, for allowing new Mg-clay, K-zeolite, and SrSO, to precipitate. The pore-water chemistry profiles (Eberli et al. 1997) show depletion in K' and Mg2' and increasing alkalinity at levels of clinoptilolite formation. 4.1
Modelling
In order to test the equilibrium of newly formed phases, a thermodynamic modelling of the saturation index of the given minerals in relation with the chemistry of the pore-fluids and at the given temperature profile (9 to 21"C, from Eberli et al. 1997), was run out. Results specify that: (1) all carbonate phases are oversaturated with respect to present fluid chemistry, but aragonite is unstable below 800 mbsf, (2) quartz and illite stay oversaturated throughout the profile, (3) the saturation index of various zeolites (at 25OC) increases with depth; analcite, a high temperature zeolite, stays undersaturated and phillipsites show lower saturation index than those of clinoptilolites (Fig. 2), (4) the Kclinoptilolite with higher Si/A1 ratio becomes more oversaturated below 969 mbsf, and particularly better still below 1100 mbsf, at level of strongest diagenetic alteration, (5) below 969 mbsf a change in the chemical composition of the smectite becoming Mgrich is needed to keep a clay mineral in equilibrium with the present day fluid system. 4.2
.........................
10
a
-
!
1
:
i ..... :
a . *............. .
a
i
U .
;
a
...............................................
~
m
0
*
I
i__...
i"
.
I
i0
o iI
-10
0
0
0 I
I
I
I
400 600 800 1000 depth (mbsf) 0 K-Clinoptilolite %Na-Clinoptilolite OK-Phillipsite OAnalcite ONa-Phillipsite 0
200
Fig. 2: Saturation index of various ideal zeolites related to pore-fluids chemistry vs. depth Strontium isotope measurements, in progress today on extracted silicates, will be used to characterise the setting and the terms of the sediment-fluid reactions, i.e. contribution of the detrital or biogenic phases, or of a deep or lateral fluid intrusion. 4.3
Geodynamic implication
The depth of massive authigenesis front in early Miocene sequence is well recorded at 1 104 mbsf, at a level of low sedimentation rate and of increased compaction features, and is linked to the position of the seismic sequence boundary P2, dated at 23.2 Ma, Eberli et al. (1997). Its follows two remarks: (1) the diagenesis should be contemporaneous of the deposits or younger, linked to the circulation of fluids and (2) the driven process could be assumed to be a consequence of a geodynamic event of the complex Caribbean region. The diagenetic front could also be recorded as a ''seismic boundary" which significance could be not only chronostratigraphic. This aspect also lies in the domain of researchers studying sealevel seismic record (i.e. Anselmetti & Eberli 1999).
Isotopic signatures
A large scale fluid circulation is not evidenced by the stable isotope composition of sediments (0 and C, on bulk rocks), which preserves primary compositions of the carbonates (Fig. 3). Likewise all documented features and records of carbonates phases diagenesis, such as isotopic composition of the pore-fluids, are related to small scale dissolution-recrystallization processes (Swart 2000; Melim et al. 2001).
0
200
400
600
800
1000
1200
depth (mbsf)
Fig. 3: Isotopic composition of bulk sediment vs. depth
625
5 CONCLUSION In clay-rich interbeds of a distal ODP Site1007 drilled on the GBB, the assemblage, the crystalline and chemical characteristics of the clay minerals and associated silicates report to various origins. Compared to the Plio-Pleistocene detrital assemblage, the Mg-smectite and clinoptilolite paragenesis of the early Miocene sequences express an intense diagenetic transformation still in equilibrium with present day pore-fluids. Without the contribution of volcaniclastic input, the water-rock reaction implies Si, A1 and Mg-rich fluids having circulated in these old sequences. The enrichment in elements o f the reacting fluids have to be search in earlier process involving either alteration of sedimentary materials (detrital and/or biogenic), or reaction with underlying basement, probably at a period of regional tectonic reorganization. Moreover, the correspondence of the depths o f the diagenetic front and o f a seismic sequence boundary P2 allows reassessment of the chronostratigraphic significance of some seismic reflections in the deeper sequences of the platform.
REFERENCES Anselmetti, F. & G. Eberli 1999. Documenting time lines on seismic sections by carrying seismic sequences across platform, slope and basin environments. In G.F. Camoin & W.-Ch. Dullo (eds), Puleoceanology of recfi and carbonates plutfbrins Miocene to Modern. ASF Publ. 32: 15-17. Aoki, S. & N. Kohyania 1998. Cenozoic sedimentation and clay mineralogy in the northern part of the Magellan Trough, Central Pacific Basin. Marine Geol. 148: 21-37. Bernoulli, D., R.E Garrison & F. Melikres 1978. Pliillipsite cementation in a foraminiferal sandstone at Hole 373A and "the case of the violated foram". In K.J. Hsu, L. Montadert et al. Init. Repts. DSDP 42: 478-482. De Ros, L.F., S. Morad. & I.S. Al-Aasni 1997. Diagenesis of siliciclastic and volcaniclastic sediments in the Cretaceous and Miocene sequences of the NW African Margin (DSDP Leg 47A, Site 397). Seclinzent(rr;v Geol. 112:137-156. Eberli, G.P., P.K Swart, M.J. Malone et al. (eds) 1997. P r m . ODP, Init. Repts. 166. Jakobsen, F., H. Lindgreen & N. Springer 2000. Precipitation and flocculation of spherical nanosilica in North Sea chalk. Claj~Miner. 35: 175- 184. Frank, Y.D. 2000. Data Report: geochemistry o f Miocene sediments, Sites 1006 and 1007, Leeward Margin Great Bahama Bank. In P. Swart, G.P. Eberli et al. Proc. ODP, Sci. Results, 166: 137- 143. Karpoff, A.M. 1992. Cenozoic and Mesozoic sediments from Pigafetta Basin, ODP Leg 129 Sites 800 and 80 1 : mineralogical and geocheniical trends of the oldest oceanic crust sedimentary cover. Ill
R.L. Larson, Y . Lancelot et al. Proc. ODP, Sci. Results, 129:3-30. Karpoff, A.M., C. Destrigneville, D. Bartier & P. Dejardin 200 1. Phyllosilicates and zeolite assemblages in the carbonate periplatform of the Great Bahama Bank: origin and relation to diagenetic process (ODP Leg 166, Sites 1006 and 1007). Marine Geol. in press Kastner, M. 198 1. Authigenic silicates in deep-sea sediments: formation and diagenesis. In C. Emiliani (ed.), The oceanic lithosphere. The Sea 7:915-980. New York: Wiley. Kastner, M., J.B. Keene & J.M. Gieskes 1977. Diagenesis of siliceous oozes - I. Chemical control on the rate of opal-A to opal-CT transformation, an experimental study. Geochim. Cosmochim. Actu 41:1041-1059. Made, B., A. Clement & B. Fritz 1990. Modelisation cinktique et thermodynamique de l'alteration: le modkle geochimique KINDIS. C. R. Acad. Sci. Paris 308:647-654. McKenzie, J.A., D. Bernoulli & S.D. Schlanger 1980. Shallow-water carbonate sediments from Emperor Seaniounts: their diagenesis and paleoceanographic significance. In E.D. Jackson, I. Koizumi et al. Init. Repts. DSDP, 55:415-455. Melim, L.A., H. Westphal, P. Swart & G.P. Eberli 200 1 . Questioning carbonate diagenetic para-digms: evidence from the Neogene of the Bahamas. Murine Geol. in press Richter, F.M., 1996. Models for coupled Sr-sulfate budget in deep-sea carbonates. Earth Planet. Sci. Lett. 141:199-211. Swart, P. 2000. The oxygen isotopic composition o f interstitial waters: evidence for fluid flow and recrystallization in the margin of the Great Bahania Bank. In P. Swart, G.P. Eberli et al. Proc. ODP, Sci. Results, 166:91-98. Swart, P. & M. Guzokowski 1988. Interstitial-water chemistry and diagenesis of periplatform sediments from the Bahamas, ODP Leg 101. In J.A. Austin, W. Schlager et al. Proc. ODP, Sci. Results, 101 :363380. Tardy, Y. & R.M. Carrels 1974. A method for estimating the Gibbs energies of formation of layer silicates. Geochim. Cosmochinz. Acta 38: 1 10 1- 1 16.
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Water-Rock Interaction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Numerical study of the coupling effect between fluid diffusion and medium deformation for subsidence calculation over deep reservoirs G.Lecca & R.Deidda Centro di Ricerca, Sviluppo e Studi Superiori in Sardegna (CRS4),C.P. 94, I-09010 Uta, Italy
G .Gambolati Dipartirnento di Metodi e Modelli Maternatici per le Scienze Applicate, Universita di Padova, Via Belzoni 7 , I20097Padova, Italy
ABSTRACT: The mechanical behavior of saturated porous media is influenced by the interaction of the solid skeleton with the water present in the pore structure. This study investigates the importance of the coupling effect between water flow and porous matrix deformation in subsidence calculations over compacting deep reservoirs, comparing two modeling approaches. The first one is an approximate uncoupled model consisting of two separate field calculations: time evolution of pore pressures and evaluation of the corresponding matrix displacements generated by the pressure gradients. The second approach is a fully coupled Biot poroelastic model that accounts for pore pressure and displacements as a unique set of unknowns to be simultaneously advanced in time. The study is carried out for a disc-shaped reservoir located in a cylindrical sedimentary basin, varying the geometrical and hydromechanical characteristics of the problem. Knowing the importance of the coupling effect can help discern geometries, ranges of physical parameters and time length scales for which the less computationally expensive uncoupled model may provide adequate results. 1 INTRODUCTION
Land subsidence caused by fluid extraction from productive reservoirs has been investigated since the beginning of the 19-th century. But it was only later that Biot (1941) developed the mathematical framework of the physical processes governing the poroelastic problem. Several authors have contributed to this subject since subsidence prediction over depleted gas, oil and water fields is of great environmental and economical relevance to many countries. A major debate involving the scientific community in the last decades concerns whether to treat the prediction of land subsidence induced by reservoir exploitation with coupled (Hsieh 1996) or uncoupled (Geertsma 1973) models. The former approach supplies the time evolution of both fluid pressures and the soil displacements by solving simultaneously the equations governing the mechanical equilibrium and the groundwater flow. The uncoupled approach, instead, first calculates the time evolution of the pressure field using the groundwater flow equation and then solves separately the equilibrium equation for the soil displacements at the times of interest. From a physical and mathematical point of view the coupled model represents the more rigorous approach to land subsidence prediction, but it is very computationally expensive, and, in addition, its numerical solutions are often affected by numerical instability and convergence problems (Hsieh 1996, Ferronato et al. 2001). On the other hand, the approximate un-
coupled model, although theoretically less accurate, requires less computational resources allowing a finer description of the medium properties. In addition, since the uncoupled model does not suffer all the numerical problems raising with the coupled model, non-specialized users can easily employ it, as long as coupling is weak. In this respect an investigation concerning the importance of coupling in the sedimentary basin of the PO river plain, Italy, has recently been perfomed by Gambolati et al. (2000). For both models a standard Galerkin finite element formulation for an axisymmetric geological setting is developed with an unconditionally stable implicit time integration scheme. Although the coupled and uncoupled models are based on a similar numerical formulation, the computational resources needed for the uncoupled simulator are significantly less than for the coupled model, both in memory and CPU requirements. For these reasons it is of great interest to investigate under which conditions and configurations the uncoupled model supplies results with acceptable errors in practical applications. The main objective of this paper is thus to evaluate the differences between the two formulations and the worsening in the accuracy of the results of the uncoupled model with respect to the coupled one. A parametric comparison, aimed to a quantitative evaluation of the land subsidence predicted by the two models, is carried out on a simplified sedimentary basin configuration for different hydromechanical properties and geometrical layouts. 627
are the vectors of the unknown nodal displacements and pressures, and their time derivatives, respec-
2 MATHEMATICAL MODEL The governing equations for the fluid flow and the poroelastic stress in a three dimensional isotrogji; porous medium are reviewed here in terms of the incremental variables: pore pressure p and components 24, v, U’ of the displacement vector & . Following Gainbolati et ctl. (2000) they read:
is the vector of nodal loads and tively; and {f ,f flow sources. The FE formulation of the uncoupled flow eq. (3) reads:
Hp+SP = f ”
1 dp dE -V(kVp) = [np + cKr(1 - M)]+at at Y
(A
while eq. (2) remains in the uncoupled formulation as: Ku = Qp + f In equations (4) and ( 5 ) time derivatives are computed using a weighted finite difference scheme with a, usually taken as 0.5, the weighting parameter. As an example, in eq. (5) the pressure field pk” at time t k i 1 = t k + Atk is obtained evolving the field pkat time tk according to: ‘ I .
aP + G)-d&+ GV% = x -
dx dX ( R + G )d& - + G v 2 v = X - aP 8Y ay
where t is the time, y and p are the specific weight and the volumetric compressibility of water, respectively, k is the hydraulic conductivity tensor, M is the porosity, cgr is the rock grain compressibility and the medium volume strain is defined as&= V -d. Denoting by E and v the Young and Poisson moduli of the bulk porous medium, and by cb the medium compressibility, the remaining symbols are defined as x = (1 - cgr/cb) the Biot coefficient, G = E/2( 1 + v) the shear modulus, and h = vE/[( I - 2v)/( 1 + v)] the Lame constant. While the coupled model solves simultaneously eqs. (1) and (2), in the uncoupled approach they are solved separately. First the time evolution of the pressure field p is obtained by the only eq. (l), eliminating the dependence on E with the equation E = aw,being a = l / ( h+ 2G) the vertical medium compressibility, and then p is supplied to the right-hand-side of eq. (2) for the calculation of the corresponding displacement field. Eq. (1) can thus be rewritten for the uncoupled model as: -V(kVp) 1
Y where S storage.
= S-
Algebraic systems of type A x = b, like eq. (6), arising from the numerical discretization of eqs. (1)-(3) are then solved using iterative solvers of the conjugate gradient family, with incomplete factorization of the coefficient matrix A (Saad 1996). Both numerical models have been tested for accuracy and robustness by comparison with analytical and numerical solutions (Lecca 1999). 3 SIMULATION RESULTS
(3)
dt =
[np + ch,.(l-n)+ a ~is ]the specific elastic
The finite element (FE) formulation (Gambolati et. crl. 2000) of the coupled eqs. (1) and (2) leads to:
(4) where K, H,(2 and SI are the elastic stiffiiess, the flow stiffness, the poroelastic coupling and the flow capacity matrices, respectively;
{U,
p]“ and
(5)
{U,
p>’
To restrict our attention on the sensitivity of the coupled and uncoupled models to the different geometrical and hydromechanical conditions considered in this study, a simplified geological setting is assumed. It consists of a disc-shaped reservoir of radius r, thickness s, and depth c interbedded between two clay layers of thickness t and hosted in a cylindrical sedimentary basin of height and radius e = 10000 m. Figure 1 shows a schematic view of the case study and defines the geometrical parameters involved. In all the siinulations the following boundary conditions were imposed: null variation of displacements and pore pressures on the domain outer boundary (AB) and basis (BC), regarded as unperturbed zones; null variation, for geometrical compatibility, of radial displacements on the cylinder axis of symmetry (CD) and null variation of pore pressure at the domain surface (AD). The reservoir is pumped with a volume-distributed rate amounting to Q = 3x 10-3 m/s, which is linearly increased from zero to Q in one day and then kept constant until steady state. Initial conditions both for incremental displacement and pore pressure were taken as zero. Two values, constant with depth, of hydraulic conductivities k, (1 0-7- 1Oe5m/s) for the aquifers (pumped or not) and
628
Figure 1. Scheme of the computational domain with the definition of the geometrical parameters (right) and enlargement of the reservoir zone interbedded between two clay layers (left).
k? (1 O-'* - 10-'I mis) for the aquitards, have been assumed. The vertical medium compressibility a and the corresponding porosity n (using the definition of oedometric compressibility) have been derived as a function of depth for two data sets (Gambolati & Teatini 1998): in situ measurements (class I) and oedoinetric laboratory tests (class 11). Figure 2 shows vertical medium compressibility a vs. depth for the two sets of measurements. The other constant parameters of the simulations are summarized in Table 1. In the sensitivity analysis a reservoir of depth c = 1500 in, radius r = 1000 m, and thickness s = 10 in, interbedded between two aquitard layers of thickness t = 1 ni, with hydraulic conductivities kl = 10-6and k-7 = IO-" m/s, was regarded as the reference case. Figure 3 compares, for this refererwe case, the time evolution of coupled and uncoupled vertical displacements (subsidence) v g at the point D, located at the intersection of the domain axis with the surface, both for class I and I1 characteristics. The uncoupled solution overestimates the coupled one, although this effect is less pronounced for class I than for class 11. The difference between coupled and uncoupled displacements is always present but tends to reduce itself until to vanish at steady state, as the poroelastic theory dictates. The duration of the transient regime, in which the coupling effect is appreciable, is longer for class I1 then for ;;:ass I due to the larger laboratory compressibilities with respect to the in situ ones Table 1. Constant parameters of the simulations.
Water specific weight Grain compressibility Water compressibility Poisson coefficient
27860 N/m 9810 N/m3 1.631x 10.' m'/N 4 . 5 ~ 1 0 ' m2/N '~ 0.3 9810 N/m3
'
Figure 2. Vertical medium compressibility CI [m2/N] vs. depth [m] for in situ (class I) and laboratory (class 11) measurements (after Ganibolati & Teatini, 1998). On the average class I1 CI is about 5 times larger than class I one.
(Fig. 2). To investigate the importance of the hydrological parameters like aquifer kl and aquitard k2 conductivities on the coupling effect, a set of simulations was run with all the parameters of the class I reference case, except for the reservoir thickness s which was set 50 in. Isotropic and homogeneous conductivity k, ranging from 10-7to 10'6 m/s was assumed, while the ratio k = kl / kz was kept to a constant value of 105.A second set of simulations was also run using the same layout in order to evaluate the importance of the radial anisotropy ( k l , > kl,) to mimic the presence of horizontal preferential flow paths, which are typical of stratified deposits, in the unpumped strata above the reservoir (overburden). Figure 4 compares the results of all these simulations, showing the time evolution, at the same point D, of the ratio of uncoupled vo in the folto coupled subsidence: vo lowing referred to as the coupling factor. Reducing kl by an order of magnitude from 10-7(empty square curve) to 10-6niis (filled circle curve) reduces the duration of the transient regime, which lasts from about 10 years to 1 year, by the same factor. The coupling effect is visible in both curves that appear geometrically similar but it is indeed of practical interest only for the less permeable configuration. Moreover, introducing a radial anisotropy of one order of magnitude in the overburden ( k l , =10-5 and k,, =10-6 m/s) contributes substantially to a further reduction of the coupling effect: as an example, at time 106s (about 2 weeks) from the start of pumping in the isotropic case (filled circle curve) uncoupled displacements exceed coupled ones by 1.9, while in the anisotropic case (empty rhomb curve) this coupling factor is reduced to 1.5. Results from other runs show that a lower aquitard hydraulic conductivity kz, while keeping k l constant, emphasizes the 629
Figure 3. Land subsidence vfj[m] vs. time [s] for class I (circle symbol) characteristics and I1 (square symbol). Comparison between coupled (filled symbol) and uncoupled (empty symbol) models.
Figure 4. Ratio of uncoupled to coupled land subsidence vfjW l d / Vf;Pld [/I vs. time [s] for two aquifer isotropic hydraulic conductivities k,: 10'7 (square symbol) and 10-' (circle symbol) m l s . The effect of a radial anisotropy k l r =10 klr =10-' m l s of the overburden is also shown (rhomb symbol).
coupling effect and increases the duration of the transient regime in which this effect is remarkable. To investigate also the importance of different geometrical parameters such as c (500, 1000, 1500, 3000 m), r (1000,2500,5000 m), and s (5,10,50 m) in the development of the coupling effect, a series of parametric simulations were performed during the very early stages (1 month) of the pumping, in order to maximize the differences between the two compared models. Each of these groups of simulations operates only on one geometrical parameter at time, while maintaining the general layout of the class I reference case. Only the depth of the reservoir c was found to appreciably affect the coupling effect. As an example, at 1 month from the start of pumping at the point D for c = 3000 m the coupling factor was 2.2 against a value of 1.6 for the depth of 500 m. This suggests that an important role is played in the compared models by a thicker consolidating overburden. All the runs were made on computational grids of about 18000 nodes, 54000 unknowns and 35000 annular finite elements with triangular cross section. The computational cost of each coupled time iteration was about a hundred times greater than the uncoupled (flow and deformation) counterparts. Also the memory requirements were very different: 350 Mbytes of RAM for the coupled model against only 30 Mbytes of both the uncoupled simulators. In conclusion, the study indicates that, for the simplified configuration adopted, the coupling effect between water flow and porous matrix deformation, represented here by the differences in the displacements computed by coupled and uncoupled models, may be significant during the early stages of the pumping and is practically important only for highly compressible sedimentary basins characterized also
by very low permeability deposits with pumped reservoirs located at large depths.
'
Acknowledgements. The authors would like to thank Claudio Zoccatelli from ENI AGIP for helpful collaboration.
REFERENCES Biot, M. A. 1941. General theory of three-dimensional consolidation. J. Appl. Phys., 12(2), 155-164. Ferronato M., G. Gambolati & P. Teatini. 2001. 111conditioning of finite element poro-elasticity equations. Int. J. Solids d Stritctures, (in press). Gambolati, G. & Teatini, P. 1998. Natural land subsidence due to sediment compaction of the Upper Adriatic sea basin. Ingegneria e Geologia degli Acqiiijeri (IGEA) XXXV, 2-3, 29-40, Torino, Italia. Gambolati, G. et al. 2000. The importance of poroelastic coupling in dynamically active aquifers of the PO river basin, Italy. Water Resour. Res., 36(9), 2443-2459. Geertsma, J. 1973 Land subsidence above compacting oil and gas reservoir. J. Pet Technol., 25, 734-744. Hsieh, P.A. 1996. Deformation-induced changes in hydraulic head during ground-water withdrawal. Ground Water, 34, 6 , 1082-1089. Lecca, G. 1999. Validazione e documentazione dei modelli accoppiato e disaccoppiato in assialsimmetria. CRS4-Tech. Rep. 99/06, Cagliari, Italia Saad, Y. 1996. Iterative Methods for Sparse Linear Systems, PWS Publishing Co.
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Water-Rock Interaction 2001, Cidu fed.), 02001 Swets & Zeiflinger, Lisse, ISBN 90 2651 824 2
Squeegee flow in Devonian carbonate aquifers in Alberta, Canada H.G.Mache1, B .E.Buschkuehle & K.Michae1 University of Alberta, Edmonton, Canada
ABSTRACT: The Alberta Basin in western Canada contains four Devonian carbonate aquifers that are interbedded with mark and evaporites The major objective of this study is to characterize the extent of squeegeetype tectonic expulsion of fluids into these aquifers close to the Rocky Mountains. A past influence of squeegee-type flow is evident in late diagenetic calcite cements. Their 87Sr/s6Sr-isotopevalues are unusually high, up to 0.7323, adjacent to the deformation front, decreasing into the foreland basin to 0.7132. The present formation fluid flow in the deep basin is generally upward and away from the deformation front. A lower-salinity brine (100-175 g/l) occurs in the south and along the deformation front and downdip from a high-salinity brine (200-300 g/l). Some of these characteristics could be relics of squeegee-type flow.
1 INTRODUCTION The Alberta Basin in western Canada contains four regionally extensive Devonian aquifers andor hydrocarbon reservoir levels, traditionally called D 1D4 in descending stratigraphic order. The D2+D3 aquifers, as defined by seismic data (Fig. l), are the main subject of this study. Due to the asymmetry of the basin, these strata subcrop or crop out near the northeastern limit of the basin and dip to depths of about 6 km close to the limit of the disturbed belt, which originated during the Laramide orogeny that formed the Rocky Mountains in the Late Cretaceous to Early Tertiary. Within the study area (stippled outline in Fig. l), the Devonian lies at depths of 2-5 k m with an average slope of 15m/km, and consists mainly of marine carbonates and shales, approximately lkm in thickness. Across much of the basin the four Devonian aquifers, which contain oil, sweet and sour gas reservoirs, are interbedded with marly and evaporitic aquitards, and are confined by tight evaporites at the base and by Carboniferous shales at the top. Central to this study is a prominent feature in the deep part of the basin, the Southesk-Cairn complex (SCC). This complex is a major Upper Devonian (D2+D3), northeast to southwest-trending platformreef sequence that extends 125 to 150 km into the foreland basin from the limit of the disturbed belt. Palinspastic restoration of exposures in the Rocky Mountains extends the complex about 300 km farther to the southwest (Fig. 1). We investigated the
SCC mainly in the subsurface but also in some outcrops. The major objective of this study is to investigate
Figure 1. Simplified distribution map of Devonian platform and reef carbonates of the Late Devonian Woodbend Group (D3) in the Alberta Basin. Light grey pattern: undifferentiated Cooking Lake andor Leduc Formations and their outcrop equivalents; medium grey pattern: reef complexes; darkest grey: basinal areas. Fat double lines denote fault systems; black arrows mark inferred paleofluid flow, as explained in the text. L.D.B. = limit of the disturbed belt. SIM = Simonette. Cross section C-D: see Figure 5.
631
whether tectonic expulsion of formation fluids sensu the so-called "squeegee model" (Oliver 1986) can be identified and/or characterized in the deep part of the Alberta Basin. To this end, we are investigating chemical as well as physical components of the present and (inferred) past fluid flow.
4 m
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2 STRATIGRAPHY AND FACIES Stratigraphic investigations and correlations, using well logs and drill core, have shown that the SCC consists of Middle- to Late Devonian reefs and carbonate platforms that are interbedded with mark and subordinately evaporites. Our as well as some previous correlations (Wendte et al. 1998; Skilliter 1999) indicate that the aquitards, although generally regionally continuous, are thin or missing in the deepest part of the basin, such that the four aquifers form a thick, contiguous I'mega-aquifer" near the limit of the disturbed belt. Furthermore, the aquitards are locally discontinuous and/or relatively permeable where thinner than about 10m, which leads to crossformational flow of water and/or hydrocarbons in some locations. 3 DIAGENESIS Most rocks of the SCC are medium to coarsecrystalline, grey dolostones with poorly preserved primary textures. These dolostones are similar to those that are common elsewhere in the basin, and that probably formed from some type of pervasive convection of chemically modified seawater at depths of about 500 to 1,500m (Mountjoy et al. 1999). Most oil and gas is contained in molds and vugs that formed during or after pervasive dolomitization. Several other diagenetic phases, including early marine calcite cements and traces of bacteriogenic sulfide, are common but not abundant. Most important in the present context are coarsecrystalline white calcite cements that occur in overall volumes of about 1 - 2 %, and that, like the hydrocarbons, occur in dissolution vugs. A plot of 613C vs. 6"O (Fig. 2) shows two trends in the data for late calcites from the subsurface at Obed, one of the sour gas fields in the SCC (Fig. 3). The first is a low-angle positive trend showing a wide range in 6"O over a small range in 613C. This pattern is typical for precipitation either from a fluid with a variable meteoric water component or over a relatively wide temperature range from water of a constant, near-marine composition. The latter probably is the more appropriate interpretation, considering that there are no other indications of the incur-
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6''o O/,,PDB Figure 2. Stable isotope data from late calcite cements of the subsurface Leduc (D-3) and Nisku (D-2) Formations in the Obed area of the SCC (filled circles), and data from late calcites in outcrop from stratigraphically equivalent sections in the Rocky Mountains (hollow and filled triangles). Early calcites formed from seawater (SWC) from various locations in the basin are plotted for reference.
sion of meteoric water in these rocks. The second trend is nearly vertical and shows a wide range in 613C (ca. 0 to -27 %O PDB) over a small range of 6l80 (ca. -7 to -12 %O PDB). The range of 613C indicates incorporation of variable amounts of isotopically light carbon derived from the oxidation of organic carbon. Considering further that these particular samples were retrieved from a sour gas field that probably has never been in contact with meteoric water, thermochemical sulfate reduction (TSR) is the only realistic interpretation for the origin of this 613C-trend. Data from two outcrop locations are similar (Fig. 2). Most of the 87Sr/86Sr-ratiosof late calcite cements from Obed and from other parts of the SCC in the subsurface are well above the Middle- to Upper Devonian seawater value of about 0.7085 and also above MASIRBAS=O.7120. The latter value is the Maximum Sr Isotope Ratio of Basinal Shales, as determined by Sr-leaching experiments (Machel and Cave11 1999) that the formation waters can obtain by circulation through these shales at diagenetic temperatures. Values higher than MSIRBAS indicate a metamo hic origin of at least some Sr. Furthermore, the 87Sr/F ?3r-ratios of these calcites are not randomly distributed. Rather, they are highest near the limit of the disturbed belt, 0.7323, and generally decrease eastward to about 0.7132 (Fig. 3). These findings suggests an extra-basinal source of Sr from the fold and thrust belt, as well as a generally easterly flow from the Rocky Mountains into the
632
Figure 3. Distribution of Sr-isotope values of late calcite cements in the Southesk Cairn Complex northeast of Jasper. Compare outline of the platform with that in Fig. 1. Arrows denote inferred flow of 87Sr-enrichedfluids.
foreland basin, while these particular calcitesformed. Furthermore, such fluid flow probably was not restricted to the SCC. The highest values measured to date are from the Simonette reef just north of The SCC (Fig. l), where late calcites range up to 0.7370 and late dolomites range up to 0.7369. The paragenetic sequences established for various parts of the SCC place the white calcite cements as the youngest events of lengthy diagenetic histories. Fluid inclusion data for some of these samples indicate crystallization temperatures ranging fiom 130 to 16OoC, the same range at which TSR is inferred to occur. However, the fluid inclusion data do not show a regional easterly trend comparable to that of the Srisotope data.
4 PRESENT HYDROGEOLOGY Formation water chemistry and pressure data from standard drillstem tests were used to interpret the present flow conditions within the four Devonian aquifers. Both data sets were culled for erroneous analyses and tests, including production influences. Formation water salinity and geothermal gradients were used to calculate water density. Maps of salinity, density and hydraulic head distributions were used in the flow analysis. The salinity and bicarbonate distributions suggest the existence of two different brines in the investi-
Figure 4. Present salinity distribution and flow pattern in the D3 aquifer, as inferred fiom pressure data and salinities.
gated Devonian aquifers. A lower-salinity brine (100-175 g/l) fills the pore spaces in the south and along the deformation front, downdip from a highsalinity brine (200-300 g/l). The low-salinity brine has a high bicarbonate content (750- 4000 mg/l). The transition zone between the two brines is characterized by relatively sharp salinity gradients. Generally, the heavier brine is located progressively farther updip. The regional-scale flow pattern is locally modified by permeability heterogeneities. The lighter brine attempts to flow updip, along large-scale ((fingers)) that may be due to channeled flow along highpermeability paths (Fig. 4). In an area of convergence of the two brines, the flow direction changes from mainly along dip to along strike. This suggests that the lighter brine is attempting to bypass laterally the heavier brine that forms a slug of almost stagnant water. Impelling forces acting on formation waters in each aquifer were calculated based on hydraulic-head and density distributions, to account for both potential and buoyancy components. The force vectors (Fig. 4) indicate regional-scale converging flow directions for the two brines, with the lighter brine moving updip and laterally, and the heavier brine directed downdip. The relatively small magnitude of the impelling force suggests that the two brines are almost stagnant, however, except for the transition zone between the two. Both types of brines are interpreted to be of mainly 'connate' origin, their high salinity being the result of seawater evaporation during the Devonian
633
and some mixing with Laramide to Pliocene meteoric waters. The marked difference in salinity between the two brines could be due to either one or a combination of the following processes. If the lighter brine is considered representative for the original connate water, then the heavier brine may be the result of halite dissolution of Elk Point evaporites strata (Prairie Fm.) that are present in the northwest part of the study area. In this case, the originally lighter brine downdip is most probably driven updip by past tectonic compression during the Laramide orogeny. On the other hand, if the heavier brine is representative of the original connate water, the lighter brine may be the result of dilution by fresher water of either meteoric origin or with a metamorphic component. At present, the Devonian aquifers cannot be recharged from above because Of the hydraulic barriers comprised of thick, competent aquitards, as well as by thick zones of underpressured (sinks) and overpressured strata that are present in the Overlying Cretaceous succession* Hence, the Devonian strata be recharged Only from the Rocky Mountain deformation front to the southwest and from the south in a through and flow 'Ystem Originating in Montana where these aquifers crop out.
Figure 5. Cross section C-D (see Fig. 1 for location) with vertical exaggeration of about 6: 1, schematically showing presentday formation fluid flow (hollow arrows) and inferred Laramide tectonically induced fluid flow (solid, black arrows). Circles with X indicate present-day flow subperpendicular to the plane of view. Present flow in the upper, Cenozoic and Mesozoic parts of the basin is essentially meteoric and topography-driven.
that squeegee-type flow was laterally rather limited, into the foreextending perhaps only 100 to 200 land basins, where the radiogenic Sr signal lldisappears". Our data, therefore, support the results of modeling studies that indicate generally low fluxes for squeegee flow.
5 CONCLUSIONS
ACKNOWLEDGMENTS
Our data indicate that the Source of the elevated 8 7 ~ ~ / 8 in6the ~ deepest ~ - ~ part ~ ~of~the~ Southesk~ Cairn complex is most probably fluids that interacted with and/or were derived from metasedimentary Precambrian rocks. One likely possibility is the Proterozoic Miette Group, which is present in outcrop near Jasper (Fig. 1) and extends at depth along the entire eastern front of the Rocky Mountains. The overall small amounts of the white calcite cements (1- 2 vol%) reflect a relatively small volume of fluid from an external source flowing through a limited rock system during or immediately after TSR. The spatial distribution of the 87Sr/86Sr-ratiossuggests a general west-to-east flow pattern through the Southesk-Cairn complex, with 87Sr/86Sr-ratiosin the fluids decreasing eastward because of dilution and increasing water-rock interaction with the host carbonates. These findings are strong evidence for squeegeetype flow in the Southesk-Cairn Complex, whereby fluids were expelled laterally into the adjoining carbonate aquifer(s) via thrust sheets and/or from Precambrian basement metasediments via faults that transect the underlying Cambrian clastics (Figs. 1, 5). The present flow system and salinity distribution in the deep part of the basin may reflect, at least in part, past squeegee flow. Our data further suggest
Funding for this study was provided by h o c o , Chevron, Crestar, Petro Canada, Alberta Geological Survey, and NSERC. We thank Stefm Bachu for advice on hydrogeological questions.
REFERENCES Machel, H.G. & Cavell, P.A. 1999: Low-flux, tectonically induced squeegee fluid flow ("hot flash") into the Rocky Mountain Foreland Basin. Bulletin Canadian Petroleum Geology 47: 510-533. Mountjoy, E.W. , Machel, H.G., Green, D., Duggan, J. & Wilhams-Jones, A.E. 1999. Devonian matrix dolomites and deep burial carbonate cements: A comparison between the Rimbey-Meadowbrook reef trend and the deep basin of west-central Alberta. Bulletin of Canadian Petroleum GeolOQ 42 (4): 487-509. Oliver, J. 1986. Fluids expelled tectonically from orogenic belts: their role in hydrocarbon migration and other geologic phenomena. Geology 14: 99-102. Skilliter, C.C. 1999. Stratigraphic and geochemical investigation of Middle to Upper Devonian ccaquitardw in westcentral Alberta, Canada. Unpublished M.Sc. Thesis, University of Alberta. Wendte, J., Dravies, J.J., Stasiuk, L.D., Qing H., Moore, S.L.O. & Ward G. 1998. High-temperature saline (thermoflux) dolomitization of Devonian Swan Hills platform and bank carbonates, Wild River area, west-central Alberta. Bulletin Canadian Petroleum Geology 46: 2 10-265.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Flat lowland paleogeography of sediment-collecting basins: Evidence from formation waters E.Mazor Eizvilannzental Sciences and Energy Research, Weizrnaniz Institute of Science, Rehovot 76100, Israel
ABSTRACT: Many formation waters reveal meteoric-like 6D values, but the dissolved salts resemble evaporitic brines, disclosing an origin as groundwaters with spray and dust from evaporitic brines. These compositions differ from the local recent groundwaters, indicating a fossil nature and confinement. The common association of formation waters with shale, sandstone, gypsum and halite indicates that the active sediment accumulation of subsidence basins was in large-scale low-land platforms, frequented by alternations of sea regressions and transgressions. The regressing sea left behind large-scale evaporitic environments of sabkhas and lagoons, and the exposed marginal marine sediments acted for a while as unconfined groundwater systems. Each paleo-transgression produced new sediments that covered the former rock systems and confined them along with their connate saline groundwaters. Thus, large-scale evaporitic facies of flat lowlands seem to have been an intrinsic part of subsidence basins, as disclosed by the evaporitic affinity of fossil groundwaters within sedimentary basins. 1 INTRODUCTION: RECONSTRUCTING THE ENVIRONMENT AT WHICH SEDIMENTS FILLING DEEP BASINS WERE FORMED We are familiar with the phenomenon of sedimentary basins, filled with thousands of meters of sedimentary rocks. Does the term “basin” indicate that sedimentation took place in large and deep bowl-shaped structures? Or, were the sediments laid down in a different environment and they subsided after their formation? If so, what was the original environment of sediment formation? The question is of great interest as sedimentary basins are found all over the globe, and the age of the sediments indicates thls type of structure was repeatedly formed from the Precambrian on. Thus, we deal with a widespread geological feature. The variety of rocks found within sedimentary basins includes shale, limestone, sandstone, gypsum and halite that indicate precipitation at alternations of a marginal marine environment and low-land platforms, frequented by alternations of sea regressions and transgressions. Gypsum and halite can not be formed at the open ocean, and they disclose formation at sabkhas and lagoons, i.e. formation on flat lowland that was frequented by sea transgressions and regressions. The large scale of many sedimentary basins indicates the mentioned flat lowlands were often of a large scale. Gypsum and halite are easily washed away from their location of initial precipitation, and only in rare cases were evaporitic sediments rapidly covered by a protective layer of other sediments, m d y clay.
Halite is highly soluble, and the fact that halite deposits of even Palaeozoic and Mesozoic ages have been preserved, indicates that the deposits were efficiently engulfed by impermeable rocks. Such an efficient engulfinent must have been a rare case, indicating that formation of evaporites was much more frequent than reflected by the preserved gypsum and halite deposits. Thus, fat lowlands with evaporitic setups were much more common. The lithological record is highly biassed, in the sense that it provides mainly evidence of the marine rocks formed during the transgressive phases. Subaerial processes are mainly connected to erosion and material removal and, hence, sea regressions can easily be overlooked in the rock records. There is a need for additional geological archives that may provide information both of the frequency of transgressions and regressions that occurred at a study site, and the type of environment that prevailed when the land was exposed. The chemical and isotopic composition of formation waters provides this information with accuracy and details, a topic addressed in the present communication. 2 FOSSIL GROUNDWATERS TRAPPED IN SUBSIDED ROCK-COMPARTMENTS Formation waters encountered in adjacent wells, or at different depths within the same well, often reveal a pronounced variability in salinity, ionic ratios and 635
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subsided with their interstitial waters, which thus reflect the large variety of evaporitic environments that prevailed at the different locations at different times. These are connate waters. The quantity of specific formation waters is generally limited, and we commonly know about them from DST samples or co-production with petroleum.
isotopic composition (Mazor 1995, 1998, Mazor et al. 1995). Figure 1, based on data by Egeberg & Aagaard (1989), reveals an example of depth profiles from Triassic and Jurassic strata in the Norwegian Shelf. A large variety of water chemistries is observed, leading to the conclusion that the formation waters were formed under different paleo-environments and, hence, are fossil. The observation of spatial variability leads to the conclusion that the formation waters are stored in distinct rock-compartments that are hydraulically sealed. This point is further elucidated by the example of enormous variety of formation waters encountered within the large West Canada basin, as depicted in Figure 2, which is based on data from Hitchon et al. (1971). None of these waters is formed there at present, and the only way these groundwaters could be introduced into the distinct rock-compartments was by recharge into the rocks while they were exposed on the surface. The rocks 100000
3 METEORIC GROUNDWATER RECHARGED INTO MARINE ROCKS THAT WERE EXPOSED DURING SEA REGRESSIONS A literature survey reveals that the isotopic composition of formation waters is light: 6D values are in the range of -120%0to -20%0.This is a clear proof that the formation waters are meteoric groundwaters, and hence, they were recharged into the present host rocks when the latter were exposed on land, during sea regressions.
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Figure 2. Composition fingerprint diagrams of petroleum-associated formation waters, western Canada sedimentary basin (data from Hitchon et al. 1971). The only way waters of such a salinity and composition diversity could be introduced into the host rocks was by recharge in evaporitic flat lands and confinement by sediments during a following sea transgression.
636
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by salts from brine-spray and evaporitic dust. Hence, the landscape at which the formation waters were injected into their host rocks was a widespread flat lowland frequented by sea transgressions and regressions. This type of environment well explains the large variety of chemical composition of formation waters - it reflects the variability observed at existing evaporiticterrain.
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6’8O%o Figure 3. Isotopic composition of formation waters from different oil fields in the eastern Michigan Basin (data from Weaver, 1995). The data reveal light isotopic compositions and they plot along the Global Meteoric Water Line (GMWL), indicating these are fossil meteoric groundwaters.
The 6l80 values are in certain formation waters light as well, and plot on the global meteoric water line (GMWL) in 6D - 6l80 plots, e.g. data from the Michigan Basin, as seen in Figure 3, which is based on data from Weaver et al. (1995). In many other cases formation waters have relatively heavy 6l80 values, plotting to the right of the GMWL This reflects isotope exchange with rocks at elevated temperatures, reached upon subsidence. The rocks hosting formation waters are mainly marginal marine. The meteoric formation waters were recharged into them during phases of sea regression, when these rocks acted as continental groundwater through-flow systems. Eventually, the rocks subsided and got confined by new marine sediments, which at their turn got exposed and filled with groundwater, and so on. Each rock unit subsided along with its interstitial groundwater.
Formation waters are by and large fossil groundwaters (differing from local recent groundwater), with a high spatial variability (differing between adjacent wells), constituting ancient meteoric groundwaters (as reflected from their light isotopic composition), that were tagged by evaporitic salts (as disclosed by medium to high salinities, ClBr weight ratios of 20 to 200, and presence of CaC12). These are saline groundwaters formed on land, recharged into exposed marine rocks, and trapped in them when they subsided. Their common occurrence in large basins discloses that flat lowlands, frequented by alternations of sea transgressions and regressions, were widespread, forming the top of evolving sedimentary basins. The spatial distribution of formation waters, and their chemical and isotopic composition, thus, provide a most informative archive that complements the records obtained from the coexisting rocks. The spatial distribution of formation waters, variability of their properties, and the lithology of the host rocks provide the tools for the identification of sea regressions and transgressions in selected study areas. REFERENCES
4 FORMATION WATERS TAGGED BY EVAPORITIC SALTS REFLECT LOWLAND PALEO-GEOGRAPHY Having reached the conclusion that formation waters were mainly groundwaters recharged into rocks that were exposed during sea regression phases, the question comes up what type of continental environment prevailed at those ancient landforms that made up the top of active sediment collecting basins?. The answer is coded in the chemical composition of the discussed formation waters: (a) the formation waters are saline, indicating the respective paleo-climate was arid; (b) C1 is the dominant ion, the weight ratio of Cl/Br is 20 to 200, i.e. different than in halite or seawater, but well in the range of evaporitic end-brines, disclosing prevalence of an evaporitic environment of sabkhas and lagoons, left behind each time the sea receded. The formation waters are, thus, recharged paleo-groundwater,enriched
Egeberg, P. K. & P. Aagaard 1989. Origin and evolution of formation waters from oil fields on the Norwegian shelf. Applied Geochemistry, 4: 131- 142. Fridman, V., Mazor, E., Becker, A., Avraham D. & E. Adar 1995. Stagnant aquifer concept: 111. Stagnant mini- aquifers in the stage of formation, Makhtesh Ramon, Isral. J. of Hydrology, 173: 263-282. Hitchon, B., Billings, G. K. & J. E. Klovan 1971. Geochemistry and origin of formation waters in the western Canada sedimentary basin -111. Factors controlling chemical composition. Geochim. Cosmochim Acta, 35: 567-598. Mazor, E. 1995. Stagnant aquifer concept I. Large scale artesian systems - Great Artesian Basin, Australia. J. of Hydrology, 174: 2 19-240. Mazor, E., Gilad, D. & V. Fridman 1995. Stagnant aquifer concept: 11. Small scale artesian systems - Hazeva, Dead Sea Rift Valley, Israel. J. of Hydrology, 173: 24 1-26 1. Mazor, E. 1998. Allochthonouos ions dissolved in recent and fossil groundwaters: identification and origins. . The 9th Int. Symp. Water-Rock Interaction, Taupo, New Zealand, April, 1998; 169-172 Rotterdam: A. A. Balkema.
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Weaver, T. R., Frape, S . K: & J. A. Cherry 1995. Recent crossformational fluid flow and mixing in the shallow Michigan Basin. GSA Bull., 107: 697-707.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets 8, Zeitlinger, Lisse, ISBN 90 2651 824 2
The influence of basement fluid upwelling and diagenesis on CaCO, stability in sediments from the eastern flank of the Juan de Fuca ridge Christophe Monnin MLcanismes de Transfert en Gkologie, CNRSIUPS, 38 rue des Trente-Six Ponts, 31400 Toulouse, France
C .Geoffrey Wheat Global Undersea Research Unit, University of Alaska P.O. Box 47.5, Moss Landing, CA 95039, USA
Mike M.Mott1 Department of Oceanography, University of Hawaii at Manoa, 1000 Pope Road, Honolulu, HW 96822 USA
Sabine Balleur Laboratoire de GLologie, CNRSIENS, 24 rue Lohmond 7500.5 Paris, France
ABSTRACT: Basement fluid discharge in zones associated with basaltic outcrops and seamounts of the eastem flank of the Juan de Fuca ridge (NE Pacific) has been investigated through heat flow measurements and coring during the Retroflux cruise (September 2000). Porewaters of the 93 collected cores have been analyzed onboard for chlorinity, Ca, Mg, pH, alkalinity, and nutrients. Upwelling basement fluids have an alkalinity lower than that of seawater. In contrast, porewaters from cores without porewater flow have alkalinities that increase with depth resulting from the oxidation of organic matter and the reduction of sulfate during diagenesis. Thus alkalinity-depth profiles reflect the competition between alkalinity production from diagenetic reactions and the transport of an alkalinity-depleted basement fluid. Thermodynamic calculations reveal that 1) cores that lack porewater flow are increasingly supersaturated with respect to calcite and aragonite as a function of depth in response to the increase in alkalinity, 2) upwelling of basement fluid at speeds of up to several c d y r results in a porewater carbonate system that is in equilibrium with aragonite, and 3) extremely rapid flow (10 c d y r ) for core PC-GC28 taken at Wuzza Bare induces the transition from equilibrium with aragonite to equilibrium with calcite.
1 INTRODUCTION Fluid circulation through mid-ocean ridge flanks plays an important role in cooling of the oceanic lithosphere and in mass transfer between the crust and the oceans. The eastern flank of the Juan de Fuca ridge has been the focus of numerous cruises (e.g. Davis et al. 1992, 1997, Mottl et al. 1998) to study the patterns of such a low-temperature hydrothermal system in young sedimented crust. During ODP Leg 168 (Davis et al. 1997), a series of ten boreholes were drilled along a 100 km-long transect perpendicular to the ridge axis. It was found that temperature at the sediment-basement interface increases from west to east from 16°C at Site 1023 to 63°C at Sites 1026 and 1027. This temperature increase is accompanied by a change in composition of basement fluids that become enriched in calcium and depleted in magnesium when compared to seawater. Data collected during ODP Leg 168 further suggest a simple hydrologic regime in which cold seawater penetrates the seafloor through fractures near the ridge axis and flows eastward in a permeable layer of the uppermost basalts buried under a thick sedimentary cover. The relatively impermeable sediments act as a thermal blanket and prevent basement fluids from discharging to the deep ocean. Fluid discharge has been found to occur either as fo-
cused flow in springs such as those found atop a basaltic mount called Baby Bare, located about lOOkm east of the ridge axis (Mottl et al. 1998); or as diffuse seepage at basement bathymetric highs, such as at ODP Sites 1030 and 1031, where the sediment cover is only 40 m thick (Davis et al. 1997). Fluid discharge resulted from drilling at ODP Site 1026, showing that basement buried under 250m of sediment is overpressured (Fisher et al. 1997). It thus appears that basaltic outcrops and basement highs covered by thin sediment layers act as preferential pathways for basement fluid discharge to the deep ocean. The simple hydrologic regime suggested by data from ODP Leg 168 is challenged by the composition of basement waters collected at four sites along a 3.5 my-old, partly buried ridge about 100 km east of the Juan de Fuca ridge. These waters become increasingly altered relative to seawater from south to north (Wheat et al. 2000). It has been inferred that this SN circulation is decoupled from the east-west circulation inferred from ODP Leg 168. These observations demonstrate the need for more information about the location of fluid discharge and recharge sites on the eastern flank of the Juan de Fuca ridge, to further constrain seafloor hydrology and estimates of mass fluxes between the oceanic crust and the oceans. Additional constraints come from heat flow 639
data on one hand and from the evolution of basement fluid composition on the other. Compositional changes result from the combined effects of molecular diffusion, advection, and chemical reactions. Basement water compositions are often deduced from those of pore water in sediments immediately overlying the basaltic basement. In this paper we examine some of the shipboard pore water data collected during the Retroflux cruise (September 2000) on the eastern flank of the Juan de Fuca ridge, in order to understand the role of the sediment pore water composition on calcium carbonate precipitation and dissolution.
the estimated residence time of the fluids within the crust (Mottl & Wheat 1994). These fluids have an alkalinity lower than that of seawater, making ridgeflank alteration at low temperature a potentially important sink for carbon on a global scale (Sansone et al. 1998). 3 THE RETROFLUX CRUISE The Retroflux cruise took place in September 2000. It had been planned to use heat flow measurements and coring to map several outcrops, subcrops, and seamounts detected during previous seismic surveys, that could be potential locations of fluid exchange between the crust and the oceans. Heat flow measurements were done at night and coring was done the next day at sites of heat flow anomalies. Ninetythree cores (up to 7 meters in length) were collected during this expedition. Pore waters were extracted at sea by centrifugation and analysed onboard for chlorinity, Ca, Mg, pH, alkalinity, and nutrients. Operations were concentrated in three main areas. The first, called First Ridge, is a buried basement ridge near ODP Sites 1030 and 1031, 39 km from the ridge axis and midway along the 103-km ODP Leg 168 transect. A second buried basement ridge is located near Sites 1026 and 1027 at the eastern end of ODP Leg 168 transect, and close to the Mama, Papa and Baby Bare seamounts. Three sites were cored there: Isita Bare, Wuzza Bare and Zona Bare. The third study area is the location of two large seamounts (Grizzly Bare and Grinnin' Bare) 25-40 km south of the eastern end of the ODP transect.
2 CHEMICAL COMPOSITION OF BASEMENT FLUIDS ON THE EASTERN FLANK OF THE JUAN DE FUCA RIDGE Uncontaminated basement fluids have been directly sampled at ODP Site 1026 during ODP Leg 168, where fluid discharge took place during and after drilling (Fisher et al. 1997, Davis et al.1997). In the same area, 25°C springs have been discovered atop a basaltic seamount called Baby Bare (Mottl et al. 1998). Fluids discharging at this location have been recovered by manned submersible, but the samples are a mixture of bottom seawater and basement fluid. In this case the composition of basement fluids is obtained from plots of a given dissolved element versus Mg. Extrapolation to zero Mg concentration gives the end-member composition (Wheat & Mottl 2000). Several areas of fluid seepage through sediments have been also found on the eastern flank of the Juan de Fuca ridge. Seepage is identified from the characteristic diffusion-advection shape of profiles of pore water composition vs. depth. For example, dissolved Mg decreases rapidly with depth and reaches a constant value representative of basement fluid (Wheat & Mottl 1994). Basement fluid composition can thus be obtained from the composition of pore waters sampled in shallow cores (a few meters in length) retrieved in zones of upwelling areas. This method supposes that molecular diffusion and advection (upwelling of basement fluids) dominate and that reactions taking place within the sediment do not alter solute concentrations. The ideal tracer would be inert with respect to the sediment. We here investigate the role played by these phenomena on calcium carbonate stability, using an analysis of the calcium content, pH, and alkalinity of sediment pore waters. Data collected earlier (Wheat & Mottl 1994, Mottl et al. 1998, Elderfield et al. 1999, Wheat & Mottl2000) show that basement fluids of the eastern flank of the Juan de Fuca ridge are enriched in calcium and depleted in magnesium, consistent with basalt-seawater interactions at low temperatures, to an extent that depends mainly on temperature and
4 COMPOSITION OF PORE WATERS Upwelling of basement water is inferred from the characteristic profiles of Mg and Ca (Fig. 1). In cores with no flow Mg is constant at the seawater value, whereas Mg decreases rapidly with depth in upwelling zones. The behavior of Ca is symmetrical to that of Mg as a result of Ca-Mg exchange during basalt alteration. A range of advection velocities has been observed in the 93 cores collected during the Retroflux cruise: the higher the upwelling speed, the faster the decrease in Mg (Fig. 1). For Core PC-GC 28 taken at Wuzza Bare, the pore water Mg concentration (3mM) is less than 10% of that of normal seawater only 5 cm below the sediment-water interface, while its Ca content increases to 63 mM. This leads to an estimated upward seepage velocity of about 1Ocdy. ODP Leg 168 data showed that sediments in this region undergo diagenetic reactions that lead to a 640
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GC40 0 A A
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GC48 GC 56 GC-PC61 GC-PC62 Gt-PC2 GC-PC 28
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2 3 4 Depth (mbsf)
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Figure 2. Aragonite saturation index for cores with increasing seepage velocities. The dashed line indicates equilibrium (SI =l).
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.
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Figure 1. Porewater Mg and alkalinity versus depth for cores with increasing seepage velocity.
total depletion of pore water sulfate within the upper 20 meters below seafloor (mbsf) or less, with a simultaneous alkalinity production due to organic matter oxidation (Davis et al. 1997, Monnin et al. 2000). We likewise observe this alkalinity increase in the Retroflux cores collected in no-flow zones (Fig. 1). When a low-alkalinity basement fluid percolates upward through the sediment, the alkalinity of the pore waters does not increase as much with depth, to the point where the effect of diagenesis is totally masked and alkalinity decreases with depth. For Core GC48 the two opposite effects balance and lead to a pore water alkalinity almost constant at the seawater value (Fig. 1).
5 CALCITE AND ARAGONITE SATURATION INDICES TN POREWATERS We have used models of the carbonate system in seawater embodied in the computer code CO2SYS (Lewis & Wallace 1999) to calculate the in situ (2"C, 260 bars) pore water saturation states with respect to calcite and aragonite from measured values of pH, alkalinity, and Ca concentration (Monnin et al. 2000). During Retroflux two bottom water samples were collected. Our calculations show that their calcite and aragonite saturation indices are 0.93 and 0.60 respectively, in accordance with the general
0
1
2 3 4 Depth (mbsf)
5
6
Figure 3 - Calcite and aragonite saturation indices for Core PC-GC 28 showing high upwelling fluid velocity (10cdy). Note calcite equilibrium below 2mbsf.
CaC03 saturation state of the Pacific Ocean (Morse & MacKenzie 1990).
Results for the 93 cores show that when equilibrium is reached it is with aragonite rather than calcite. Figure 2 shows cores with increasing flow velocities: more rapid flow reduces aragonite supersaturation toward equilibrium at depth. Core PC-GC 28 collected at Wuzza Bare shows a very high upwelling speed of basement fluid, the alkalinity of which is 0.4 meq/kg. This induces equilibrium with calcite at depths below 2mbsf (Fig. 3). Our calculations also indicate a calcite-aragonite transition in Wuzza Bare Core GC 79 (no flow, constant alkalinity, small pH increase from 7.8 at the seafloor to 7.9 at Imbsf). 6 DISCUSSION: REASONS FOR CALCIUM CARBONATE SUPERSATURATION Differences in basement fluid composition and local variations in the extent of diagenesis are primary factors influencing the chemistry of pore waters and water-rock interactions. But there are a number of
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causes linked to the mechanisms of calcite/aragonite precipitation that can lead to supersaturation (see discussion in Morse & MacKenzie 1990). Besides phosphate and sulfate, the magnesium ion is a strong inhibitor of calcite precipitation. Our data deal with pore waters with Mg concentrations varying from the seawater value 53 mM to about 3 mM. Equilibrium calculations should be discussed in light of the formation of magnesian calcites or even dolomite, and of the relative stability of aragonite and calcite. On the other hand, the increase in alkalinity should favor CaC03 precipitation, we yet observed steady increases in calcite and aragonite supersaturation, generally indicating a kinetic hindrance to precipitation. Core GC 87 was retrieved at Papa Bare in an area where carbonate concretions have been observed and where fluid flow occurs at a moderate rate. Our results show a plateau in the SI values, linked to a pH that changes from 7.8 at the seafloor to a level value of 8.7 below 1 mbsf. This would indicate that precipitation occurs only when a given level of supersaturation is reached. This study is for now limited to the analysis of the Retroflux shipboard pore water data. More insight will be gained from trace element data and from the observation and analysis of solids in the sediment samples.
Mottl, M.J., Wheat, C.G., Baker, E., Becker, N., Davis, E., Feely, R., Greham, A., Kadko, D., Lilley, M., Massoth, G., Moyer C. & F. Sansone 1998. Warm springs discovered on 3.5 Ma oceanic crust, eastern flank of the Juan de Fuca Ridge. Geology, 26( 1): 5 1-54. Sansone, F.J., Mottl, M.M., Olson, E.J., Wheat, C.G. & M.D. Lilley 1998. C02-depleted fluids from mid-ocean ridge flank hydrothermal springs. Geochim. Cosmochim. Acta. 62( 13): 2247-2252. Wheat, C.G. & M. Mottl 1994. Hydrothermal circulation, Juan de Fuca Ridge eastern flank: factors controlling basement water composition. J. Geophys. Res., 99: 3067-3080. Wheat, C.G., & M. Mottl 2000. Composition of pore and spring waters from baby bare: global implications of geochemical fluxes from a ridge flank hydrothermal system. Geochim. Cosmochim. Acta. 64(4): 629-642. Wheat, C.G., Elderfield, H., Mottl, M.M. & C. Monnin 2000. Chemical composition of basement fluids within an oceanic ridege flanc: implications for along-strike and across-strike hydrothermal circulation. J. Geophys. Res., 105(B6): 13437- 13447.
REFERENCES Davis, E.E., Chapman, D.S., Mottl, M.M., Bentowski, W.J., Dadey, K., Forster, C., Harris, R., Nagihara, S., Rohr, K., Wheat, C.G. & M. Whiticar 1992. FlankFlux: an experiment to study the nature of hydrothermal circulation in young oceanic crust. Can. J. Earth Sci., 29: 925-952. Davis, E.E., Fisher, A.T., Firth, J.V. et al. 1997. Proc. ODP, Init Repts., 168, College Station, TX (Ocean Drilling Program). Elderfield, H., Wheat, C.G., Mottl, M.M., Monnin, C. & B. Spiro 1999. Fluid and geochemical transport through oceanic crust: a transect across the eastern flank of the Juan de Fucaridge. Earth Planet. Sci. Lett., 172: 151-165. Fisher, AT., Becker, K. & E.E. Davis 1997. The permeability of young oceanic crust east of the Juan de Fuca Ridge, as determined using borehole thermal measurements. Geophys. Res. Let., 24: 1311-1314. Lewis, E. & D. Wallace 1999. CO2SYS: program developed for CO2 system calculations, the program and user's guide are available online on the CDIAC (Carbon Dioxyde Information and Analysis Center) web site (http://cdiac.esd.ornl.gov). Monnin, C., Karpoff, A.M. & M. Buatier 2000. Calcium carbonate stability in the sediments of the eastern flank of the Juan de Fuca ridge, in Fisher A. Davis E.E. and Escutia C. (Eds.). Proc. ODP Sci. Results, 168: 95-103, College Station TX (Ocean Drilling Program). Morse, J.W. & F.T. Mackenzie 1990. Geochemistiy of sedimentaiy carbonates. Elsevier, 707 p. Mottl, M.J. & C.G. Wheat 1994. Hydrothermal circulation through mid-ocean ridge flanks: fluxes of heat and magnesium. Geochim. Cosmochim. Acta, 58( 10): 2225-2237.
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Water-Rock Interaction 2001, Cidu (ed.),02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Fluid flow in the Cantabrian Zone (NW-Spain) - contributions to the diagenetic evolution JSchneider, T.Bechstadt & S .Zeeh Geological-PaleontologicalInstitute, University of Heidelberg, Irn Neueizheimer Feld 234, 69120 Heidelberg, Germany
M .Joachimski Institute of Geology a i d Mineralogy, University of Erlangen, Schlossgarten 5, 91054 Erlangen, Germany
ABSTRACT: The Devonian carbonate formations of the Somiedo-Correcillas Unit (NW-Spain) are the main topic of this diagenetic study. The objective is to reconstruct the different fluid flow events to investigate the basin evolution before, during and possibly after the Variscan orogeny. The starting point of our research is a cement stratigraphy to distinguish different fluid events affecting the Devonian to Lower Carboniferous sedimentary series. During the basin stage, rock buffered fluids circulated, precipitating cements characterized by decreasing 6180-values. The formation of non-ferroan calcites marks the onset of deep burial diagenesis and is followed by the precipitation of ferroan saddle dolomite. Finally, ferroan blocky calcites crystallized, showing the most negative 6"O-values which indicate deepest burial conditions. Non-ferroan saddle dolomite precipitated in post-Variscan time or later. Two generations of blocky calcites formed subsequently from solutions with low fluid-rock interaction.
1 INTRODUCTION Our diagenetic studies within Lower to Middle Devonian carbonates (La Vid, Santa Lucia, and Portilla Formations) of the Somiedo-Correcillas Unit of the Cantabrian Zone show a complicated diagenetic stratigraphy of deep burial cements. This cement stratigraphy was established by a combination of petrographic, cathodoluminescence (CL), stable isotope, and trace element analyses.
Devonian and Carboniferous. Units deposited further to the north, in close proximity to the Cantabrian block, are characterized by less continuous Devonian successions. Although the different palaeogeographic units were thrusted and folded during the Variscan orogeny, temperatures in the Siluro-Devonian to Lower Carboniferous sediments of the southern Cantabrian Zone mostly did not exceed the field of diagenesis (Keller & Krumm 1993).
2 GEOLOGICAL SETTING The Cantabrian Zone belongs to the northern part of the Iberian Variscan orogen (Fig. 1). It consists of different palaeogeographic units. The Sobia-Bodon, the Aramo, and the Somiedo-Correcillas tectonic units of the Cantabrian zone (Fig. 2) display a shelf to basin transition (Julivert 1971). The distal Somiedo-Correcillas Unit is characterized by the most complete sediment series, spanning Lower Cambrian to Namurian strata (Keller & Krumm 1993). In this unit, carbonates of the Lower Cambrian are overlain by siliciclastic sediments of Cambrian to Silurian age. The Lower Devonian La Vid Formation marks the return to carbonatedominated successions that prevailed during the
Figure 1. Tectonic framework of the Cantabrian Zone (after Julivert 1971, and Perez-Estaun et al. 19SS), with the working area in the Somiedo-Correcillas Unit.
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Figure 2. Stratigraphic section of the Cantabrian Zone showing the pre-orogenic (Cambrian to Devonian) and syn-orogenic successions (according to Marcos & Pulgar, 1982). The thrust fronts to the left of each wedge represent only their symbolic positions. For the position of the cross-section see Figure 1.
3 ANALYTICAL RESULTS Different calcite and dolomite cement phases can be distinguished in the Somiedo-Correcillas Unit. Their distribution within the stratigraphic succession of Devonian and Carboniferous rocks enables us to determine a relative cement stratigraphy. The carbonate formations show a diagenetic evolution from shallow to deep burial stages and finally post-orogenic cement precipitation. Early diagenetic calcite cements occur in fossil moulds, and as fibrous cements within cavities. The latter show the same CL zonation as syntaxial overgrowths, from non- to orange luminescent. Rhombic dolomites were precipitated in cavities and along compactional stylolites during the final stage of shallow burial diagenesis. A significant hint for a strong diagenetic overprint of the whole rock samples is the large spread in 6l80 (Fig. 3) ranging from -2 to - 8 % 0 VPDB. Values around -2%0were measured for lowmagnesium calcite of Devonian brachiopod shells that may represent 6180-values in equilibrium with Devonian seawater. The first cement phase indicating deep burial conditions is a non-ferroan calcite (NFC) observed on the inner rim of fossil moulds and fissures. This orange-coloured luminescent phase shows 6l80- and 613C-values of -6 to -10%0 and around +0.5%0 VPDB, respectively. Non-luminescent iron-rich saddle dolomite (ISD) can be observed in several Devonian and Early Carboniferous formations up to the turbidite series of the Olleros Formation (Namurian B). The 644
cloudy crystals show an undulatory extinction, affected by tectonic stress. This cement has a high iron content of about 2% and a low MdFe-ratio of 0.006. ISD crystallized in cavities and in fissures. In the upper La Vid Formation, relics of ISD are intergrown with barite. Both minerals are fractured and crosscut by micro-fissures. 6l80- and 613Cvalues vary from -7 to -10%0 and -1 to +4%0 VPDB, respectively. The iron-rich saddle dolomite (ISD) is crosscut by iron-rich blocky calcite (IBC) reflecting varying CL colours. These iron-rich calcites fill fractures and cavities and show the most depleted 6180-values ranging from -13 to -l6%0 VPDB. 613C-valuesare around - 1 to +1 %O VPDB. A generation of non-ferroan saddle dolomite (NFSD) is interpreted to be later in origin. No tectonic overprint can be recognized. NFSD might therefore be post-Variscan or even younger. The cement crystals show a zonation from dull to bright red luminescent. The 6180-valuesare between -6 to -8%0, whereas 613C is around +2%0 VPDB, slightly heavier than the isotopic values of the iron-rich saddle dolomite. However, manganese and iron contents of NFSD are significantly lower (MdFeratio of 0.014). The non-ferroan saddle dolomite may be followed by clear blocky calcite cements (CBC) that are non-luminescent. The crystals crosscut the NFSD and crystallized in cavities after the precipitation of earlier calcites (IBC). This cement phase has low iron contents. 6l80 varies from -7 to - 1 1 % 0 VPDB. 613~-values(-1 to -4.5%0 VPDB) are depleted in 13C in comparison to the respective host rock.
Figure 3. 613C/6'80-isotopeplot of the different carbonate cements in the Somiedo-Correcilla Unit.
increasing burial, record lighter 6'*0- and 6I3Cvalues and become simultaneously enriched in iron. Cement IBC is interpreted as the latest cement of this stage formed during deep burial, possibly in syn-Variscan time.
The latest cement generation is a nonluminescent calcite which occurs as sinter and blocky crystals in karst moulds (KBC). 6"Ovalues range from -7 to -9%o VPDB. 613~-values are strongly depleted with values ranging from -7 to -10%0 VPDB.
4.2 Post- Variscan Cementation The field and petrographic observations of the subsequent NFSD indicate its post-Variscan emplacement. NFSD differs significantly from the preVariscan saddle dolomite (ISD) by its minor manganese and iron contents as well as by the cathodoluminescence and zonation patterns of the crystals. However, the 6'3C-values suggest that the diagenetic system is still rock-buffered. The two latest cement generations are depleted in I3C. This argues for a change to a diagenetic system partly open to carbon from external sources. The slightly negative 613C-valuesof the CBC can be attributed to the contribution of oxidized organic carbon. A possible source rock might be organic matter from Stephanian coal basins. Precipitation of the KBC phase is clearly related to karstification as evidenced by field observations and low 6I3C-values. The latter are due to the incorporation of isotopically light CO2 from karst cave systems. The fluid rock interaction between the KBC and the Devonian host rock is relative1 small, indicated by the distinct differences in C.
4 DISCUSSION The described cement succession sheds some light on the diagenetic evolution of the SomiedoCorrecillas Unit. Two clearly different fluid systems were present. Carbonate cements, precipitated during the first diagenetic stage, indicate an open system due to their 6"0-values, decreasing with increasing temperature, whereas the 6I3C-values give evidence for a rock buffered diagenetic system. Post-Variscan carbonate cements are characterized by depleted 6' 3C-values suggesting a diagenetic system partly open for carbon derived from the remineralisation of 13Cdepleted organic carbon. 4.1 Pre- to syn- Variscan stage
6l80 signals of whole rock samples range from the assumed isotopic composition of Devonian seawater to highly depleted values. The same trend is observed for the different cement generations (NFC, ISD, and IBC). Cements precipitating during
Y
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REFERENCES Julivert, M. 197 1 . Decollement tectonics in the Hercynian Cordillera of Northwest Spain. Am. J. of Sci. 270: 1-29. Keller, M. & S. Krumm 1993. Hercynian versus Caledonian and Precambrian metamorphic events in the Cantabrian Mountains, Northern Spain. Z. dt. geol. Ges. 141:88-103. Marcos, A. & J.A.Pulgar 1982. An approach to the tectonostratigraphic evolution of the Cantabrian foreland thrust and fold belt, Hercynian Cordillera of NW Spain In J. Kullmann, R. Schonenberg, & J. Wiedmann (eds), Subsidenz-Entwicklung im Kantabrischen Variszikum und an passiven Kontinentalrandern der Kreide; Teil 1 , Sonderforschungsbereich 53 Tubingen: Neues Jahrbuch f i r Geologie und Palaontologie. Abhandlungen: Stuttgart, Federal Republic of Germany, E. Schweizerbart’sche Verlagsbuchhandlung, 256-260. Perez-Estaun A., Bastida, F., Alonso, J.L., Marquinez, J., Aller, J., Alvarez, M.J., Marcos, A. & J.A. Pulgar 1988. A thin-skinned tectonics model for an arcuate fold and thrust belt; the Cantabrian Zone (Variscan IberoAnnorican Arc). Tetonics 7 5 17-537.
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Sulfate reduction rates and organic matter composition in sediments off Namibia C.J.Schubert, T.G.Ferdelman, B .B .J@rgensen& GKlockgether Department of Biogeochemistry, Max Planck Institute for Marine Microbiology,0-28359 Bremen, Germany
ABSTRACT: Sulfate reduction rates in relation to organic matter composition have been investigated in sediments off Namibia. Sediment samples were taken with RV Poseidon in 1999 on a transect from 40 to 2000 m water depth and therefore span a wide range from the shelf to the deeper slope. Organic carbon concentrations range from 3 to 14% and C/N values from 3 to 22. Low organic carbon values with low C/N values are measured on the shelf whereas higher organic carbon and C/N values occur on the deeper slope. Total chlorins, degradation products of chlorophyll, show higher concentrations on the shelf are interpreted together with lower C/N values as indicating the deposited organic matter to be more labile than on the deeper slope. This is also confirmed by a new chlorin index CI that is lower on the shelf and higher on the slope indicating the difference in lability of the organic material. further destroyed by chemical treatment. Namibian shelf and slope sediments provide an excellent environment for test the relation between SRR, TOC, C/N, and Chlorins since the area from where we got our cores is far away from any river input, e.g. the closest river the Orange is about 800 km away. The hinterland of Namibia is mainly represented by desert or land with little vegetation. We therefore assume the input of terrestrial organic carbon to the core sites as negligible if existing at all. Changes in C/N ratios should therefore not indicate variations in the input of marine versus terrestrial organic carbon but should indicate the difference in the state of diagenesis of the buried organic material since nitrogen containing compounds are much more labile than non-nitrogen organic compounds. Hence, higher C/N ratios should indicate more refractory “old” organic material, whereas low C/N ratios stand for labile “fresh” organic material. In the shallow sediments with c 200 m water depth, sulfate reduction is the main process of organic matter degradation since oxygen is not available in the bottom water. Methanogenesis is only observed in deeper sections of recovered gravity cores but does not play a role in our short cores.
1 INTRODUCTION The shelf region off Namibia with the Benguela upwelling system is characterized by extremely high phytoplankton productivity over the continental shelf and by oxygen depletion and nitrate enrichment in the lower water column. The sediment cores underlying the core of upwelling are very rich in organic material that is preferentially degraded anaerobically by sulfate reduction and deeper in the sediments by methanogenesis. As a result, huge amounts of hydrogen sulfide and methane build up in the sediments and are eventually released into the atmosphere. We compared sulfate reduction rate measurements with the organic carbon composition of sediment cores. We investigated the organic carbon composition in respect to its lability and refractory, respectively. Until now, we have measured TOC content, C/N values and total chlorins but we aim at getting more parameters such as carbon isotope values and protein content as well as the lipid content. Additionally, we tested a new chlorin index indicating the relatively freshness of the organic material deposited on the shelf. This index is the relation between chlorins that could be measured fluorometrically after the extraction with acetone and the extract that has been successively treated with hydrochloric acid. The index is low when the hydrochloric acid treatment leads to a severe destruction of the relatively fresh chlorins, but is high when the chlorins are already heavily degraded and can not be
2 METHOD Seven multicorer cores (Stations 2,4,7,8,9,11,12) from the shelf off Namibia were collected in May 1999. The sediment cores were recovered from a 647
depth of 40 to 2000 m water depth and span therefore a wide range from the shelf to the deeper slope (Fig. 1). Total carbon (TC) was determined by combustion/gas chromatography (Carlo Erba NA-1500 CNS analyzer) with a precision of 51.2 % on the bulk samples. Carbonate carbon (CC) was measured on a Coulomat with a precision of 21.1 %. The organic carbon (TOC) content was calculated by subtraction of TC and CC. For the determination of chlorins, which include a whole suite of degradation products of chlorophyll, freeze dried sub-samples (- 0.6 g) were extracted by threefold sonication and centrifugation in acetone. The samples were cooled with ice under low light conditions during extraction to prevent decomposition of the chlorins. Sediment extracts were measured fluorimetrically (Hitachi F-2000 fluorometer) immediately after extraction. Chlorophyll a (Sigma) which was acidified with a few drops of hydrochloric acid was used as a standard. The precision of the method was 210 %. To calculate the freshness index CI we have acidified the sample extract after the measurement with a few drops of hydrochloric acid (10%) and remeasured. The CI index is the relation between acidified sample to non-acidified sample. Sulfate reduction rates were determined using a whole-core 35SO4-2 incubation method (Jgrgensen 1978). Reduced 35s was determined using the onestep Cr-I1 destillation (Fossing & Jgrgensen 1989). Sulfate was measured using nonsuppressed ion chromatography with a Waters 510 HPLC pump, Waters WISP 712 autosampler, Waters IC-Pak anion exchange column (50 x 4.6 mm), and a Waters 430 conductivity detector (Ferdelman et al. 1997).
Figure 1. Core locations off Namibia
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3 RESULTS At Station 2 in 40 m water depth, organic carbon concentrations vary from 3 % to 6 % with relatively small changes downcore (Fig. 2). C/N values vary from 7 to 8.5 which basically could be interpreted as a mainly marine organic matter source. Total chlorin concentrations are only available for the lower part of the core and vary from 1.8 to 2.8 * 104 ng/g (CI= 0.53, mean of 0-15 cm), but are the highest measured in all of our cores. The areal sulfate reduction rate of 20.3 mmol/m2/d (integrated over the uppermost 15 cm) is amongst the highest values measured in the world ocean so far. At Station 4 in 100 m water depth, TOC are highly variable between 5 and 14 % including the highest value of all our cores (Fig. 2). C/N values increase steadily from the surface downcore from 6.5 to 8.5. Total chlorin concentrations vary highly and inversely to the TOC concentrations from 0.4 to 2.4 * 104 ng/g (CI= 0.45). The areal sulfate reduction rate is still very high with 9.1 mmol/m2/d. At Station 7 in 200m water depth, TOC values vary from 9 to 11.5 %, C/N values are very high around 13 to 17, and total chlorin concentrations decrease rapidly downcore from 1.9 to 0.9 * 104 ng/g (CI= 0.84). The areal sulfate reduction rate is 3.1 mmol/m2/d. At Station 8 in 300 m water depth, organic carbon concentrations are relatively stable and vary from 4 to 5 % (one value at 1.5 cm shows 3 %, Fig. 2). C/N values vary between 9 and 12 with one value of 6.5 at 1.5 cm. Total chlorin concentration range between 0.6 to 1.4 * 104 ng/g (CI= 0.90). The areal sulfate reduction rate is 1.5 mmol/m2/d. At Station 9 in 400 m water depth, the lowest TOC values are measured (1 to 3 %, Fig.2), the C/N values are following the TOC values and vary between 3 and 8.5. Total chlorins are low and vary between 1.2 * 104 ng/g at the surface to 0.3 * 104 ng/g deeper in the core (CI= 0.92). The areal sulfate reduction rate is 0.4 mmol/m2/d. At Station 11 in 1000 water depth, TOC values are somewhat higher and very stable between 3 and 4 %. C/N values follow the TOC signal and vary between 7 and 9. Total chlorin concentrations are only available for the uppermost samples which vary from 0.7 to 1.1 * 104 ng/g (CI= 0.85). The areal sulfate reduction rate is 0.8 mmol/m2/d. At Station 12 in 2000 m water depth, TOC values increase dramatically and vary between 8 and 11 %. C/N values in this core are the highest of all measured cores and vary between 17 and 22. Total chlorin concentrations are around 2.0 * 104 ng/g at the surface but decrease sharply and reach values below 0.6 * 104 ng/g below 3 cm core depth (CI= 0.85). The areal sulfate reduction rate is 0.7 mmol/m2/d.
Figure 2. Total organic carbon and chlorins concentrations, TOC/Ntot ratios of cores 2,4, 7, 8, 9, 11, 12.
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4 DISCUSSION Concerning the organic carbon concentrations there is no general trend in the cores from the shallowest to the deepest station. Therefore, we have to look for the other parameters. The two shallowest cores (Station 2 and 4) show relatively similar low C/N values and the highest total chlorin concentrations. This is interpreted as indicating relatively freshly deposited organic material from the water column. Higher C/N values are found in cores from the deeper stations (with the exception of Station 9) which might indicate that the organic material is more refractory i.e. diagenetically more altered. This is explained by relatively high bottom currents responsible for the transport of shelf material down the slope. Support comes also from the total chlorin concentrations that are decreasing drastically from the shallower to the deeper cores. Whereas the shallower cores show values of 2 to 3 * 104 ng/g, the deeper cores only show values in the 0.5 to 1 * 104 ng/g range. Additionally, the relatively freshness of the organic material deposited on the shelf is also indicated by the new chlorin index that range between 0.4 and 0.5 on the shelf and between 0.8 and 0.9 deeper on the slope. Area1 sulfate reduction rates (SRR, 0-15 cm) in the cores vary between 0.43 and 20.3 mmol/m2/d with the higher values being among the highest SRR ever measured. Whereas the organic carbon concentrations show almost no correlation (r2 = 0.06) with the SRR, the total chlorin concentrations as well as the chlorin index CI correlate very well with the SRR with an r2 of 0.91 and 0.83, respectively. This confirms further, as indicated before in sediments of Chile (Schubert et al. 2000) that sulfate reduction (as a measure of organic reactivity) in shelf and slope sediments is strongly determined by the input of fresh organic material, here represented by the relative labile chlorine fraction. REFERENCES Ferdelman, T.G., Lee, C., Pantoja, S., Harder, J., Bebout, B.M. & H. Fossing 1997. Sulfate reduction and methanogenesis in a Thioploca-dominated sediment off the coast of Chile. Geochim. Cosmochim.Acta 61(15): 3065-3079. J~rgensen,B.B. 1978. A comparison of methods for the quantification of bacterial sulfate reduction in coastal marine sediments. I. Measurement with radiotracer techniques. Geomicrobiol. J . 1: 29-47. Fossing, H. & B.B. J~rgensen1989. Measurement of bacterial sulfate reduction in sediments: Evaluation of a single-step chromium reduction method. Biogeochem. 8: 205-222. Schubert, C.J., Ferdelman, T.G. & B. Strotmann 2000. Organic matter composition and sulfate reduction rates in sediments off Chile. Organic Ceochemistiy 31: 351-361.
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Origin of Ordovician organogenic dolomite concretions: Significance for the 6 l 8 0 of Lower Paleozoic SMOW J.Shah & R.Hesse Earth and Planetary Sciences, McGill University, Montreal, Quebec, H3A 2A7, Canada R. Hesse presently: Institlit fiir Geologie, Miizeralogie und Geophysik, Ruhr- Universitat Bochum, 0-44801 Bochuin Germany
S .Islam Akzo Sult Inc., P.O. Box 6920, Cleveland, Ohio, 44 101, U S A .
ABSTRACT: Estimated depth of formation of authigenic dolomite concretions, based on centre-to-margin variations in minus-cement porosity, is between 4 and -200 m below seafloor (mbsf), mostly 4 - 2 5 mbsf, in Middle Ordovician Cloridorme Formation, Quebec (80-90% to 45-75% porosity), and >350 mbsf (25-17% porosity) in the Lower Ordovician Levis Formation. Outward decreasing 6l3Cp~*values (from 10.2 to 2.1%0) suggest precipitation in the methane-generation zone with an increasing contribution of light carbonate supplied by advection from thermocatalytic reactions at greater depth. Corresponding 6l8O values are anomalously low (centre-to-margin variations: -0.4 to -7.5%0). They give reasonable burial temperatures only assuming that Ordovician paleo-SMOW was negative (-6%0) and, in addition, or anic-matter decomposition, volcanic-ash alteration andor advection of 6I8Odepleted water have lowered the 6l 0 of the pore waters.
8
1 INTRODUCTION
the discrepancy of isotopic temperature estimates and burial depths (e.g., Mozley & Bums 1993); (3) the primary or replacement origin of organogenic dolomite (Lawrence 1991). These questions were addressed in studies of dolomite concretions and beds from the Ordovician Cloridorme and Levis formations of the Quebec Appalachians in order to establish their detailed physical and chemical environments of formation based on comparison with DSDP and ODP data (see also Hesse et al. 2001). The chosen formations represent basin-plain deposits (Hesse 1989) of a foreland basin (with an estimated sedimentation rate on the order of 400 &a, Hiscott et al. 1986) and slope deposits of the Ordovician southeastern continental margin of North America (Landing et al. 1992), respectively.
The formation of authigenic dolomite from organicmatter in argillaceous deep-sea sediments was first suggested by Bramlette (1946) and Spotts & Silverman (1966) for occurrences in California which were subsequently studied in great detail (e.g., Garrison et al. 1984). Since 1970, the organogenic origin of deep-sea dolomite has been confirmed with evidence from the Deep Sea Drilling Project (DSDP) and Ocean Drilling Program (ODP) (e.g., Kelts & McKenzie 1982; Lumsden 1988). Direct precipitation of dolomite from aqueous solution is a less common mode of dolomite formation, compared to the widespread replacement origin in shallow-water carbonates. Like other authigenic carbonates, dolomite is an early diagenetic pore-filling cement which may occur in the form of scattered rhombs or as concretions and diagenetic beds in argillaceous host sediments mostly in deep-water continental-margin settings. The concretions and beds turned into well cemented hard rocks, while the surrounding host sediment generally still remained in state of soft mud. Studies on land and under the sea have since illuminated many facets of deep-water organogenic dolomite formation, but did not resolve some long standing questions concerning its origin (e.g., Curtis & Coleman 1986). These include (1) the quantitative interpretation of the physical conditions of dolomite formation in terms of burial depths and temperatures based on petrographic and geochemical evidence (e.g., carbonate content, carbon and oxygen isotopes); (2)
2 RESULTS AND INTERPRETATIONS 2.1 Minus-cementporosity Based on X-ray diffraction results the Cloridorme and Levis dolomite concretions and beds consist of wellordered dolomite of near stoichiometric composition (5 1:49% = CaCO3 : MgC03 on average). Iron content averages 2% as measured by microprobe and checked by staining of thin sections. For analytical methods see Hesse et al. (2001). Calcite constitutes 0-39% of the concretion carbonate and, where present, is considered detrital in origin. Higher calcite contents (>30%) are found in 65
Cloridorme concretions which grew in carbonate-rich mud turbidites. Carbonate content of concretions is a good estimator for the porosity present at the time of precipitation, if (i) the host sediment was originally carbonate free, and (ii) no carbonate recrystallization has taken place. Detrital calcite was therefore subtracted from total carbonate content in order to obtain the authigenic carbonate as a measure of porosity at the time of concretion growth ("minuscement porosity" of Taylor, 1950). The absence of authigenic calcite is supported by thin-section evidence showing non-ferroan calcite grains which, like detrital silt-sized quartz, are on average twice as large as the 10-15 pm-diameter sucrose ferroan dolomite crystals. Preservation of the isotopic zonation of the concretions and beds speaks against recrystallization. With these precautions in mind, the variation in authigenic carbonate content from centre to margin of the concretions and beds is interpreted to indicate a porosity decrease of the host sediment during concretion growth due to compaction from 87 to 40% for the Cloridorme Formation (83-44% for the diagenetic beds) and 25 to 17% for the Levis Formation (Fig. 1). 2.2
d3cpDB
values
In the concretions and beds, 6 1 3 C p ~ values ~ decrease outwards from 10.2 to 0.8 and 7.7 to 1.2%0, respectively. This suggests precipitation in the methane-generation zone with an increasing contribution of light carbonate supplied by advection from thermocatalytic reactions at depth (Fig.2). 2.3
6 1 8 0 p values ~ ~
The corresponding 6l80 values are anomalously low (-0.4 [centre] to -6.6%0 [margin] for the concretions and -0.5 to -7.5%0for the beds, Fig.3). They can give reasonable burial temperatures only assuming an Ordovician SMOW value of -6%0 (Veizer et al. 1997) and a bottom-water temperature of 15OC. In addition, the 6l80 of the pore waters must have been lowered by mechanisms such as organic-matter decomposition
Figure 2A Variations of carbon-isotope ratios (6I3cpDB values) across and along Cloridorme concretions (H = horizontal samples from center plane of concretion). B. Traverses across diagenetic beds of Cloridorme Formation.
Figure 3A Variations of oxygen-isotope ratios values ( 6 l ' O p ~ ~ values) across and along Cloridorme concretions (H - horizontal samples as in Fig. 2A). All concretions except yl-C1 show systematic centre (sample 3 ) to margin (samples 1 and 5) variations of -3%0 requiring a 15'C temperature increase which is excessive for <25 (max. 75) m burial depth. See text for discussion. B. Same for diagenetic beds of Cloridorme Fm.
(Sass et a1.1991), volcanic-ash alteration (Perry et al. 1976), and advection of "0-depleted water. Values for the Levis Formation of -4.1%0(center) to -5.4%0 (margins) give temperatures of 32-38OC in accordance with the assumed burial depths of up to 600 mbsf (see section 3.1). The pore-water 6l80 was not lowered by volcanic ash alteration. (Notice that the Levis Fm. is located on the Ordovician non-volcanic passive margin of south eastern North America; St. Julien & Hubert 1975).
3 DISCUSSION 3.1 Burial depth of concretiongrowth. We attempted to quantify the burial depth of concretion growth by assuming that the porosity-depth trend (Fig.4) in the DSDP drill site 262, Timor Trough, is applicable to the Cloridorme Formation. This is justified by the similar depositional environment and tectonic setting of the Cloridorme Formation an the Timor Trough, a foreland basin located between the active Banda Arc and the Sahul Shelf of the northwestern Australian passive margin. Based on this comparison, the paleo-porosity (minuscement porosity) values of 80-90% for the centre of
Figure 1A. Carbonate content in vertical traverses across Cloridorme (p6, p7, y1 members) and Levis (L) concretions (sample 1 = bottom of concretion, highest sample no. = top), B. across diagenetic beds of Cloridorme Fm. (PI, p6, y2 members).
652
most concretions and at least one of the beds of the Cloridorme Formation (Fig.1) suggest that concretion growth started at very shallow sub-seafloor depth. Specifically, the maximum measured porosity is 87% - in this case growth should have started at <1 mbsf. Porosities for the margins of most concretions and some beds range from 70-80%, suggesting that termination of growth occurred at <25 mbsf, certainly at <75 mbsf. Some of the beds and one of the concretions have lower margin porosities in the 4070% range and may have continued to grow to burial depths of 150-200 mbsf. Diagenetic beds of the Cloridorme show much greater variability in minus-cement porosities than the concretions (Fig. 1). There are two explanations for this difference. If diagenetic beds, elsewhere called 'sheet concretions' (Pirrie & Marshall 1991), result from the coalescence of concretions during prolonged growth then this greater variability is to be expected because some parts of the beds become cemented much later than others. Although the samples of Figure l b are from traverses across different beds, random sampling of these beds apparently has resulted in different growth stages being represented - in some traverses the beds had completed most of their growth early, in others growth started much later. Another reason is the variation in host lithology. Beds p6-C2 and -C3 have relatively low minus-cement porosities (45-60%) throughout. Bed p6-Cl represents a dolomite-cemented, coarse silty mud turbidite, which probably started with a significantly lower initial porosity than the hemipelagic muds of the Cloridorme. Therefore, it may not be representative for the compaction history of the normal argillaceous host rocks of the concretions; the same may apply to p6-c3. Estimated minus-cement porosities for the Levis Formation concretions (Fig.1a) of 25 (centre) to 17% (margins) suggest that these concretions started to grow at depths (350 mbsf) where the growth of the Cloridorme concretions and beds had long stopped and that they continued to grow as deep as 600 mbsf. 3.2 Chemical environment of concretion growth The more or less uniform decrease of the 6I3C values of the Cloridorme concretions and beds from centre to margin by 2-5%0 provides clear evidence that dolomite precipitation took place at or after peak methane generation. Combined with the porosity data (>80%), the maximum 613C values in the centre of some concretions (10 %o) and beds ( 8 % 0 ) suggest that peak methane generation took place at very shallow subsurface depth (<5-10 mbsf in many cases). Correspondingly, the overlying sulfate reduction zone was very thin; perhaps as thin as a few centimeters (cf. Murray et al. 1978) in the case of the concretion with 87% centre-porosity. The outward decrease in the
Figure 4. Porosity-depth data for DSDP foreland-basin Site 262 (eom Rocker 1974 and Shipboard Scientific Party 1974).
6I3C values reflects the contribution from thermocatalytically generated carbonate advected from greater subsurface depth. Thermocatalytic reactions start at temperatures >7O-8O0C, which are much higher than those encountered at the burial depths of concretion growth in the Cloridorme Formation. 3.3 Subsurface precipitation
temperatures
during
dolomite
Assuming an Ordovician SMOW value of -6%0 (Veizer et al. 1997), the observed range of 6l8O values gives burial temperatures between 17'and 45-5OoC. A temperature of 17OC determined for the centre 6l80 values of some of the concretions (i.e., concretions of the p6-rriember) is in line with an assumed Ordovician bottom water temperature of 15OC. However, temperatures of 25OC for the centre and 4O-5O0C for the margins of many concretions and beds are by far too high in view of the shallow burial depths determined from the porosity data. Average paleogeothermal gradients for Ordovician sedimentary sequences of the St. Lawrence Lowlands, the passivemargin flank of the Taconian foreland basin to the south (St. Julien & Hubert 1975), are 35OCh-1 (Yang & Hesse 1993). With these gradients isotopic temperatures for the concretion margins should not exceed 18-20°C for the shallow burial samples (25-75 mbsf) and 25OC for the deeper ones (150-200 mbsf), corresponding to 6l80 values in the 0 to -2%0range. This leaves 4-6%0of the measured 6l8O values to be explained by alternative mechanisms. The most likely alternative is devitrification of volcanic glass, as the Cloridorme contains abundant bentonites (Enos 1969) and volcanic rock fragments (KO 1985). However, the 653
extent of the unexplained 6”O anomaly (4-6%0) probably requires additional sources, for which organic-matter degradation is a possibility. In the Cloridorme, this would not have occurred in the sulfate reduction zone which was too thin. However, the same mechanism continues in the thermokatalytic reaction zone, from where waters depleted not only in I3C but also in “0 have probably been advected to shallower subbottom depths. Although it is not possible to separate the effects of the different processes (volcanic glass alteration, organic-matter degradation, advection), none of them would have had a dramatic effect at very shallow subsurface depths, where the oxygen-isotopic composition of the pore water is dominated by the SMOW value. Therefore, the combined effect of the various processes cannot have lowered the 6”O of the pore waters in the upper 25 mbsf to -6%0 as required by the crystalline-phase data. This supports the assumption of a negative Ordovician SMOW value of that magnitude, although the subject is still controversial (e.g., Land 1995 vs. Veizer 1995). FEFERENCES Bramlette, M.N. 1946. The Monterey Formation of California and the origin of its siliceous rocks: U. S. Geol. Sum. Proj Paper 212: 57 pp. Curtis, C.D. & M.L. Coleman 1986. Controls on the precipitation of early diagenetic calcite, dolomite and siderite concretions in complex depositional sequences. In D.L. Gautier (ed.), Roles of organic matter in sediment diagenesis. Soc. Econ. Paleontol. Mineral., Spec. Publ. 38: 23-33. Enos, P. 1969. Cloridorme Formation, Middle Ordovician Flysch, Northern Gaspe peninsula, Quebec. Geol. Soc. Amer. Spec. Paper 1 17: 6 1 pp. Garrison, R.E., Kastner, M. & D.H. Zenger (eds.) 1984. Dolomites of the Monterey Formation and organic-rich units. Soc. Econ. Paleontol. Mineral., PaciJic Section, Spec. Publ. 41: 215 pp. Hesse, R. 1989. The Cloridorme Formation of the Taconian belt. In: Sedimentology,paleoenvironmentsand paleogeography of the Taconian to Acadian rock sequence of Gaspe Peninsula (Ed. by P.-A. Bourque, R. Hesse & B.R. Rust) Geol. Soc. Canada - Mineral. Soc. Canada., Field trip guide book B8: 185-208. Hesse, R., Shah, J. & S. Islam. 2001. Ordovician organogenic deep-sea dolomite: Physical and chemical conditions of early diagenetic concretion growth and the 6I80 of Ordovician SMOW. J. Sedim. Res. (submitted). Hiscott, R. N., Pickering, K.T. & D.R. Beeden 1986. Progressive filling of a confined Middle Ordovician foreland basin associated with the Taconic Orogeny, Quebec, Canada. In: Allen, P.A & P. Homewood (eds.), Foreland basins, Spec. Publ. Int. Ass. Sediment. 8: 309-325. Kelts, K. & J.A. McKenzie 1982. Diagenetic dolomite formation in Quaternary anoxic diatomaceous muds of Deep Sea Drilling Project Leg 64, Gulf of California. In: Init. Rep. Deep Sea Drilling Project 64, part 2: 553-569. KO, J. (1985). Controls on graywacke petrology in Middle Ordovician Cloridorme Formation: tectonic setting of source areas versus diagenesis. Unpubl. M.Sc. thesis, McGill University, 235 pp.
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Land, L.S. 1995. Comment on “Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater”. Geochim.Cosmochim.Acta59: 2843-2844. Landing, E., Benus, A.P. & P.R. Whitney 1992. Early and early Middle Ordovician continental slope deposition: Shale cycles and sandstones in the New York Promontory and Quebec Reentrant region. New York State Museum/ Geological Survey Bull. 474: 40pp. Lawrence, M.J.F. 1991. Early diagenetic dolomite concretions in the late Cretaceous Herring Formation, eastern Marlborough, New Zealand. Sediment. Geol. 75: 125-140. Lumsden, D.N. 1988. Characteristicsof deep-marine dolomite. J. sedim. Petrol. 58: 1023-1031. Mozley, P.S. & S.J. Bums 1993. Oxygen and carbon isotopic composition of marine carbonate concretions: An overview. J.sedim.Petro1. 63: 73-83. Murray, J.W., Grundmanis, V. & M. Smethic Jr. 1978. Interstitial water chemistry in the sediment of Saanich Inlet. Geochim. Cosmochim. Acta 42: 101 1- 1026. Perry, E.A., Gieskes, J.M. & J.R. Lawrence 1976. Mg, Ca and 0’8/0’6 exchange in the sediment-pore water system, Hole 149, DSDP. Geochim. Cosmochim.Acta 40: 4 13-423. Pirrie, D. & J.D. Marshal1 1991. Field relationships and stable isotope geochemistry of concretions i?om James Ross Island, Antarctica Sedim. Geol. 71: 137-150. Rocker, K. 1974. Physical properties measurements and test procedures for Leg 27. Init. Rep. Deep Sea Drilling Project 27: 433-443. Sass, E., Bein, A. & A. Almogi-Labin 1991. Oxygen-isotope composition of diagenetic calcite in organic-rich rocks: Evidence for I 8 0 depletion in marine anaerobic pore water: Geology 19: 839-842. Shipboard Scientific Party 1974. Site 262. Init. Rep. Deep Sea Drilling Project 27, 193-278. Spotts, J.H. & S.R. Silverman 1966. Organic dolomite from Point Fermin, California.Amer. Mineral. 51: 1144-1155. St.Julien, P. & C. Hubert 1975. Evolution of the Taconian orogen in the Quebec Appalachians. Amer. Jour. Science 275-A: 337-362. Taylor, J. 1950. Pore space reduction in sandstone. Amer. Assoc. Petrol. Geol. Bull. 34: 701-716. Veizer, J. 1995. Reply to the comment by L.S. Land on “Oxygen and carbon isotopic composition of Ordovician brachiopods: Implications for coeval seawater”. Geochim. Cosmochim.Acta 59: 2845-2846. Veizer, J., Bruckschen, P., Pawellek, F., Diener, A., Podlaha, O.G., Carden, G.A.F., Jasper, T., Korte, C., Strauss, H., Azmy, K. & D. Ala, 1997 Oxygen isotopic evolution of Phanerozoic seawater. Palaeogeography,Palaeo-climatology, Palaeoecology 132: 159-172. Yang, C. & R. Hesse 1993. Diagenesis and anchi-metamorphism in an overthrust belt, external domain of the Taconian Orogen, southern Canadian Appalachians - 11. Paleogeothermal gradients derived from maturation of different types of organic matter. Organ.Geochem. 20(3): 38 1-403.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Illite crystallinity an expandability: XRD and HRTEM studies of Gasp6 Peninsula mudstones and slates Salah Shatal & Reinhard Hesse Department of Earth & Planetary Sciences, McGill University, Montreal, Qukbec, H3A 2A7, Canada IPresent address: Geology Department, Faculty of Science, Suez Canal University, Ismailia, Egypt
ABSTRACT: In mixed-layer illitehmectite (VS) containing >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 <2 pm and <0.1 pm fractions before and after treatment with short- and long-chain alkylammonium cations (nc = 8 and nc = 18, respectively) with their XRD patterns as well as XRD patterns of the glycolated samples.
3.2XRD and HRTEM characterization of individual phases 3.2.1Expandable illite Long-chain alkylammonium cations (1 2 I nc I 18) are capable of selectively removing interlayer K and other cations from a wide range of micas, such as illite, glauconite, and the high-charged variety of I/S mixed-layers. The latter does not respond to ethylene glycol or glycerol and does not expand upon treatment with nc 4 2 (Vali et al. 1991). Illite and some varieties of mica possessing interlayers intercalated with nc=l8 in a paraffin-type
3.2.3Rectorite-like RI -ordered 2: 1 layer-structure The RI-ordered structure consists of an alternation of one non-expandable 2:1 silicate layer (10 A) and one 2:1 layer with an nc=18 expanded interlayer of 27 A (10 + 27 = 37 A). The diffraction pattern of this phase is characterized by a sharp peak at 36.236.9A forming part of a set of integral reflections at 36.9(OOl)*, 18.5(002)*,12.3(003)*,9.4(004)*, and 7.38(005)* A (Fig. 1). Lattice-fringe images of this phase treated with w=18show the presence of a coherent sequence of double layers with alternating non-expandable 2:1 layers and layers with nc=l8 expanded interlayers (Fig. 3). The intercalated nc=l8 interlayer suggests an intermediate to high interlayer charge-density typical of vermiculite with a paraffin-type arrangement of the alkylchains. 3.2.4Corrensite-like R1 -ordered layer structure The diffraction pattern of this phase is characterized by a sharp 43-448, peak along with a set of integral reflections. The reflection of 44.4(001)* suggest a superlattice reflection with integral higher-order reflections of 22.2(002)*,14.8(003)*, 11.1 (004)*, and 8.9(005)*A. The RI-ordered structure consists of double layers, one 2:1 layer silicate (14 A) and one 2:1 layer silicate with nc=18 expanded interlayers of 30 8, yielding a super reflection of 44 A. The intercalated 656
3.2.6 Discrete illite and chlorite The fine fractions of both high-grade diagenetic and VLGM samples comprise thin packets of discrete non-expandable 2: 1 layer-silicates along with expandable phases. In contrast, the coarse fractions exhibit thick packets of an illitic or micaceous phase either as expandable illite or non-expandable discrete illite. Similarly, chlorite forms very thick packets in the coarse fractions, consisting of thick 2: I layers of 14 A spacing. 4 DISCUSSION
Figure 1. XRD patterns of sample TM39 showing the response of different size fractions to glycolation and alkylammonium ion intercalation. Sr-Sat stands for air-dried, Sr-saturated; EG = ethylene glycol; nc = 8, octylammonium cations; nc = 18, octadecylammonium cations.
nc=18 interlayer (30 A spacing) suggests a high density of the layer charge typical of vermiculite. 3.2.5 "Kalkberg" R3-ordered layer structure The diffraction patterns of this phase is characterized by a sharp reflection of 51.6 A forming part of a set of integral reflections at 51.6 (001)*, 25.8 (002)*, 17.2 (003)" and 10.5 8, (005). An R3-ordered 2: 1 layer-structure consists of three 10 8, 2:l layers (30 A) followed by one 2:l layer with an nc=l8-expanded interlayer of -22 A.
Figure 2. XRD patterns of sample TM29, showing the response of different size fractions to glycolation and alkylammonium ion intercalation (for abbreviations see Figure 1).
During burial diagenesis the percentage of expandable layers in mixed-layer clays decreases systematically as a fimction of many factors, among which temperature with increasing burial depth is most important (Srodon 1984). Below 10-20% smectite (S), expandable components become difficult to detect with conventional methods. For samples formed under high-temperature conditions (>165OC), EG solvation causes only a slight sharpening (or none at all) of the reflections (Hunziker et al. 1986). The conventional methods to assess expandability, such as the peak position of glycolated samples (%SxRD, Srodon 1981) or the intensity ratio (Srodon 1984), are no longer suitable under anchizonal conditions. In ion-thinned samples studied under the HRTEM, expandable components cannot be detected because they collapse under the electron beam in the high vacuum. It is for these reasons that the role of expandable components in high-grade diagenetic and VLGM rocks has remained obscure and poorly documented. The present study for the first time demonstrates the presence of expandable components in high-grade diagenetic and anchizonal samples. However, such components are considered metastable and not expected to last indefinitely during prograde recrystallization. With prograde changes in the VLGM zone, expandability is expected to reach a minimum, beyond which stacking faults shall be the main cause of the remaining lattice strain. Consequently, the continuous sharpening of the 10 A X-ray reflection in a prograde alteration sequence in anchizonal rocks is interpreted as gradual conversion of the last remaining expandable phases to illite, to the point at which stacking faults remain the sole cause of imperfect crystallinity. The occurrence of different types of expandable components and consequently variations in interlayer charge density can be detected by XRD and HRTEM in n-alkylammonium-treated samples, whereas ethylene glycol treatment of XRD mounts is indifferent to these variations. The arrangement of alkylammonium ions exchanged for inorganic interlayer-cations in smectite, vermiculite, illite and 657
treated clays. This method offers a unique opportunity to identify the presence of expandable components in high-grade diagenetic and VLGM illitic clays and to distinguish different types of illites that cannot be achieved with XRD upon ethylene glycol treatment. The illite phases identified with this technique in a prograde sequence of hgh-grade diagenetic to VLGM rocks are rectorite-like R1-ordered layer structure either with an intermediate or high-charge vermiculite interface, and R3-ordered structure found in the fine fraction while expandable illite is found mainly in the coarse fraction. Expandable 2: 1 clay minerals decrease in abundance with increasing metamorphic grade. Stacking faults may be the main source of lattice strain only when phyllosilicates consist of thick non-expandable layer silicates. Chlorite evolution in response to high-grade diagenesis and VLGM appears to follow reaction pathways parallel with US, but with fewer amounts and types of intermediate products. The chlorite mixed layers consist of a corrensite-like R1-ordered layer structure.
Figure 3. HRTEM image of the <2 pm size fraction of sample TM39 after treatment with nc=18 showing packets of expandable illite interlayers (solid diamonds) occurring together with thin packets of three to five non-expandable 2: 1 layers of illite (open diamond) and a sequence of double layers with expandable interfaces (hollow triangle) interpreted as "rectorite-like RI -ordered structure".
REFERENCES
Figure 4. HRTEM images of the <2 pm size-fraction of the sample TM 29 after treatment with nc = 18: coherent packets of expanding 2.1 layers with average spacing of 25 A (solid diamonds) alternating with very thin packets of layers of nonexpandable illite (hollow diamond). The solid arrow indicates the non-expandable packet termination.
mixed-layer minerals and the degree of expansion are a function of alkyl chain length (6 1.nc I 18) and the density of interlayer charge.
5 CONCLUSIONS The presence of expandable layers in illite is a key problem for the evaluation of illite crystallinity (IC) at high (>8O%) illite levels. Although high-grade diagenetic and VLGM samples do not expand in response to ethylene glycol treatment, the IC improves significantly with increasing diagenetic and metamorphic grade as a consequence of the loss of still present expandable components. The contribution of expandable components to IC in the high-grade diagenetic to VLGM zone can be detected with the aid of XRD patterns and HRTEM lattice-fringe images of n-alkylammonium-cation
Arkai, P. & M. Toth 1983. Illite crystallinity: Combined effects of domain size and lattice distortion. Acta Geologica Hungarica, 26(3-4): 341-358. Hesse, R. & E. Dalton 1991. Diagenetic and low-grade metamorphic terranes of Gaspe Peninsula related to geologic structure of the Taconian and Acadian orogenic belts, Quebec Appalachians. J Metamorphic Geol. 9: 775-790. Hunziker, J.C., Frey, M., Clauer, N., Dallmeyer, R.D., Friedrichsen, H., Flehmig, W. & K. Hochstrasser 1986. The evolution of illite to muscovite: mineralogical and isotopic data from the Glarus Alps, Switzerland. Contrib. Mineral. Petrol. 92: 157-180. Merriman, R.J., Robert, B. & D.R. Peacor 1990. A transmission electron microscope study of white mica crystallite size distribution in a mudstone to slate transitional sequence, NorthWales, UK. Contrib. Mineral. Petrol. 106: 27-40. Nieto, F. & A. Sanchez-Navas 1994. A comparative XRD and TEM study of the physical meaning of the white mica "crystallinity" index. Eur. J. Mineral. 6: 6 1 1-62 1. Shata, S. 2000. Illitization and chloritization of illitelsmectite and chlorite/smectite mixed-layer clays in high-grade diagenetic and very low-grade metamorphic environments of Gaspe Peninsula, Quebec Appalachians, Canada: Problems and Solutions. Ph.D. Thesis, McGill Univ., 18Op. Srodon, J. 198 1. X-ray diffraction of randomly interstratified illite-smectite in mixtures with discrete illite. Clay Miner. 16: 297-304. Srodon, J. 1984. X-ray diffraction of illitic materials. Clays Clay Miner. 32: 337-349. St. Julien, P. and C. Hubert 1975. Evolution of the Taconian Orogen in the Quebec Appalachians. Atner. Jour. Sci., 27514: 337-362. Vali, H., Hesse, R. & E.E. Kohler 1991. Combined freezeetched replicas and HRTEM images as tools to study fundamental particles and multi-phase nature of 2: 1 layer silicates. Amer. Mineral. 76: 1973-1984. Warr, L.N. & F. Nieto 1998. Crystallite thickness and defect density of phyllosilicates in low- temperature metamorphic pelites: a TEM and XRD study of clay mineral crystallinity-index standards. Can. Mineral. 36: 1393-14 13.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
H,S in North Sea oil fields: importance of thermochemical sulphate reduction in clastic reservoirs R .H .Worden Department of Earth Sciences, University of Liverpool, Brownlow Street, Liverpool, M9 3GP, UIK.
P.C .Smalley BP, Sunbury Technology Center, Chertsey Road, Sunbury Middlesex, TW16 7LN, U.K.
ABSTRACT: H2S in North Sea basins is broadly limited to the Central and Wytch Ground Grabens where it occurs at concentrations of up to 20,000 ppm. H2S is likely to be due to thermochemical sulphate reduction occurring between sulphate enriched waters, derived from Zechstein evaporites, and oil. Although the sulphur content of oil decreases with increasing reservoir temperature, H2S and the oil-sulphur content are not inversely correlated: H2S cannot have originated from S-compound cracking in oil. There is a broad correlation between the sulphur isotope value of oil and H2S concentration suggesting that TSR-H2S has been incorporated into the oil. H2S is only notably enriched in oil fields at about 120°C proving that thermochemical sulphate reduction is likely (and bacterial reduction unlikely). H2S is highly enriched only in Upper Jurassic, clean quartz arenites suggesting that in more normal subarkosic and sublithic sandstones, TSR-H2S is scrubbed by Fe-minerals producing late-stage pyrite cement, common in North Sea reservoirs. H2S and that reservoir lithology and distribution of subsurface evaporites seem to play important roles.
1 INTRODUCTION It is widely assumed that clastic reservoirs and clastic-dominated basins contain little, or no, H2S in oil and gas fields because it is recognised that clastic rocks have an abundance of Fe-bearing minerals that can naturally scrub the H2S leading to the growth of pyrite. This assumptions seems reasonable as long as: (1) H2S is removed faster than it is added to a reservoir and (2) the sandstones in question actually have Fe-minerals. In fact, many sandstone reservoirs have very small amounts (1 or 2 ppm; Whittingham & Jones 1987) of H2S in associated gas that may be intrinsic to the system or, more commonly, that is the result of production activities leading to bacterial sulphate reduction. There are significant amounts of H2S in some North Sea reservoirs. Values of H2S in associated gas in North Sea oil fields range naturally from below detection (< 0 ppm) to 20,000 ppm. There are four natural potential sources of HIS: direct generation from sulphur-rich kerogen, break down of S-rich oil, bacterial sulphate reduction (BSR) and thermochemical sulphate reduction (TSR) (Orr 1977). The first two result from Fe-poor source rocks, the second two require a supply of aqueous sulphate. In this paper, we will examine H2S distribution in North Sea oil fields and link H2S data to oil inspection properties, and sulphur and carbon isotope geochemistry and conclude that TSR has been the dominant mechanism that produced
2 DATA SOURCES AND TYPES Data on H2S and oil sulphur contents have been collated from company reports, database searches and from the open literature. The geochemical data have been compiled over many years and represent analysis of over 500 oil samples from exploration, appraisal and production wells. Note that most fields have no H2S reported suggesting that it is below detection limits. Such data are not added to any of the plots. It should be appreciated that the H2S data reported represent only those fields for which there was a definite and measured H2S content in the associated (or free gas) phase.
3 THE NORTH SEA BASINS Most North Sea oil fields sit within the Central, Viking, and Wytch Ground Grabens which form a rified triple junction. The Wytch Ground Graben extends west towards the UK coast from the main rift axes. The dominant source rock is the Kimmeridge Clay Formation which is a type II/III source rock (predominantly clastic marine source with minor coal components). Typical maturation products are dominated by sweet (low-S) petroleum 659
that can vary from a wet gas through condensates to single liquid oils. The reservoirs are mainly clastic and range from Triassic redbeds to Middle Jurassic deltaic and shallow marine sandstone and submarine fans of the Upper Jurassic and Tertiary. The Upper Jurassic sandstones tend to be the cleanest sandstones approaching a quartz arenite classification. The Permian salt of the Zechstein rich in anhydrite and extends over most of the southern North Sea including the Central Graben but has a northern subcrop that largely coincides with the Wytch Ground Graben.
especially at northern subcrop limit of the Zechstein evaporites. Southern North Sea gas fields have no reported H2S. There is very little intrinsic H2S in northern North Sea reservoirs. Although very low H2S concentrations (1 -2 ppm) have been reported for the Northern North Sea oil fields, these are likely to be due to bacterial contamination, sea water injection elevating the formation water sulphate concentration and cooling the near to the well bore to within a range tolerated by active bacteria. 5 LINK OF H2S TO PETROLEUM GEOCHEMISTRY
4 TEMPERATURE-EFFECTS It is widely expected that the high pressure/high temperature (HPHT) reservoirs, presently being investigated in the North Sea, will be where H2S will be a potential problem. Analysis of the data do not support this notion. The maximum H2S concentrations tend to be found in (more or less normally pressured) reservoirs that are at or about 120°C (Figure 1). The highest pressure (and most overpressured) fields and highest temperature fields (up to 19OOC) do not contain exceptional H2S contents. Although the temperature for the maximum H2S is not exactly the same as that found in some gas-dominated basins in the Middle East (Worden et al. 1995; 1998). The temperature for the elevated H2S fields precludes bacterial sulphate reduction since bacteria struggle to survive (let alone be effective) at such high temperatures. The temperature seems to put the H2S-generating process firmly within the TSR temperature range. The highest H2S concentration are found in fields in the northern Central and Wytch Ground Grabens,
North Sea oils are mostly sweet crudes although the bulk sulphur content can be as high as 2 wt %. There is a good correlation between the H2S concentration of the associated gas and the sulphur content of the oil (Figure 2). A simple interpretation of this is that high sulphur content source rocks have locally led to the generation of high sulphur crudes and abundant H2S. It is also possible that high sulphur oils have decomposed during reservoir heating resulting in H2S. However, the picture is not quite so simple. Sulphur and nitrogen contents of North Sea oils broadly correlate although there are notable deviations to high sulphur at relatively low nitrogen contents (Figure 3). Both the sulphur and nitrogen contents of oil tend to decrease during oil heating and cracking leading to a good correlation of the two (for a given source rock). It is thought that deviation of the correlation towards relatively high sulphur at low nitrogen contents is the result of thermochemical sulphate reduction with backincorporation of some of the H2S into the oil (Orr 1974; Thompson 1994). Thus heating of oil in the
Figure 1. Effect of temperature on H2S abundance in N Sea oil fields (fields with no detectable H2S not represented)
Figure 2. Correlation of bulk oil sulphur content and the H2S content of the associated gas.
660
TSR has been shown typically to result in fractionation of the carbon isotopes of the reacting petroleum (Krouse et al. 1988; Worden & Smalley 1996; Worden et al. 2000) whereby the I2C isotope reacts faster than the I3C isotope. This is a kinetic effect resulting in progressively elevated 6I3C values in the remaining petroleum and CO2 resulting from the oxidation of petroleum with very low 613C values. Whole oil 634S data correlate well with oil fraction 613C data suggesting that the processes that lead to the elevated sulphur isotope ratio also lead to elevated carbon isotope ratios (Figure 5). This seems to confirm the occurrence of oxidation of the oil, presumably by TSR. The modifications of the carbon and sulphur isotopes in the oil are likely to occur at different stages of the oil and sulphur compound reactions. The oil will react initially with sulphate, generating H2S, and resulting in carbon isotope fractionation but with no initial effect on sulphur isotopes. Reaction of the TSR-H2S with oil may happen next and would be the stage at which the sulphur isotope signature of the oil became modified. The following reactions (RI and R2) and the sulphur isotope values reflect the interpreted sequence of reactions and their effect of on Sisotope ratios:
Figure 3. Bulk oil sulphur and nitrogen contents form N Sea oils. N and S decrease with oil-cracking. TSR tends to cause sulphurisation and deviation from the maturation trend.
presence of sulphate can lead to competing desulphurisation (cracking) and sulphurisation (reaction of TSR-H2S with the oil). Nitrogen is lost by cracking regardless of any other processes. If H2S was sourced from sulphur compounds in oil, then the two might be expected to be inversely related (which is not the case). North sea oils with elevated sulphur (and thus with elevated H2S in associated gas) tend to have elevated 634Svalues (Figure 4). Sulphur in oil is thought to have a very similar stable isotope ratio to the sulphur in the parent kerogen. Similarly, organicderived H2S is thought to have a very similar stable isotope to parent S-rich kerogen and oil. Marine kerogen contains sulphur that has a 634Sabout 15%0 lower than the contemporary sea water. Thus Upper Jurassic kerogen (and resulting S in oil and organicderived H2S) should have a 634S of about O%O (Claypool et al. 1980). The higher sulphur contents in some oils thus seems to be unlikely to be due to a normal marine kerogen given that the 634S values are up to 8%0 enriched relative to the anticipated value. Note that low S-oils have a 634Svalue typical of normal marine kerogen sourced oils. TSR seems not to result in sulphur isotope fractionation even though sulphate and sulphur have a significant fractionation factor. The feedstock sulphate 634S is typically adopted by the resulting H2S. Orr (1974) concluded that, during TSR, oil tends to adopt the sulphate 634S due to sulphurisation of oil by back reaction with the TSR-H2S. Thus the elevated 634S of the oil is likely to be due to TSR between oil and sulphate and then reaction of the H2S with the remaining oil.
compound
intial oil
6"s CDT
0%0
compound 6"s CDT
I
+ aqueous
altered O%o Oil
so4
=
altered
+
H2S
+
CO2
oil
(RI) 15%0
+
15%0
O%o
H2S
15%0
=
S-enriched altered oil >O%o
(R2) but <15%0
Figure 4. Correlation of sulphur content of oil and its sulphur isotope ratio
261
8 CONCLUSIONS
1. There has been widespread thennochemical sulphate reduction in the Central and Wytch Ground Grabens of the North Sea. 2. TSR seems to have occurred predominantly at 120°C. 3 . The sulphate was sourced from the Zechstein evaporites. 4. H2S has accumulated only in clean Upper Jurassic sandstones, largely free of Fe-minerals. 5. Oil has reacted with some of the H2S and has partly adopted the 634Sratio of Zechstein evaporites. REFERENCES
Figure 5. Correlation of carbon and sulphur isotope ratios from N Sea oils.
6 SOURCE OF SULPHATE FOR TSR The sulphate that caused the TSR is likely to have been derived from the Permian Zechstein evaporites. There is no other major reservoir of mineral sulphate in the North Sea basins and many studies of sulphur isotopes from reservoir sandstones have concluded that the Zechstein has been the dominant source of diagenetically late anhydrite cement (e.g., Gluyas et al. 1997). The distribution of fields with elevated H2S seems to define the subcrop of the Zechstein with little or no H2S found in fields north of the Zechstein subcrop and much less south of the subcrop limit. This may be a simple coincidence or may suggest that more sulphate can be mobilised (by water) from the edge of the (impermeable) Zechstein than from underneath and through the Zechstein. 7 RESERVOIR EFFECTS High H2S concentrations are only found in very clean sandstones of the Upper Jurassic play fairways. Dirtier (more Fe-rich) sandstones of for example the Brent Group allow easy pyritisation of Fe-minerals. It should be noted that diagenetically late stage pyrite is common in many reservoirs possibly suggesting that TSR-derived H2S is ubiquitous but most sandstones have the capacity to removed by natural scrubbing. The sulphur isotope composition of late stage pyrite (even in the Wytch Ground Graben reservoirs) is typically complex although the most enriched values reach (and exceed) those of the Zechstein sulphate (e.g., Riciputi and Hendry 1997).
Claypool, G.E., W.T. Holser, I.R. Kaplan, H. Sakai, I. Sak 1980. The age curves of sulfur and oxygen isotopes in marine sulfate and their mutual interpretation. Chem. Geol. 29: 491-501. Gluyas, J., L. Jolley and T.J. Primmer 1997. Element mobility during diagenesis: sulphate cementation of Rotliegend sandstones, Southern North Sea. Mar. Petrol. Geol. 14: 1001-1011. Krouse, H. R., Vian, C. A., Eliuk, L. S., Ueda, A., and Halas, S. 1988. Chemical and isotopic evidence of thermochemical sulfate reduction by light hydrocarbon gases in deep carbonate reservoirs. Nature 333: 415-419. Orr, W. L. 1974. Changes in sulfur content and sulfur isotope ratios during petroleum maturation - study of Big Horn basin Paleozoic oils. Amer. Assoc. Petrol. Geol. Bull., 50: 2295-23 18. Orr, W. L., 1977. Geologic and geochemical controls on the distribution of hydrogen sulfide in natural gas. In: R. Campos. and J. Goni, Advances in organic geochemistry, 1975, Madrid, Spain: 571-597. Riciputi, L. and J.P. Hendry 1997. Cross formational flow during burial of Cretaceous sandstones (North Sea): implications from high resolution 634S analysis of mesogenetic Fe-sulphide. In: (Hendry J., Carey P., Parnell, J.,Ruffell, A. and Worden, R.H. eds.) Extended Abstract volume of the Geofluids 11 ‘97 conference: 331-334. Geological Society, London. Thompson, K.F.M., 1994. A classification of petroleum on the basis of the ratio of sulfur to nitrogen. Org. Geochem. 2 1: 877-890. Whittingham, K P and T.J. Jones 1987. Reservoir souring. Royal Soc Chem. 67: 228-243. Worden, R.H. and P.C. Smalley 1996. H2S-producing reactions deep carbonate gas reservoirs: Khuff Formation Abu Dhab Chem. Geol. 133: 157-171. Worden R.H., P.C. Smalley and M.M. Cross 2000. The influences of rock fabric and mineralogy upon thermochemical sulfate reduction: Khuff Formation, Abu Dhabi. J. Sed. Res. 70: 1218-1229. Worden, R.H., P.C. Smalley and N.H. Oxtoby 1995. Gas souring by thermochemical sulphate reduction at 140OC. Assoc. Petrol. Geol. Bull. 79: 854-863. Worden, R.H., Smalley, P.C. and Oxtoby, N.H. 1998. Gas souring by thermochemical sulphate reduction at 140°C: reply. Amer. Assoc. Petrol. Geol. Bull. 82: 1874-1875.
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Magmatic, metamorphic and minerogenetic processes
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Metasomatic reaction bands - a key to component mobility at metamorphic conditions R.Abart Imtitute of Mineralogy and Petrology, Karl-Franzens- University, Graz, Austria
ABSTRACT: In the Ivrea Zone @-Italy) metasomatic reaction bands comprising a monomineralic clinopyroxene layer, a grandite-garnet clinopyroxene layer, and a scapolite-clinopyroxene layer developed at hornblendite-marble interfaces during granulite facies metamorphism (T=75OoC, P=700 to 1000 MPa). Formation of the reaction bands was controlled by diffusive redistribution of silica, aluminum, calcium, and magnesium between the two mutually incompatible rocks. The phenomenological Onsager diffusion Coefficients are constrained to: Lsisi/Lcaca>2.5, LAIAI/LCaCa1. The reaction bands accommodate an oxygen isotope gradient from about 18%o(SMOW) in the marble down to 7%0 in the hornblendite. Minimum estimates for the phenomenological oxygen diffusivities are on the order of 10-l2to 10-13m2/s within the individual layers of the reaction bands. This is several orders of magnitude faster than oxygen volume diffusion in the corresponding minerals suggesting that oxygen diffusion was enhanced by coupling of oxygen transport with major component diffusion and/or by grain boundary diffusion. 1 GEOLOGIC SETTING The Ivrea zone is an approximately northeastsouthwest striking unit primarily comprised of metasediments and metabasites. It was metamorphosed at upper amphibolite facies to granulite facies conditions at about 270 Ma (Vavra et al. 1999). In Val Fiorina some 12 kilometers southeast of Domodossola, a granulite facies metamorphic sequence of the basal Ivrea zone is exposed in the riverbed. Among others, the sequence comprises marbles, which behaved very ductile during granulite facies metamorphism. In the course of synmetamorphic deformation, fragments of hornblendite country rock were incorporated into the marble. The hornblendite xenoliths and the marble were incompatible at high grade conditions and complex multilayer reaction bands formed at their interfaces.
2 PETROGRAPHY OF THE REACTION BANDS The reaction bands comprise unaltered hornblendite, a clinopyroxene layer, a garnet-clinopyroxene layer, a scapolite clinopyroxene layer, and calcite marble. A schematic sketch of the reaction band is given in Figure 1. The original hornblendite is comprised of
pargasitic hornblende, almandine-pyrope-grossular garnet and tschermakitic clinopyroxene. The hornblendite is replaced by a monomineralic clinopyroxene layer. The clinopyroxene layer is, in turn, replaced by a grandite-garnet clinopyroxene layer. The outermost portion of the reaction band is represented by a scapolite clinopyroxene layer. The inner portion of this layer exhibits a symplectite texture, the outer portion has a polygonal equigranular texture. The occurrence of two texturally distinct domains in the scapolite clinopyroxene layer suggests, that the original hornblendite-marble interface was located within this layer.
3 MASS BALANCE CONSIDERATIONS The major element compositions of the reactant rocks and of the individual mineral layers are given in Table 1. There is a monotonic trend of increasing calcium content from the hornblendite to the calcite marble and a reverse trend for iron and magnesium. The silica content reaches a maximum and the aluminum content shows a minimum in the clinopyroxene layer. Neither of the compositions of the mineral layers nor any positive linear combination of these 665
Figure 1. Schematic illustration of the reaction band, reaction bands are up to 12 centimeters wide.
compositions can be produced from the reactant rock compositions. This indicates that reaction band formation involved material transfer across the phenomenological limits of the reaction band. At the high grade conditions of reaction band formation neither of the components can a priori be assumed as immobile, nor can the volume change be quantified. An unambiguous derivation of material fluxes across the boundaries of the reaction band is thus impossible. The stoichiometry of the bulk metasomatic reaction can, however, be constrained from qualitative arguments. From mass balance considerations a relation between the bulk reaction stoichiometry and the volume factor may be derived for the conditions of conservation of component i:
amount of component i contained in phase j . In Figure 2 the conditions for the individual component conservations are shown. Any point to the right of a component i conservation curve represents a scenario with net introduction of this component into the reaction band, any point to the left represents a net loss of this component. For the geological situation at hand, an outward diffusion of MgO and a net inward diffusion of CaO are the most likely scenario. Feasible mass balance scenarios are thus those to the right of the CaO- and to the left of the MgO conservation curve. This limits the molar proportion of reactant marble over hornblendite to less than 0.4.
4 RELATIVE COMPONENT MOBILITIES -
m:v,
'hornblendire
-
ntmarble
f v vmarhie
where and Vhornb[end,reare the stoichiometric coefficients of the reactant marble and homblendite and fvis the volume factor. m: is the molar amount of phasej in the mineral layer k,
i, is the
In a system with more than two components, the sequence of mineral layers in a metasomatic reaction band depends on the relative component mobilities (e.g. Joesten 1977). The observed reaction band may volume factor 0.6 0.8 1.0
molar
0.1
Table 1. Selected major element concentrations.
Wt%
horncpx blendite layer
41.45 Si02 13.85 A1203 F e 2 0 3 ~ ~ 13.93 ~ 11.17 MgO 13.78 CaO 1.63 Na2O L.O.I. 0.84
47.30 8.18 10.68 8.90 20.76 0.91 0.80
grtcpx layer 41.64 13.48 9.71 4.15 27.48 0.09 0.52
1.2
0.0
volume of phase or rock j and n;' is the molar
scp- marble cpx layer 38.05 9.22 17.45 3.43 5.75 3.00 2.23 1.26 32.45 48.64 0.79 0.39 2.46 32.75
0.2 0.3
0.4
0.5 Figure 2. Mass balance scenarios for reaction band formation, curves indicate component conservation.
666
be described in a simplified four component system with components Si02, A1203, MgO, and CaO. Based on the assumption that transport occurred by diffusion and that local equilibrium prevailed, constraints on the relations between the phenomenological Onsager diffusion coefficients of the individual components may be derived form “steady state diffusion modeling” (Joesten 1977). Conservation of component i in the metasomatic reaction at layer boundary 1 may be expressed by:
Figure 3. Results of steady state diffusion modeling, valid and preferred solutions indicated by light and heavy dots, insert: bulk reaction stoichiometry.
where srct is a source/sink term that accounts for the liberatiodconsumption of component i in the course of the metasomatic reaction at layer boundary 1, v: is the stoichiometric coefficient of phase j in the metasomatic reaction at layer boundary 1, and nl is the molar amount of component i contained in phase j . The continuity of diffusive fluxes is expressed by:
ratios are arbitrarily specified. The strategy is now to systematically vary the diffusivity ratios and classify the results with respect to the stoichiometric coefficients obtained. A solution is regarded as a valid solution, if the stoichiometric coefficients are compatible with growth of all mineral layers. A solution is classified as a preferred solution, if the stoichiometric Coefficients also approximately reproduce the relative layer thickness and the modal mineral contents. The result of a model calculation for fixed LSiSi/LCaCa is shown in Figure3. Calculations were done for the range of feasible mass balance scenarios and for a variety of LSiSi/LCaCa ratios. The influence of the bulk metasomatic reaction stoichiometry, i.e. the boundary fluxes is relatively modest and the general pattern of valid and preferred solutions remains the same. Valid solutions are only found for values of LSiSi/LCaCa>2-5,LAIAl/LCaCal. This indicates that diffusion of silica was at least by a factor of 2.5 faster then CaO diffusion, MgO diffused at least as fast as CaO and A1203 may have diffused faster than CaO by a factor of 10 or slower.
JF - J,”-’= srct , where J,“ is the diffusive flux of component i in mineral layer k. The condition of local equilibrium requires that the Gibbs-Duhem equation is fulfilled at any point:
where dp, is the total differential of the chemical potential of component i and the sum is over all system components. The relation between the gradient in chemical potential of component i and its diffusive flux is given by:
=-L-----!dP Idx’ where L, is the phenomenological Onsager Diffusion coefficient. From these relations a total of 37 linear equations in 41 unknowns is derived for the observed layer sequence. The unknowns are the stoichiometric coefficients of the reactant rocks and of the mineral phases that take part in the metasomatic reactions at layer boundaries, the source/sink terms for liberatiodconsumption of material at layer boundaries, the diffusive fluxes of components within the individual mineral layers, and the relative component diffusivities expressed as diffusivity ratios: LSiSi/LCaCa, LAlA&aCa, LMgMg/LCaCa. The system of equations may be solved, if the stoichiometric coefficient of reactant hornblendite is set to -1, and the three diffusivity J
5 OXYGEN ISOTOPE SYSTEMATICS The unaltered hornblendite has an oxygen isotope composition of about 7%o(SMOW)the marble has a composition of 6180(SMOW) = l8%0. The oxygen isotope transition across the reaction band is illustrated in Figure 4. The observed pattern may be interpreted in terms of a moving boundary diffusion problem. The individual mineral layers represent media with distinct transport properties and the replacement fronts represent the moving boundaries. The front geometry reflects the relative oxygen bulk diffusivities in the different media. The transport 667
6 CONCLUSIONS Formation of granulite facies multilayer reaction bands at homblendite-marble interfaces was controlled by diffusive redistribution of Si02, A1203, MgO and CaO, where Si02 and MgO diffused fast as compared to CaO, A1203 may have diffused up to ten times as fast as CaO or slower. Oxygen diffusion across the reaction band was significantly enhanced as compared to volume diffusion by either fast diffusion along grain boundaries or due to coupling with major element transport. Figure 4. Oxygen isotope systematics of the reaction bands.
parameters derived from the front geometry are given in Table 2. The effective bulk oxygen diffusivities vary over at least one order oaf magnitude within the reaction band. The bulk oxygen diffusivity was slowest in the monomineralic clinopyroxene layer and it was fastest in the marble. This indicates, that oxygen diffusion was controlled by the mineral content of the respective medium. The maximum duration of oxygen isotope exchange between the hornblendite and the marble matrix is 2.7*107 years, that is the age of granulite facies metamorphism. For a duration on the order of 107 years effective oxygen isotope diffusivities of 10-13 m2/s are derived for the monomineralic clinopyroxene layer. This is six orders of magnitude faster than experimentally determined rates of oxygen volume diffusion in clinopyroxene. The discrepancy between the estimated oxygen bulk diffusivity and rates of oxygen volume diffusion suggests that oxygen diffusion was significantly enhanced in the reaction bands. This is ascribed to either one or a combination of two processes: Oxygen diffusion may have been significantly enhanced by diffusion along grain boundaries (e.g. Joesten 1991), and, in addition, it may have been coupled to major element diffusion that was driven by gradients in the chemical potentials of the major system components. Table 2. Effective bulk oxygen diffusivities in the different domains of the reaction bands.
Do*t
Do DO 103years 1 hornblendite 690 2.2* 1O S 2.2* 10-" cpx-layer 40 1.3* 1OW9 1.3* 10-' grt-cpx layer 370 1.1*10-' l.l*lO-'o scp-cpx layer 500 1.6*10-s 1.6*10-" marble 2200 7*10-' 7*10-"
years
'
DO 107years 2.2" 1O-'* 1.3* 10-13 1.1*10-'2 1.6*10-12 7*10-12 668
ACKNOWLEDGEMENTS This study was hnded by the Austrian Science Foundation, grant Nr. P12903-GEO. R. Schmid, D. Harlov and R. Sperb are thanked for their contributions to this work.
REFERENCES Joesten, R. 1977. Evolution of mineral assemblage zoning in diffusions metasomatism. Geochim. Cosmochim. Acta 4 1 : 649-670. Joesten, R. 199 1. Grain-boundary diffusion kinetics in siicate and oxide minerals. In: Diffusion, atomic ordering and mass transport, J. Ganguly Ed. 345-395. Vavra, G., Schmid, R. & D. Gebauer 1999. Internal morpholoy, habit and U-Th-Pb microanalysis of amphibolite to granulite facies zircons: geochronology of the Ivrea zone (Southern Alps). Contrib. Mineral. Petrol. 134: 380-404.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Petrology and alteration of basalts from the intraplate rises, Indian Ocean A .V.Artamonov, V.B Kumosov & B .P.Zolotarev Geological Institute, Moscow, Russia
ABSTRACT: Basalts recovered during ODP Leg 115 on Mascarene Plateau, Chagos Bank, and Maldives Ridge are all tholeiitic in composition. Abundance trace and rare earth elements (REE) suggest the presence of normal and enriched varieties (N-MORB and E-MORB) among the rocks from studied holes. Basalts from all holes of Leg 1 15 are slightly altered. Secondary minerals are represented mainly by smectites and calcite, Feoxides and hydroxides. This secondary mineral assemblage is significantly poorer than that in basalts from Ninetyeast Ridge. 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. The main peculiarity is in that oxidative alteration is accompanied mainly by accumulation of major and minor elements while the non-oxidative alteration is accompanied mainly by loss of elements. The various ages of sub-sea volcanic edifices in Indian Ocean did not change significantly the degree of basalt alteration.
1 INTRODUCTION The detailed investigations of intraplate rises (aseismic ridges, plateaus, and seaniounts) is important for understanding of formation and geodynamics of ocean floor. Two groups of large rises are located in the different segments of Indian Ocean: Kerguelen Plateau, Broken Ridge and Ninetyeast Ridge - in eastern segment; Mascarene Plateau, Chagos Bank, Maldive Ridge and Laccadive Ridge - in western segment. We have studied samples of basalts recovered during Leg 26 DSDP and Leg 121 ODP on Ninetyeast Ridge (Artamonov et al. 1998), and Leg 11 5 ODP on Mascarene Plateau (Holes 706C and 707C), Chagos Bank (Hole 713A), and Maldives Ridge (Hole 715A). Hole 706C is located on the northeastern margin of Nazareth Bank. Hole 707C is located in the northwestern part of the Mascarene Plateau. Hole 713A represents the northern part of the Chagos Bank. Hole 715A is located on the northeastern margin of the Maldives Ridge (Beckman et al. 1988, Fig. 1). Volcanic activity occurred at these sites at 34, 64, 49. and 57 Ma, respectively (Duncan & Hargraves 1990). The Ninetyeast Ridge and Mascarene-Chagos-MaldivesLaccadive Volcanic lineament are an aSeiSmiC ridge systems in the Indian Ocean basin.
Figure 1. Location map of the western Indian Ocean showing ODP sites (Leg 115). Radiometric ages of basal& (Ma) are given in parentheses.
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2 PETROLOGY Basalts recovered from Holes 706C, 707C, 713A and 71 5A are all tholeiitic in composition. Chemical composition of basalts of upper and lower lava units from Hole 706C are different (Baxter 1990). Trends of the REE distribution for aphyric basalts of upper lava units are similar to Deccan Traps (Lightfoot & Hawkesworh 1988) and basalts of Reunion Island (Fisk et al. 1988), but La,,/Yb,, ratio in island’s basalts and traps are higher. Concentrations of the rare earth and trace elements and trends of the REE distribution in plagioclase-phyric basalts of lower lava units from Hole 706C and olivine-phyric and aphyric basalts from Hole 715A are similar to E-MORB. Plagioclase-phyric basalts from Holes 707C and 713A are similar to N-MORB. The geochemical data suggest that basalts of studied holes are derived from mantle melts generated at different depth. 3 ALTERATION Alteration of tholeiitic basalts from MascareneChagos-Maldives (Leg 1 15) and Ninetyeast (Legs 26 and 121) lineaments demonstrate both similarity and some differences. Maximum degree of alteration of basalts from Leg 115 is significantly lower (H20c up to 1.46, 1.69, 2.12, and 2.89 wt%, Holes 707C, 71 3A, 706C, and 715A, respectively) than in basalts from Ninetyeast Ridge (H20’ up to 1.4, 3.6, 4, and 6.2 wt%. Holes 756D, 758A4,254, and 757C, respectively). Moreover, each hole drilled in Leg 115 contains basalts as fresh or almost unaltered rocks with H20’ variation from 0.15 to 0.79 wt%. Low degree of basalt alteration in all holes of Leg 115 is resembled by alteration of basalts from Holes 756D and 758A (Ninetyeast Ridge, Leg 121). Slight alteration of rocks from Holes 706C, 707C, and 713A is probably due to the following main reasons: low temperature of alteration and relatively low permeability for the seawater of sampled sections of basalts. The temperature of calcite formation in basalts (Hole 707C) was about 12 to 2OoC (Burns et al. 1990). Basalts from these holes show low permeability. They are mainly massive and sparsely vesicular, especially deep-water lava flows from Hole 713A. This bacame the reason to create low water-rock ratio. The latter explain both slight alteration of basalts from Holes 706C, 707C, and 713A and the absence of oxidizing environment. Similar environments controlled alteration of the dominant portion of basalts from Hole 758A (Ninetyeast Ridge, Leg 121). All basalts from Holes 706C, 707C, and 713A are dark gray to black in color. Besides, they demonstrate Fe203/FeO ratio from 0.4 up to 1.25. This is higher than in typical non-oxidized basalts. This, in turn, reflects the presence of some slight alteration in the past. Also,
“non-oxidized” basalts from Hole 758A demonstrate signs of slight oxidation. FelO3/FeO ratio varies from 1 up to 1.6. This is slightly higher than in basalts from Holes 706C, 707C, and 713A. Two types of low-grade “non-oxidative” (Fe203/FeO < 1.60) and “oxidative” (FezOj/FeO > 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).
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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
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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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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 ~ <+6.5%0 ) can not have developed these values by differentiation from any normal magma. The low 6 l 8 0 ( w ~rocks ) do not show distinctive differences in majodtrace element geochemistry compared with normal-6"O equivalents and only poorly correlate with proximity to DZb shear zones (in as much as low 6 l 8 0 ( w ~ ) values only occur at crustal levels that contain migmatitic D2b shear zones). The first observation suggests that the highly variable but low 6 l 8 0 ( w ~ ) values do not result from variable contamination of parental magmas by 10w-'~O crustal material. Rather, they probably reflect heterogeneous fluidrock interaction with heated low-6l 8 0 fluid. Such a fluid may have been derived directly from the surface or from dehydration of l 0 w - 6 ~ ~ rocks 0 in the terrain. The latter possibility is unlikely given Figure 2. Whole rock 6l80 values (vs SMOW) of metagranites that metagranites and mafic dykes in the deepest and mafic dykes, inside and outside of DZbshear zones. Data crustal levels of the Mann Terrane have 6l80 values for continental intraplate basalts and MORE3 from Hoefs that are little reset from igneous precursors, and that (1 997). these rocks, and Musgrave Inlier granulites, contain of likely precursors, and require either derivation little or no water. from abnormally low- " 0 magmas, contamination of Limited mineral oxygen isotope data show: 1) high-temperature fractionations between quartz and these magmas by 10w-'~Ocrust or interaction with 1 0 w - 6 ~ ~fluids 0 at some time. The lowest 6180(wn) minerals such as hornblende or clinozoisite, which formed through partial melting reactions in the shear values occur within the D2b shear zones. However, the following observations suggest that fluid flow zones; and 2) negative Ai80~quartz-k-fe~dspar) values that during shearing has not lowered 6'%(WR) values 1) indicate isotopic disequilibrium, even within the metagranites in any one shear zone show a migmatitic D2b shear zones. The former, together considerable variation in 6l 'O(WR) values (for with a lack of evidence of low-temperature alteration example, -4%0to +3.5%0, with one sample at -16.1%0 a n y h e r e in the terrain, suggests that the low within the Mt. Le Hunte Shear Zone); 2) there are 6' O(WR)values were set prior to, or during, the variable but small increases (0-2%0) in 6 I 8 0 ( w ~ ) Petermann Orogeny. The latter may reflect the difficulty of isotopically equilibrating the large (2-4 values from D2b shear zones into country rocks on a tens of metre scale; 3) 10w-'~Ometagranites in the cm) IS-feldspar megacrysts during fluid infiltration. 679
One migmatitic granite has an exceptionally low ~ " O ( W Rvalue ) of -16.1%0. This value has been verified in duplicate at Monash, and by analysis of mineral separates ( ~ " O ( Q 7 ~ )-15.4 %o; 6180(czo)=: 17.1 %o) at Cape Town. It is unclear why only one sample has such an extremely low 6 l 8 0 ( w ~value. ) However, such a low 6 1 8 0 j ~value ~ ) requires infiltration of extremely low- '0 fluid over any reasonable temperature range (6180(H20) = c.-16 to c.~ O ) 300-600°C; Clayton 21%0 based on A ~ ~ O ( Q ~ - Hover et al. 1972). Such extreme estimates require that either the Mann Terrane was at a high palaeolatitude, or that mountainous topography existed, at the time that infiltration took place. In the Arunta lnlier (Fig. l), Cartwright & Buick (1998) described shear zones developed during the 300-400 Ma Alice Springs Orogeny that contain similarly 1 0 w - 6 ~ ~ 0but , lower grade, sheared metagranites. In these greenschist-grade shear zones lowering of 6"O values was accompanied by metasomatism, and adjacent granitic country rocks had normal igneous 6 " 0 ( w ~ )values. Both isotopic resetting and metasomatism were inferred to have resulted from long-distance, polythermal flow of surface-derived fluids (Cartwright & Buick 1998). However, in the present case, l 0 w - 6 ~ ~granites 0 also occur well away from shear zones, suggesting that there is not a simple causative link between isotopic resetting, fluid flow, deformation and anatexis. One possible explanation for this lack of a simple relationship is that the igneous rocks of the Mann Terrane underwent hydrothermal alteration, and variable isotopic resetting, at shallow cmstal levels, prior to the Petermann Orogeny. Hydrothermal alteration may have been most intense along normal fault systems. These hydrothermally altered igneous rocks were then metamorphosed during the intraplate Petermann Orogeny, with reactivation of normal faults as D2b thrusts, and localisation of melting in these hydrous, and therefore more fertile deformation zones (M. Hand, pers. comm., 2000). Such a model explains the apparent decoupling of isotopic resetting, deformation and partial melting. The occurrence of low-6"O amphibolites in the shear zones is consistent with hydrothermal alteration occurring after dyke emplacement i.e. as late as 0.8 Ga, but before 0.56 Ga. Hydrothermal alteration could have been driven by emplacement of the dyke swarms. Alternatively, isotopic alteration of the granites may have pre-dated dyke emplacement, with partial, metre-scale isotopic equilibration occurring between recrystallised dykes and granites during Petermann-age D2b shearing. The heterogeneous patterns of isotopic resetting in the Mann Terrane are not consistent with the introduction of a low-''O fluid to great crustal depths to facilitate high-pressure metamorphism. Instead, the current dataset suggests that early
hydrothermal alteration may have exherted an important control on terrain reactivation, and partial melting, during intraplate orogenesis. ACKNOWLEDGEMENTS Research finding from the ARC (Large Grant A00000278 and Senior Research Fellowship to ISB, Large Grant A39030662 to IC), the NRF (CH) and a University of Cape Town Postdoctoral Fellowship (JM) are gratefully acknowledged. Marlen Yanni and Martin Hand are thanked for help with isotope analyses, and constructive discussions, respectively.
REFERENCES Bamicoat, A.C. & I. Cartwright 1997. The gabbroeclogite transformation: an oxygen isotope and petrographic study of west Alpine ophiolites.Journal of Metamorphic Geology 15: 93-104. Cartwright, I. & 1.S Buick 1998. The flow of surface-derived fluids through Alice Springs age middle-crustal ductile shear zones, Reynolds Range, central Australia Journal of Metamorphic Geology 17: 397-4 14. Clayton, R.N., O'Neil, J.R. & Mayeda, T.K., 1972. Oxygen isotope exchange between quartz and water. Journal of Geophysical Research 77: 3057-3067. Hoefs, J. 1997. Stable Isotope Geochemistry: 201. Berlin: Springer-Verlag. Johannes, W. & F. Holtz 1996. Petrogenesis and experimental petrology of granitic rocks: 334. Berlin: Springer-Verlag,. Rubie, D.C. 1990. Role of kinetics in the formation and preservation of eclogites. In D.A. Carswell (ed.), Eclogite facies rocks: 1 1 1-140. New York: Chapman & Hall. Scrimgeour, I. & D.Close 1999. Regional high-pressure metamorphism during intracratonic deformation: the Petermann Orogeny, central Australia. Journal of Metamorphic Geology 17: 557-572. Sun, S.S., Sheraton, J. W., Glikson, A.Y. & A.J. Stewart 1996. A major magmatic event during 1050-1080 Ma in central Australia, and an emplacement age for the Giles Complex. Australian Geological Survey Organisation Research Newsletter 24: 13-15. Taylor, H.P. & S.M.F. Sheppard, 1986. Igneous rocks: I Processes of isotopic fractionation and isotope systematics. In: Valley, J.W., Taylor, H.P. & O'Neil, J.R. (eds.). Stable Isotopes in High Temperature Geological Processes. Reviews in Mineralogy 16: 227-272. White, R.W., Clarke, G.L. & D.R. Nelson 1999. SHRIMP UPb dating of Grenville-age events in the western part of the Musgrave Block, central Australia. Journal of Metamorphic Geology 17: 465-48
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Natural zeolites from Cenozoic pyroclastic flows of Sardinia (Italy): evidence of different minerogenetic processes P.Cappelletti, G .Cerri & M.de’Gennaro Dip. di Scieiize della Terra, Universitd di Napoli Federico II, Nupoli, Ituly
A .Langella Facolta di Scienze, Universitd del Sannio, Benevento, Italy
S.Naitza, G.Padalino & R.Rizzo Dip. di Geoingegneria e Tecizologie Ainbieiitali, Universita di Cagliari
M .Palomba Centro Studi Geoininerari e Mineralurgici del CNR, Universitd di Cagliari
ABSTRACT Different ininerogenetic processes leading to the crystallization of clinoptilolite and inordenite have been recognized in the rhyolitic to dacitic volcaniclastic deposits of Sardinia For Loyudoro region zeolitization is due to late circulation of hydrothermal fluids favored by the fault system of the area Bosa region was affected by a sin- and post-depositional zeolitization due to the phreatoinagmatic character of the volcanic event Finally, for Central Sardinia the clear mineral zonation suggests a late zeolitization process in an open hydrologic system
1 INTRODUCTION
Recent studies carried out on the Clenozoic volcanosediinentaiy sequences of Sardinia recognized largescale zeolitization processes as a consequence of water/rock interaction phenomena, leading to cryst a11izat i o n of cli no pt i 1o 1it e, soin et i ines associated to mordenite These deposits, unique as diffusion on the national territory, constitute a source of industrial minerals of potential economic interest The areas so far identified (Fig 1) are localized in CentralNorthern (I,oXiiu’oro. Cerri et al 200 l ) > CentralEastern (Ho.w, Cappelletti et al. 2000a) and Central ( A / / L I I - . ~ . sCappelletti ~ ~ I ~ I , et al 2000b) sectors of Sardinia, where zeolitization affected both the daciticrhyolitic volcaniclastites and the i-elated epiclastic products The present research compares the data so far collected in order to identify the possible zeolitization mechanisms, at present defined only for the L o ~ u doro sector (Clei-ri et al 200 1 ) 2 GEOLOGY OF THE STIJDIED AREAS 2. I I.ogdoro i q y o t i , Northcii.ii LSciidiiiiu
1,ogudoi-o region is located at the eastern edge of the “Fossa Sar-da”, formed by Burdigalian NNW tensive basins; supeninposed on older ENE trending Oligo.4quitanian strike-slip basins The volcano-sedimentary products filling these basins belong to a calk-alkaline to Khigh calcalkaline succession Porphync andesite domes, over-lainby pyroclastic ash and pumice flows with a different welding degree, fonn the oldest volcanic sequence. Re681
cent surveys allowed to distinguish a basal rhyodacitic pyroclastic flow characterized by a deeply welded middle level. overlain by an epiclastic unit showing fluvio-lacustiine sedimentation features The succession is closed by a 1-hyoliticpyroclastic flow Epiclastic unit and unwelded levels of the two flows are often deeply zeolitized (Cerri et al 200 1 )
3.2Ho.w .sc>c’tol; N o i . t h ~ t i ~ ~ , ~Stiidiriiu t~~i.ii I n the Bosa sector, the Great Ignimbrite of Bosa (GIB) is a rhyodacitic compound large volume igniiiibrite i-elated to phreatomaginatic volcanic activity (Assorgia et a1 1094) emplaced at 18 Ma. The GIB, up to 150 in thick and extending over =30 km‘, consists of three overlapped poorly welded, ash-and-pumice pyr-oclastic flows resulting froin r-epeated highly explosive events and forming a single cooling unit. The tvpical volcanological layering of the igniinbrites, as defined by Sparks et al ( 1 973) was identified in eacli of the three flow units In particular, ground surge and ash cloud surge levels were interpreted as sti-atigraphical markers (Cappelletti et al 2000a) The upper flow shows increasing welding towards the top 2.3
7 c t l t l l ! / sLrldllllLl
I n this sector (East of the “Fossa Sarda”), a coinplex v~~lcan~~-sedii~ientary succession, LIP t o 300 in i n
thickness, crops out (Assorgia et a1 1998) The lowermost volcanic formations, the Luzzana (LU) and Allai (.4L) ignimbrites, are poorly welded, ash-andpumice rhyodacitic flows, deriving from phreatomagmatic explosive volcanic activity, and form a sequence up to 100 in thick These volcanic episodes are followed by the sedimentary epiclastic Terrigenous-TufTaceous Complex (TT), and by the Dacitic lava Complex (DL), consisting of dacite lava domes and lava flows The overlying Ruinas Unit (RU) is a poorly welded to welded dacitic ignimbrite, up to 40 in thick, and is covered by the Asuni Unit (AS), about 50 in thick, consisting of a dacitic poorly welded ash-and-pumice flow, followed by an upper epiclastic arenaceous sequence formed by erosion of previous volcaniclastic formations The latest pyroclastic products in the area form the Monte Ironi LJnit (MI), a welded to incipiently welded dacitic ignimbrite 10-20 in thick 3 MATERIALS AND METHODS Volcaniclastic units were sampled in detail in all the studied areas In each section, all layers were distinguished and sampled, aiming at investigating the possible relationships between the volcanological layering and the qualitative and/or quantitative differences in the mineral assemblages Whole-rock mineralogy was determined by X-ray powder ddfraction (XRPD), using a Philips PW 1730/3710 instrument Scanning electron microswpy was carried out with a IEOL JSM 53 10 A LINK AN 10000 microanalyzer system was used for EDS analyses Quantitative mineralogical analyses were pertbnned using the Reference Intensity Ratio (RIR) technique 4 RESIJLTS AND DISCUSSION
4.I /,O~~lldOI.O l.C~).yO1I I n Logudoro region geological-stiuctural, chemical and ininero-petrographical data indicate that the zeolitization of two pyroclastic flows and the interlayered epiclastic unit is due to a circulation of hydrothermal fluids mainly controlled by the fault system of the area Evidences of this process are rep[resented by a general silica enrichment, derived either from fluids OJ from glass dissolution, within the zeolitized sectors which led to a supersaturation and consequent precipitation of opal-CT, the late epigenetic phase in the glass + smectite -+clinoptilolite 3 opal-CT sequence Within the welded layers, the emplacement features and the thermodynamic conditions favored the crystallization of an adularia-like. feldspar, and a lower silica concentration in the circulating fluids allowed quartz to precipitate mainly in the fracture system In addition, this described proc682
ess is supported by a preferential occurrence of quartz, adularia, sometimes mordenite, as well as calcedony veins close to the fault system Notwithstanding this minerogenetic mechanism settled a well defined lateral variability all over the investigated area (about 140 kin2), some statistical considerations can be drawn On average, the epiclastite shows the highest and less variable zeolite grades if compared to the other units (Fig. 2a). This is likely due to the presence of a more abundant, fine-grained and homogenously distributed glassy matrix that increases the specific surface of the system favouring the water-rock interaction process and, consequently., the zeolitization process Finally, the space-time variatior! of chemistry and temperature of the circulating fluids determined a wide variability of the exchangeable cationic population of clinoptilolites that was also recorded laterally within each single unit (Fig. 3) On the other side, the influence of the composition of the starting material cannot be disregarded, in fact, the ciinoptilolites of the two pyroclastic flows are generally Krich, whereas those from epiclastic units are enriched in divalent cations deriving by the partial alteration of andesites and metamorphic rocks which constitute the clastic component of these sediments, along with elements of the lower ignimbrite 4.2 Ho.ar .sec1or’
As well as Logudoro region, a detailed research was carried out on the Great Ignimbrite of Bosa A minerogenetic hypothesis was fiirmulated considering al! the geological> volcanological and mineralogical features This unit shows high homogeneity either in terms of minei-alogical or chemical composition of the authigenic phases Within the fifty collected samples only few showed minor amount of mordenite, being all the others charactei-ized by clinoptilolite as prevailing authigenic phase along with minor amounts of opal and smectite The Figure 2b shows the variations of quantitative contents of clinoptilolite, opal-CT and smectite as a function of the stratigraphic position of the corr-esponding sample It is evident that the welded layers. such as the top of the sequence, are totally sterile as already reported in other volcanic sequences ( e y Logudoi-o). The remaining part of the formation shows a quite constant clinoptilolite content, on average higher than SO%, even if the occurrence of high grades of zeolite is concentrated in the lowerground surge layer, characterized by a finer and glass-rich grain size fraction Chemical analyses of clinoptilolites also evidenced a marked compositional homogeneity, which retlects the chemistry of the original rhyolitic-rhyodacitic glass, as shown in the diagram of Figure 3, where cationic compositions gather in a quite narrow area shifted towards the K vertex.
F i g u r e 2 R e c o n s t r u c t e d s t r a t i g r a p h l c s e q u e n c e of t h e s t u d i e d s e c t o r s - a ) L o g u d o i - o \!I=Upper ignimbrite EI’I=Epiclastite, l,l==Louer igiiimbrite b ) B o s a and c ) A l l a i facies subdi\rision according to Sparks e t a1 ( 19731, d ) . A s u n i .+I--L o w e r ‘ I a m i n a t e d l a e~r . R = I / l a s s i \ e b o d y , C = - U p p e r p u m i c c - c o n c e n t ! - a t e d 1 a ) ’ c r
683
The following Na-depletion in the system shifted the equilibrium towards zeolites characterized by a ditrerent cationic population (Ca, K, Mg) with similar or slightly lower Si/AI ratios such as clinoptilolite, which become:; the only occurring zeolite in the lowermost layers Transitional conditions, with coexistence of the two zeolites, occur in the middle portion of the deposits
5 CONCLUSIONS The data so far collected for Sardinia region turned to be quite interesting, not only for the relevant occurrence on its territory of clinoptilolite-rich volcaniclastic deposits but also for the new minerogenetic knowledge it provides I n fact, the scenario which is taking shape refers to a mostly heterogeneous system of water-rock interaction processes that led to the zeolitization of these volcaniclastic and epiclastic pr-oducts All the studied areas were characterized by almost ccxval active volcanisrn but, in each sector a ditlf'erent ininerogenetic path was followed. all leadiny to similar results.
Figure 3 Mole plot of exchangeable cation content in clinopti lolites from the studied areas /ill these data, along with the volcanological evidences which account for a phreatomaginatic activity. agi-ee for a sin- and post-depositional zeolitization process favored by the massive presence of wztei- during the eruption The considerable thicki~essof the deposit assured a partial thermal insulation of the system necessaiy for the enhancement of the zeolitization, a process similar to that proposed by de'Gennaro et al (2000) for- pyroclastic tlows of C'ampi Flegrei 4.-7
REFERENCES Assorgia A . Chaii L.S.. Delno A . Garbariiio C . Montaiiari A . Rizzo R. & S. Tocco 1904. Prcliiiiiiiar!. volcaiiogeiiic and paleoiiiagnctic studics oii the Cciiozoic calc-alldic eniptive scquciize of Monte FLII-TU (Bosn. mid-n cstcrii Sardinia) I n Coccioiii R . Mont;iiiari A 22 Odin G . cds . Miocciic Stratigraph> of Ital! and Adjacent Regions. (;iormlc cli (;colo,y/u. spccinl i s s ~ i ~ . 5(>. 17-2!, Assorgia. A . Barca. S . Balogh. I< . Porcu. A . Spaiio. C' & R RIZZO1')OX The Ollgocciic-MloccjicSCdliiiciit;li-! niid volcanic siicccssioiis of Central Sardinia. Ital!. /(om .I , ~ l r ~ i l i ~ ~78. r ~ 9-23 ~~~~i.~i. Cappellctti P . lbba A . Langclla A . Naitza S . & R Rizzo 2OOOn Geological and iiiiiieralogical studies oii the Grcat Igiiiiiibritc of Bosn First evaluation of resources for tcclinological application. In Proc V Con\vgno Nazioiinlc di Scieilza c Tcciioiogia dcllc Zcoliti Ra\ cllo. Oct 2000. 7 7 9-2 3 2 Cappcllctti. P . dc-Gciinaro. M . Lniigclln. A . Nnitzfi. S Padalino. G . Palomba M 22 R Rizzo 2OOOb .ki\:aiiccs i i i mineral exploration for zcolitcs i n Central Sardinia. :lie scctor of Asuiii (Central Sai-dinia. Ital!,) Proc V Conwgiio Nazioiialc di Sciciizn c Tcciiologia dcllc Zcoliti K:i\;cllo. Oct 2000. 233-236 Ccrn G . C'appcllctti P . Laiigclln A CY: M dc'Cicimnro 200 i Zcolitizatioii of Oligo-Miocciic volcaniclastic rocks fmii Logtidoro (northern Sardinia. Ital!.) 'ot7/r/h A / f i m r l i / / ) c / r f ) l 140 (4) 1-15 dc'Gennaro M.. Cappcllctti P . Langclla A . Pcrrotta .4 22 C Scarpati 2000 Gcncsis of zcolites i n the Ncapolitnii Yullo\\ TLIE.geological. \.olcaiiological and iiiincralogicril cvldcIiccs ( 'ot~tt./hM I I ~ C I/)c/rol ~ ~ I I I30 ( 1 ) 17-35 Sparks R S J . Self. S . K: G P L Walker 1073 Products of lgiiimbritc Eniptions. (;colo,yii I 1 15-1 1 S.
~ ~ ~ ~ l l, iS iO~l ~c~iI ll l l O
Several ignimbi-iteunits with interbedded epiclastic episodes, form the volcanic succession outcropping in Central Sardinia A lower rhyolitic-rhyodacitic ignimbrite sequence (Luzzana and Allai Units) and an upper dacitic igninibrite (Ruinas and M te lroni Units) are interlayered by two main epiclastic episodes The sequences considered in this study particularly refer to the volcanic subunits of the Asuni and Allai Units Both deposits tui-ned t o be strongly afected by minerogenetic post-depositional processes that led to the crystallization of clinoptilolite and inordenite along with analciine and smectite Figures 2c, d compare the two stratigi-aphic successions and the 1-elated content of authigenic phases The most striking feature is the presence of a similar pattern of mineral zonation In particular, moi-denite occurs in the upper layers and gradually disappear at the bottom of the deposit whereas clinoptilolite tollows an opposite trend being totally absent at the top ofthe subunit with the highest grades at the bottom Moreover, at least in the Asuni sector, the intermediate layers are characterized by the concomitant occurrence of both zeolites These mineralogical evidences allow to propose an open hydrologic system as the most suitable yenetic inodel .4 supposed starting Na-rich solution interactin2 with a rhyolitic-rhyodacitic glass could have settled conditions favorable to inordenite crystallization 684
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Amphibole evolution in uitramafic amphibolites from NE Sardinia, Italy A .M .Caredda, G .Cruciani & M .Franceschelli Dipartimento di Scienee della Terra, via Trentino 51, 0912 7, Cagliari, Italy
G . Carcangiu Centro Studi Geominerari e Mineralurgici C N R . , Piazza d’Armi, 09126, Cagliari, Italy
ABSTRACT: Ultramafic amphibolites from NE Sardinia consist of relics of igneous phases and metamorphic minerals formed during the granulite, amphibolite and greenschist facies. Igneous olivine and plagioclase reacted to produce coronas consisting of orthopyroxene, clinopyroxene, green spinel and garnet. The amphibolite stage is dominated by formation of amphiboles, clinochlore, plagioclase and Cr-bearing spinel and talc. Greenschist minerals consist of amphiboles, fayalite, epidote, albite, calcite, dolomite and serpentine. Four types of clinoamphibole have been recognised: brown amphiboles (Caml : magnesiohornblende to pargasite), green amphiboles (Cam2: tschermakite to magnesiohornblende), zoned magnesiohornblende (Cam3), and late stage tremolite (Cam4). Two textural types of orthorhombic amphiboles have been detected. The P-T path of ultramafic amphibolites is characterised by a decrease in both pressure and temperature fiom the granulite to the greenschist stage. The compositional variation and cation distribution of the amphiboles can be related to the P-T evolution of the ultramafic amphibolites during the regional Hercynian metamorphism. (011) and plagioclase (P11); (mineral abbreviations as in Kretz 1983 and Spear 1993). Olivine grains are surrounded by an annular shell of Opx2 and a discontinuous irregular patch of Cpx2. Igneous plagioclase is rimmed by abundant small grains of green spinel (Spll) forming an irregular shell followed by a layer of garnet (Grtl). The amphibole shows a variety of morphological and textural features. Four types of clinoamphiboles have been distinguished in these rocks: brown amphibole (Caml), green amphibole (Cam2), green zoned amphibole (Cam3) and late stage Cam4 amphibole. Caml grows on igneous orthopyroxene and clinopyroxene and very rarely forms large grains enclosing olivine and plagioclase. Cam2 encloses large coronas around coarse-grained igneous plagioclase (P11). Cam2 grows on pyroxene and/or as a thin layer between Grtl and Cpx2/Opx2 of the coronas. Cam2 also rims the Caml amphibole with a sharp boundary between the two amphiboles. In samples from Layer C, Cam2 occurs as medium-sized crystals associated with very coarsegrained garnet (Grt2). Cam2 and porphyroblastic garnet (Grt2) form worm-like microstructures with a Cr-bearing spinel (Sp12). Cam3 amphibole occurs in veins or in the chlorite-rich amphibolites as zoned medium grained crystals. The late stage Cam4 amphibole replaces pyroxene, garnet and Cam2 amphibole, and/or forms a thin rim around Cam2 amphibole.
1 INTRODUCTION The migmatite of the Hercynian basement of NE Sardinia contains metabasites with relics of granulite facies assemblage. At Montiggiu Nieddu, there are two types of metabasites: ultramafic amphibolites and plagioclase-banded amphibolites (Ghezzo et al. 1979). The ultramafic amphibolites form a darkgreen to black, massive to weakly schistose body, made up of three main compositional layers (A, B, and C). Layer A (-20 m thick) consists of grey to brownish moderately schistose chlorite-rich amphibolites; Layer B (- 5 m thick) of plagioclase-rich greenish amphibolites; Layer C (up to 10 m) of garnet-rich dark-green amphibolites. Microstructures and mineral compositions indicate that the ultramafic amphibolites were granulites, which underwent various re-equilibrations during the amphibolite and greenschist facies. The aim of this communication is to shed some light on the amphibole evolution in the ultramafic amphibolites during the metamorphism that affected the Hercynian chain of Sardinia. 2 MINERALOGY AND PETROGRAPHY The ultramafic amphibolites are made up of relics of igneous phases and metamorphic minerals in variable amounts. Primary igneous minerals are: orthopyroxene (Opx l), clinopyroxene (Cpx l), olivine 685
Two textural types of orthoamphiboles have been detected: Oaml is present as small grains at the contact between 011 and the surrounding shell of Opx2. Oam2 occurs as small grains or interleaved with chlorite (Chll) and is usually associated with small crystals of Cam2 and orthopyroxene. Other observed minerals are: albite, blende, calcite, chlorite (Ch12), corundum, dolomite, epidote, fayalite, pyrite, serpentine, talc, titanite, and Fe-Ti oxide.
Table 1. Selected microprobe analyses of clinoamphiboles and anthophyllite from Montiggiu Nieddu. Type
Caml
Cam2 Cam3c Cam3r
Cam4
Ath
SiOz TiOz
3 MINERAL CHEMISTRY
Cr203 FeOtot MnO MgO CaO NazO K20
46.15 0.80 12.15 0.15 7.40 0.09 16.30 12.40 2.05 0.12
43.81 0.12 15.50 0.00 8.90 0.00 14.55 12.53 2.40 0.10
52.50 0.14 5.75 0.01 7.25 0.16 18.83 12.17 0.70 0.09
47.80 0.30 10.90 0.00 8.90 0.18 16.84 12.31 1.50 0.08
57.05 0.01 0.38 0.00 2.60 0.15 22.40 13.20 0.01 0.00
57.50 0.00 1.85 0.00 10.91 0.06 24.99 0.17 0.0 1 0.22
3.1 Amphibole
Total
97.61
97.91
97.60
98.81
95.80
95.7 1
Selected Caml, Cam2, Cam3, Cam4, and anthophyllite analyses are listed in Table 1. Caml, Cam2, Cam3, and Cam4 have been calculated to 13 cations except Ca+Na+K, while anthophyllite has been calculated to 23 oxygens. The composition of brown amphiboles (Caml) ranges from magnesiohornblende to pargasite (nomenclature after Leake et al. 1997 They are characterised by a large variation in Al' (1.15- 1.70) and AIV' (0.60-0.80) contents, a high XMg (0.60-0.70), and N a ~ 4content of 0.10-0.45. The sum of site A is up to 0.60. The composition of the Cam2 amphiboles widely ranges from tschermakite, magnesiohornblende to pargasite. The are characterised by high A1" (1.101.90) and A1"7 (0.40-1.20) contents, lower Ti (0.010.09), and N a ~ 4content of 0.07-0.45. XMg ranges from 0.60 to 0.80. Variations in composition in the Cam2 amphibole grains have been detected within a thin section during the microprobe analyses. Cam3 is a zoned magnesiohornblende. From core to rim Ca is almost constant, while Si and Mg decrease; Ti, AIrV,AI"', and NaA increase (Table 1). cam4 is a tremolite with very low NaM4 content (up to 0.10). Oaml and Oam2 are anthophyllites with Xrvlg = 0.70-0.80.
Si AI" AIV' Cr Fe3+ Ti M FegMn Ca Na~4 NaA
6.54 1.46 0.57 0.02 0.36 0.08 3.44 0.52 0.01 1.88 0.12 0.45 0.02
6.24 1.76 0.84 0.00 0.39 0.01 3.09 0.67 0.00 1.91 0.09 0.58 0.02
7.29 0.71 0.23 0.00 0.63 0.01 3.90 0.21 0.02 1.81 0.19 0.00 0.02
6.64 1.36 0.42 0.00 0.79 0.03 3.49 0.25 0.02 1.83 0.17 0.24 0.01
7.96 0.04 0.02 0.00 0.07 0.00 4.66 0.23 0.02 1.97 0.00 0.00 0.00
8.0 1 0.00 0.30 0.00 0.00 0.00 5.19 1.27 0.0 1 0.02 0.00 0.00 0.04
15.47
15.60
15.02
15.25
14.97
14.84
'41203
t.
K Total
a gradual decrease in XMg,and a progressive decrease in Ca content. Spll and Sp12 are members of the spine1 s.s-hercynite series (Hc=45-60%). Chll is a clinochlore and Ch12 a Fe-rich chlorite. 4 DISCUSSION AND CONCLUSION 4.1 Amphibole-forming reactions The plot of A1" versus the (Na + K) atoms (Fig. 1) in calcic amphiboles from Montiggiu Nieddu shows a linear trend of isomorphic substitution between tremolite and an intermediate member between pargasite and tschermakite. The tschermakite and pargasite substitutions are approximately in the ratio 1:l. Caml and Cam2 overgrow Cpxl or Opxl, and develop between the garnet and pyroxene layers of the coronas via the following schematic reaction:
3.2 Other minerals The main minerals coexisting with the amphiboles have the following composition: 011 olivine: Fo-70%; P11 plagioclase: An>80%, P12 plagioclase: An=36-84%; Cpxl and Cpx2: Di=80-88%, Hd=2012%; Opxl and Opx2: En-70%. Grtl is an almandine-rich garnet (49-62 mol%) with high grossularite (16-25 mol%), pyrope (18-27 mol%) and minor spessartine (1-3 mol%) content. Grt2 garnet is an almandine-rich garnet with a higher pyrope and a lower grossularite content compared to Grtl. Grt2 is zoned, and characterised towards the rim by an increase in Mg and Fe content,
(1) Grtl,2 + Cpxl,2 + Opx1,2 + V = Caml,2 + Sp12
In the garnet-free rocks the close association of Cam2 amphibole and chlorite (Chll) also suggests the following schematic Cam2 amphibole-forming reaction: (2) 011 + Cpxl,2 t Opx1,2 + Spl+ V = Cam2 + Chll
686
Figure 1. Plot of the AI'" vs. (Na+K) atoms for the calcium-rich amphiboles from Montiggiu Nieddu.
Caml and Cam2 amphiboles show a spread in AI" and (Na-tK) content, but Caml amphibole is characterised by a narrower range in the (Na+K) content than Cam2. When Cam2 rims the Caml, the transition from brown to green amphibole is always sharp, and several compositional features indicate control by the microtextural site on Caml composition. In fact the lower Al,,, content, similar XMgand NaM4amount of Caml, compared to Cam2, is presumably due to Caml growth in microdomains at the expense of igneous pyroxene. Cam 2 amphibole shows extensive solid solution and its composition is controlled by the host rock chemistry and P-T metamorphic conditions. However, the spread in composition of the Cam2 amphibole may be due to the continuous compositional reequilibration, and/or to the progress of reaction (1) and (2) during the retrograde P-T evolution of the rocks. In fact, in the most re-equilibrated samples, i.e. in samples with rare relicts of igneous and/or granulitic phase, the Cam2 amphibole is poorer in total alkali content and richer in tschermakite andor tremolite substitutions. This type of compositional variation is observed among the various grains within the same samples. Continuous chemical zoning in the (Na+K), and AI" content of Cam3 suggest that it has crystallised during a single continuous stage of metamorphism. Semi- uantitatively the zonation of Cam3 amphibole in Al", NaA, and A1"' documents a prograde metamorphic evolution in the amphibolite facies. The destabilitization of the granulitic assemblage is also documented by the growth of Oaml amphibole surrounding shell of Opx2 or by very rare formation of Oam2 amphibole interleaved with Chll chlorite during the amphibolite stage.
Figure 2. Variation of Na/(Na+Ca) vs. AV(Al+Si) content of calcic amphiboles from Montiggiu Nieddu. Envelopes delimit the composition for medium-pressure and low-pressure amphiboles from Dalradian (Scotland), Haast River (New Zealand) and Abukuma (Japan) respectively (Laird & Albee 1981). Symbols as in Figure 1.
A new formation of amphibole is supported by the growth of Cam4 coexisting as patches or thin rims on Cam2 through the following reaction: (3) Cam2 + H 2 0 = Ep + Ch12 + Cam4 (Tr) + Qtz
The very low NaM4,(Na+K), and A1 IV of Cam4 suggest that it developed in greenschist facies P-T conditions. In the Na/(Na+Ca) vs. Al/(Al+Si) diagram of Figure 2, all the calcic amphiboles plot within the mediumpressure field of Dalradian (Scotland) amphiboles. 4.2 P-T conditions The P-T conditions estimated for the granulite stage are in the range 680-750°C and 8-10 kbar (Caredda et al. 1999). No amphibole growth has been observed in the granulite stage. The conditions estimated for the amphibolite stage by the garnet-hornblende thermometers (Graham & Powell 1984) and the garnet-amphiboleplagioclase-quartz barometers (Kohn & Spear 1989) agree with microtextural observations, i.e. lower pressure (4-6 kbar) and temperature (580-640°C) compared to the granulite stage. Coexistence of Ch12 chlorite, tremolite and epidote suggests reequilibration during the greenschist stage at temperatures ranging from -300 to 400°C. The empirical barometer based on the NaM4content in the amphibole (Brown 1977) on Cam4 yields pressures below 2-3 kbar. The P-T path of ultramafic amphibolites is charac687
terised by a decrease in both pressure and temperature from the granulite to the greenschist stage. This type of P-T evolution is also supported by the composition and cation distribution in the amphiboles formed during the various metamorphic stages. The zoning observed in the Cam3 amphibole may be tentatively related to a fluctuation in the metamorphic temperature during the final exhumation stage of the rocks, which is probably due to the intrusion of late tectonic Hercynian granitoids. The P-T path of the ultramafic amphibolites of Montiggiu Nieddu took place over a considerable range of P-T conditions, and is characterised by a prevalent decompression regime. The P-T path is typical for continental collision where rocks were exhumed both due to the tectonic movement and erosion (England & Thompson 1984).
Spear, F.S. 1993. Metamorphic phase equilibria andpressuretemperature-time paths. Washington DC: Mineralogical Of America.
ACKNOWLEDGEMENTS Financial support was provided by the Italian MURST project: " Terrestrial materials and synthetic analogues at high pressure and high temperatur: physical,chemical, and rheologic properties". Cofin 98 (National Coordinator C.A. Ricci; Local Coordinator M. Franceschelli).
REFERENCES Brown, E.H. 1977. The crossite content of Ca-amphibole as a guide to pressure of metamorphism. J. Petrol. 18: 1853-72. Caredda, A.M., Franceschelli, M., Memmi, I. & M. Zucca 1999. Mineralogical re- equilibration and P-T path in amphibolitic rocks from the Hercynian metamorphic basement of NE Sardinia, Italy. Plinius, 22: 9 1-92. England, P.C. & A.B.Thompson 1984. Pressure -temperature time path of regional metamorphism I. J. Petrol. 25: 894920. Ghezzo, C., Memmi, I. & C.A. Ricci 1979. Un evento granulitico nel basamento metamorfico della Sardegna nordorientale. Mem. Soc. Geol. It. 20: 23-38. Graham, C.M. & R. Powell 1984. A garnet-hornblende geothermometer: calibration, testing and application to the PeIona Schist, southern California. J. Metamorphic Geol. 2: 13-2 1 Kohn, M.J. & F.S. Spear 1989. Empirical calibration of of geobarometers for the assemblage garnet + hornblende + plagioclase + quartz Am. Mineral. 4: 77-84. Kretz, R. 1983. Symbols for rock-forming minerals. Am. Mineral. 68: 277-279. Laird, J. & A.L. Albee 1981. Pressure, temperature and time indicators in mafic schists: their application to reconstructing the polymetamorphic history of Vermont. Am. J. Sci. 281: 127-175. Leake, B.E., Woolley, A.R., Arps, C.E.S., Birch, W.D., Gilbert, M.C., Grice, J.D., Hawthorne, F.C.,. Kato, A., Kisch, H.J., Krivivichev, V.G., Linthout, K., Laird, J., Mandarino, J.A., Maresch, W.V., Nickel, E.H., Rock, N.M.S., Scumacher, J.C., Smith, D.C., Stephenson, N.C.N., Ungaretti, L., Whittaker, E. & G. Youzhi 1997. Nomenclature of amphiboles: report of the subcommittee on amphiboles of the international mineralogical association, commission on new minerals and mineral names. Can. Mineral. 35: 219-246.
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Water-Rocklnferaction2001, Cidu (ed.),02001 Swefs & Zeitlinger, Lisse, ISBN 90 2651 824 2
Fluid geochemistry of Tieluping Ag ore and its implications for the CPMF model Y.J.Chen, Y.H.Sui & X.L.Gao Department of Geology, Peking University, Beijing 100871, China State Key Laboratory of Endogenetic Deposits of Nanjing University, Nanjing 21 0093, China
ABSTRACT: The Tieluping silver deposit occurs in the Qinling Orogen, China. Its metallogenic process included three stages, with average temperatures of 373 "C, 223 "C and 165 "C respectively. The middle stage minerals have most fluid inclusions with highest content of ions, highest H2O/C02 ratio, and likely, have crystallized rapidly under conditions of fluid boiling and mixing, containing most ore elements compared to those of other stages. The early stage fluid, with 6D = -90 %o, d3Cc02 = 2.0 %O and 8'0 = 8.94 %o, was generated from reworking-metamorphism of carbonate-rich formation; the late one, with 6D = -73 %o, d3CCo2= -1.3 %o, 8'0 = -0.92 %o, was a meteoric hydrothermal solution; and the middle might be a hybrid of metamorphic fluid and meteoric hydrothermal solution. This three-stage fluidization and mineralization were caused by a Mesozoic intracontinental subduction, and are quite consistent with the tectonic model for collisional petrogenesis, metallogenesis and fluidization. 1 INTRODUCTION
Xionger Group, the cover of the basement. In the south of the Machaoying fault, the carbonaceous carbonate-silicalite-argilliteformations (CSAFs) of the Guandaokou and Luanchuan Groups are extensively developed. The mineralizing process included three stages, marked by mineral associations of quartz-pyrite, polymetallic sulfides and fine carbonate veinlets respectively.
The Tieluping Ag (Pb, Au) deposit (> 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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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
2.3 Rhyolite outcrops The rhyolitic lavas of the Monte Arci volcanic complex represent the more evolved products of a complete subalkaline sequence, from subalkaline basalt through basaltic andesite, andesite and dacite. They occur as very thick lava flows generally spreading westwards. Small pyroclastic pumice fall deposits outcrop near the venting area (e.g. Scala Antruxioni). The rhyolite lavas consist, from bottom to top, of obsidian, obsidian-perlite, perlite and vesicular flow-banded glassy rhyolite. In terms of texture, the perlite exhibits abundant
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intersecting, concentric, curviplanar cracks from a few millimetres to a few decametres in diameter. Macro-scale spheroidal cracks are well exposed in the Scala Antruxioni quarry. In the perlite-toobsidian transition zone, perlite encapsulates cores of intact obsidian glass. 2.4 Sampling The complete volcanic sequence, from obsidian to vesicular, flow-banded lava, through perlite with obsidian core and perlite, were carefully sampled in the two well-exposed outcrops at Conca Cannas and Scala Antruxioni perlite quarries. 3 ANALYTICAL METHODS Thirteen samples of whole rock were analyzed for major and trace elements (Table 1) by means of XRF, using the data reduction method proposed by Franzini et al. (1975) and Leoni & Saitta (1976). For samples containing the obsidian-perlite pair, the two facies were carefully separated. Electron probe micro-analysis (EPMA) was carried out with a WDS ARL-SEMQ device on selected samples of perlite and obsidian using natural silicate and oxide as standards and operating at 20 kV, and 20 nA. Raw data were corrected with the ZAF method.
The purpose of these analyses was to define compositional changes on the micro-scale from the obsidian core to the perlite rim, and to investigate the inhomogeneity of obsidian glasses. In order to minimize heating of the glass and resulting vapor exsolution, a large electron beam (about 10-15 pm in diameter) was used and the sample stage was continuously moved around the analyzing point, following the procedure adopted by Jezek & Noble (1978). Four samples of obsidian and perlite were prepared for 6"O and 6D analyses. Oxygen was extracted from volcanic glasses via laser fluorination with C1F3 as described by Sharp (1990). Hydrogen was liberated via heating to 1400°C in vacuum, and water reduced via uranium at 650°C (Bigeleisen et al. 1952). Isotope analysis was performed by mass spectrometry. Fourier transform infrared spectroscopy (FTIR) was performed on double-side polished shards from the two samples P8Pr and P80b to obtain point data on water concentration and speciation both in obsidian and perlite glasses. Data were obtained using a Nicolet Magna 560 FTIR spectrometer (Globar IR source, KBr beamsplitter) equipped with a NinPlan IR microscope (MCT/A detector). Three analyses were done for each shard (256 scans, 4 cm-' resolution). Data reduction was performed for molecular water and OH' related peaks at 1630, 3530, 4500 and 5200 cm-', using the molar absorptivity
Table 1. Major and trace element analyses of whole rock samples of perlite (Pr), obsidian (Ob) and vesicular flow-banded glassy rhyolite (V Rhy). D.I. = differentiation index; A.I. = agpaitic index. Pr Pr Rock type P5 P1 Sample 72.83 72.98 Si02 0.22 0.21 TiO2 13.26 13.39 A1203 0.66 0.11 Fe203 0.98 1.59 Fe0 0.04 0.04 MnO 0.24 0.23 MgO 0.68 0.68 CaO 3.13 3.06 Na20 5.57 5.55 K20 0.07 0.08 p205 2.31 2.09 LOI D.I. 90.11 90.15 0.84 0.83 A.I. Trace elements (ppm) Zn 50 50 247 248 Rb 63 57 Sr 305 306 Ba 150 151 Zr 29 28 Pb 28 26 Nb Y 30 29 34 35 La 66 58 Ce
Scala Antruxioni Pr Pr Pr P4 P6 P2 72.89 73.17 72.81 0.21 0.21 0.22 13.55 13.37 13.41 0.52 0.42 0.62 1.07 1.13 1.08 0.04 0.04 0.05 0.25 0.25 0.27 0.67 0.66 0.72 3.07 3.11 3.14 5.62 5.53 5.79 0.08 0.08 0.08 2.05 2.05 1.83 90.21 90.44 90.44 0.82 0.83 0.85 49 247 58 299 153 28 27 31 36 50
46 245 58 296 149 26 26 29 37 55
50 248 57 296 145 28 26 28 29 56
Pr P3 73.32 0.21 13.41 0.58 1.03 0.04 0.25 0.70 3.20 5.47 0.07 1.74 90.56 0.83
(3; 75.11 0.14 12.80 0.72 0.85 0.06 0.10 0.53 4.07 4.80 0.03 0.80 93.70 0.93
49 242 61 306 149 29 27 28 36 58
93 280 18 115 194 36 68 54 74 148
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1
Ob P10 74.79 0.12 13.66 0.41 1.09 0.07 0.11 0.62 3.40 5.18 0.07 0.49 92.30 0.82
Ob P90b 74.53 0.11 13.64 0.04 1.40 0.07 0.10 0.60 3.60 5.11 0.06 0.74 92.47 0.84
85 265 25 132 94 31 51 46 24 50
82 265 19 120 96 33 49 45 15 41
Conca C'annas Pr Ob P9Pr P8Ob 73.28 74.78 0.1 1 0.11 13.35 13.72 0.02 0.51 1.40 0.98 0.07 0.07 0.10 0.13 0.60 0.60 3.63 3.20 5.26 5.08 0.07 0.06 2.45 0.43 90.47 92.74 0.82 0.84 80 263 19 117 94 36 48 46 17 34
86 264 20 122 93 32 49 44 19 34
Pr VRhY P8Pr P11 73.52 72.03 0.1 1 0.26 13.44 13.64 0.6 1 0.40 1.08 1.14 0.07 0.04 0.40 0.1 1 0.79 0.60 3.04 3.3 1 5.53 5.22 0.10 0.06 2.41 2.08 90.92 88.58 0.8 1 0.83 81 265 17 127 90 32 48 42 19 38
51 233 89 489 186 31 26 27 45 86
values suggested by Zhang (1999) and Stolper (1 982). 4 GEOCHEMISTRY Whole rock analyses show a rhyolite composition (Table 1) with Si02 content ranging from about 72 to 75%. Na20 content ranges from 3.04 up to 4.07%, K20 from 4.80 to 5.79%, with total alkali of between 8.46 and 8.93%. Volatile contents (loss on ignition = LoI) show two distinct ranges: from 0.43 to 0.80% for obsidian and from 1.74 to 2.45% for perlite. All analyzed samples have normative corundum (up to 1.51) suggesting that alteration phenomena occurred (also resulting in Na2O removal). The close negative correlation between normative corundum and Na20 seems to validate this hypothesis. Microprobe analyses (about 60 analyzed spots, not reported here) show a trend similar to that of the rock as a whole, with inhomogeneity on the microscale. Based on these analyses, the difference to 100% (according to Jezek & Noble 1978) is assumed as a rough estimate of LoI, and varies from 0.02 to 3.56%. Si02 content ranges from about 72 to 76%, Na20 from 2.49 to 4.32%0,K20 from 5.05 to 6.16%, with total alkali of between 8.39 and 9.48%. The P7 obsidian sample, which yielded 0.80% LoI and an agpaitic index (A.I.) of 0.93 (on whole rock), shows micro-scale inhomogeneity in the glass which is clearly reflected in the A.I. value. For microprobe analyses, this A.I. defines a subaluminous to peralkaline character with values ranging from 0.88 to 1.02. According to Middlemost (1975), samples appear to belong either to the K-series or to the high-K series. The K20 vs Na2O diagram highlights the compositional variation from K-rhyolite to high-K rhyolite from nonhydrated obsidian to hydrated perlite glasses respectively (Fig. 1). At least two distinct
Figure 2. LoI vs Na20 diagram. Dots circles = whole rock analyses.
=
microprobe analyses;
populations clearly emerge from microprobe analyses. We suggest that this trend cannot be attributed to magmatic processes, thus indicating deuteric phenomena. The Na2O vs LoI diagram (Fig. 2) shows a significant negative correlation between obsidian and perlite. Compared to whole rock, the scatter of microprobe data is accentuated by the assumption (mentioned above) that the estimated LoI in the microanalysis is obtained via the difference method. 5 ISOTOPE AND WATER SPECIATION DATA 5.1 Isotope data The 6l80 and 6D of perlite and obsidian are reported in Table 2. 8l8O values range between 8.01*0.02 for obsidian and 10.23rt0.02 for perlite, with intermediate overlapping values. 6D values increase from obsidian (-102.45*0.02) to perlite (-85.06*0.02), the latter richer in deuterium. The generally lower 6l80 values found for obsidian sug est, according to Hoefs (1997), a change in the l80/ 530 ratio of perlite related to low-T alteration processes. A similar enrichment is shown by 6D values of perlite. Alternatively, this trend may be explained by degassing fractionation phenomena or interaction with exogenic water. Due to the inconsistency of 6D vs H2O values of the obsidian-perlite pair in the degassing models defined by Dobson et al. (1989), the first hypothesis can be rejected. According to Taylor (1974), 6D and 6I8O values of perlites, formed during hydration, show a better isotopic equilibrium with Sardinian meteoric water Table 2. Isotope analyses. STD = Standard deviation. ~~
sample P10 P80b P8Pr P5
Figure 1. K 2 0 vs Na20 diagram; solid line separates obsidian from perlite. Dots = microprobe analyses; circles = whole rock analyses.
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facies H20% Obsidian 0.43 Obsidian 0.35 Perlite 1.69 Perlite 1.77
6’*0%0 8.01 9.66 10.23 9.62
STD
0.02 0.04 0.02 0.02
~ D % o STD -97.51 0.02 -102.45 0.02 -87.19 0.02 -85.06 0.02
composition (Caboi et al. 1993) than with parent obsidian facies. On the whole, isotopic data suggest obsidian was progressively hydrated by meteoric water, leading to perlite genesis.
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. DAmico; Local Coordinator G. Macciotta).
5.2 Water speciation data
REFERENCES
FTIR data allow us to make some obsevations as to the nature of the water dissolved in the glass. Shards from obsidian fragments and from small (cm- to mm-sized) obsidian cores in the perlite, show a constantly low water content (0.24% k 0.03), irrespective of the size of the core and of its position therein. Nearly all the dissolved water (0.23% * 0.06) is composed of hydroxyl groups according to what observed in volatile-poor rhyolitic magmas (Zhang 1999). On the contrary, shards from perlitic lava have a higher, though still constant, water content (3.20% k 0.3 l), and more than 90% of the dissolved water is present in its molecular form (2.99% 2 0.29), as observed in glasses that have undergone secondary, low-T hydration (Newman et al. 1988). Furthermore, the homogeneous, high water content of perlitic glass, irrespective of its position within the rock, and the sharply different water content of small-sized obsidian cores versus their perlitic shell, suggest that secondary glass hydration proceeded inward in these lavas, progressively saturating the obsidian glass.
Assorgia, A., Beccaluva, L., Di Paola, G.M., Maccioni, L., Macciotta, G., Puxeddu, M., Santacroce, R. & G. Venturelli 1976. Geo-petrographic map of Monte Arci volcanic complex (Sardinia). Scale 1:50,000. Parma: Grajiche STEP. Beccaluva, L., Civetta, L., Macciotta, G. & C.A. Ricci 1985. Geochronology in Sardinia: results and problems. Rend. Soc. It. Min. Petr. 40: 57-72. Beccaluva, L., Maccioni, L., Macciotta, G. & G. Venturelli 1974. Dati geologici e petrochimici sul massiccio vulcanico di Monte Arci (Sardegna). Mem. Soc. Geol. It. 13 (2): 387405. Bigeleisen, J., Perlman, M.L. & H.C. Prosser 1952. Conversion of hydrogenic minerals to hydrogen for isotopic analysis. Anal. Chemistry 24: 1356-1357. Caboi, R., Cidu, R., Fanfani, L., Zuddas, P. & A.R. Zanzari 1994. Geochemistry of the high-P CO2 waters in Logudoro, Sardinia, Italy. Appl. Geoch. 8: 153-160. Dobson, P.F., Epstein, S. & E.M. Stolper 1989. Hydrogen isotope fractionation between coexisting vapor and silicate glasses and melts at low pressure. Geoch. Cosm. Acta 53: 2723-2730. Franzini, M., Leoni, L. & M. Saitta 1975. Revisione di una metodologia analitica per fluorescenza X basata sulla correzione completa degli effetti di matrice. Rend. Soc. It. Min. Petr. 31: 365-378. Jezek, P.A. & D.C. Noble 1978. Natural hydration and ion exchange of obsidian: an electron microprobe study. Am. Miner. 63: 266-273. Hoefs, J. 1997. Stable isotope geochemistry. Berlin: SpringerVerlag Leoni, L. & M. Saitta 1976. X-ray fluorescence analysis of 29 trace elements in rock and mineral standards. Rend. Soc. It. Min. Petr. 32: 497-5 10 Middlemost, E.A.K. 1975. The basalt clan. Earth Sci. Rev. 11: 337-364. Newman, S., Epstein, S. & E.M. Stolper 1988. Water, carbon dioxide and hydrogen isotopes in glasses of the ca. 1340 A.D. eruption of the Mono Craters, California: constraints on degassing phenomena and initial volatile content. J. Volcanol. Geotherm. Res. 35: 75-96. Sharp, Z.D. 1990. A laser based microanalytical method for in situ determination of oxygen isotope ratio of silicates and oxides. Geoch. Cosm. Acta 54: 1353-1357. Stolper, E. 1982. Water in silicate glasses: an infrared spectroscopic study. Contrib. Mineral. Petrol. 8 1: 1- 17. Taylor, H.P.Jr 1974: The application of oxygen and hydrogen isotope studies to problems of hydrothermal alteration and ore deposition. Economic Geology 69: 843-883. Zhang, Y. 1999. H 2 0 in rhyolitic glasses and melts: measurement, speciation, solubility, and diffusion. Rev. of Geophysics 37: 493-516.
6 CONCLUSIONS The geological, geochemical and isotopic data enabled us to single out the phenomena underlying the formation of the Monte Arci perlite deposits. The process involves interaction between degassed volcanic glasses (obsidian) and meteoric water, to form hydrated glasses (perlite). Water-rock interaction implies the following geochemical changes in the original glasses: - different LoI in obsidian and perlite, indicated by a compositional gap; - lower Na2O content in perlite, also detected by microprobe analyses; - different hydrogen and oxygen isotopic composition in obsidian and perlite; - sharply different hydroxyl and molecular water content in obsidian (OH-) and perlite (H2O). In this context, degassing magma processes may have played only a minor role, masked by more important secondary exogenic hydration.
ACKNOWLEDGMENTS This work was carried out with the support of the Local Scientific Research Programme (Coordinator M. Marchi) and
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Trace element mobility in tourmalinite veins and surrounding metapelites from the Crummock Water aureole (Lake District, England) C .Corteel Vukgroep Geologie & Bodenzkunde, Uiziversiteit Gent, Belgium
N .J .Fortey British Geological Survey, U K
ABSTRACT: The Crummock Water aureole consists of Ordovician slates which have undergone retrogressive metamorphism and contains several fractures displaying tourmalinite wallrock alteration zones, which range in thickness from about a centimetre to a few metres. Several larger tourmalinite veins often occur together in anastomosing vein systems, continuing over several hundreds of m. Both rock types display low field strength element depletions which in the host rocks decrease on decimetre scale with increasing distance from the veins. REE, especially the LREE, display different degrees of mobility due to tourmalinisation, both in tourmalinites and host rocks close to the veins. For the tourmalinites, differences in REE mobility are possibly related to the position of the rock on the fluid circulation path along the fractures.
1 INTRODUCTION
2 RESULTS
The Crummock Water aureole is an ENE-trending elongate area of bleached slates around Crummock Water in the English Lake District (UK). According to Cooper et al. (1988) it represents a zone of lowgrade retrogressive contact metamorphism. The metamorphic event hydrothermally altered and indurated the dark coloured Skiddaw Group silt- and mudstones of the Kirkstile Formation. Gravimetric modelling (Lee 1986) indicates that the metamorphism was probably caused by an inferred and concealed highly evolved granitic intrusion, forming part of the Lake District batholith. The aureole is transected by a series of steep fractures, often sealed by quartz veins, which are occasionally brecciated and regularly accompanied by wallrock alteration zones consisting of fine-grained black tourmalinites. The tourmalinites equally occur without ‘central’ quartz vein and range in thickness from about a centimetre to a few metres. Often several large tourmalinite veins form anastomosing vein systems, continuing over hundreds of metres. In some cases the veins pass into lobes penetrating over several m along the bedding into the host rocks. Tourmalinites extent over 7 km and vertically they occur over some 750 m.
2.1 Analytical methods Solutions for determining trace elements were prepared by dissolving melts of 500 mg sample and 1.5 mg lithium-metaborate in a 100 ml aqueous solution containing 5 ml HNO3 (65 %) and 1 ml HF (48 %). All trace elements were determined on an Elan 5000 Perkin-Elmer ICP-MS. 2.2 Petrography Tourmalinite-associated quartz veins sporadically contain tourmalinite fragments reflecting vigorous hydrothermal circulation. In some tourmalinites sedimentary bedding can be traced. When examined in thin section most tourmalinites - especially those in the wallrock alteration zones - show a fine-grained irregular mixture of quartz and tourmaline. When present, banding is expressed by rapid changes in the tourmaline/quartz ratio. Tourmaline is generally the most important component and according to Fortey & Cooper (1986) it has an intermediate schorl-dravite composition. Tourmalines are predominantly anhedral and green, but larger prismatic crystals frequently posses other colours (mostly brown), espe697
Table 1. Representative trace element abundances in 3 host-rocks and 4 tourmalinites (in ppm). Analyte
V Cr CO Ni cu Ga Rb Sr Y Nb cs Ba La Ce Pr Nd Sm Eu Tb Gd DY Ho Er Tm Yb Lu Hf Ta Th
Host rocks KRE.2 111 92 15.7 46.5 26.5 29.6 11 1.0 91 17.2 17.0 3.5 434 8.3 47.8 2.5 9.4 2.55 0.67 0.53 2.87 3.40 0.77 2.46 0.84 2.71 0.40 5.7 1.2 12.4
KRG.l 114 101 13.1 37.3 3.8 27.6 123.6 191 22.3 18.8 3.3 368 16.8 55.1 4.2 15.3 2.94 0.96 0.57 3.20 3.96 0.89 2.86 0.96 2.94 0.45 5.7 1.3 13.4
KRTlH.8 118 103 21.4 44.0 63 .O 34. I 181.6 93 33.9 18.3 4.5 615 20.2 95.6 5.9 21.3 4.88 1.24 1.01 5.32 6.30 1.31 3.85 1.24 3.71 0.52 5.4 1.4 14.5
Tourmalinites LC.304B KRTlB.5 106 102 86 62 4.0 9.6 32.7 38.4 11.4 b.d. 30.5 22.1 86.4 13.7 44 56 14.2 31.6 16.6 7.2 0.9 0.2 186 19 5.3 32.0 12.8 130.7 1.6 10.4 6.2 38.4 1.39 8.76 0.29 1.20 0.33 1.25 1.70 8.15 2.37 7.03 0.55 1.32 1.83 3.68 0.63 1.09 2.15 3.04 0.33 0.40 4.7 2.4 1.2 0.5 13.2 7.1
cially when they display optical zonations. Accessories are restricted to some percentages (max. 10 %) and are mainly rutile, chlorite and to a lesser extent muscovite. All of these lie scattered in the tourmalinite mass, although chlorite and muscovite mostly occur in veins. Other textures are mainly restricted to veins and are either poikiloblastic, with disseminated tourmaline in larger quartz crystals, or brecciated. Last mentioned texture consists of metasomatic tourmalinite fragments in a quartz and/or poikiloblastic quartz-tourmaline matrix. The host rocks consist of pale (and dark) grey banded metapelites, which lost their dark colour and most of their original cleavages during retrograde metamorphism. In thin section a bedding-parallel lepidoblastic fabric can be observed, mainly consisting of muscovite and chlorite. Quartz and some rutile are present as coarser grains (> 64 pm) and the occurrence of quartz is mostly bedding-related. 2.3 Trace element compositions Table 1 gives the abundances of 30 trace elements in three metapelites and four tourmalinites of the Crurnmock Water aureole. Figure 1 shows North
KRTlH.5 104 88 19.2 54.7 1.8 28.0 63.3 54 33.2 16.7 1.1 175 83.O 149.5 20.1 74.0 14.16 2.43 1.73 13.30 7.91 1.36 3.49 1.06 3.23 0.48 5.4 1.2 13.4
KRTIM.I 89 80 13.1 52.3 b.d. 22.3 12.5 39 14.1 20.4 0.4 17 1.1 2.5 0.3 1.o 0.36 0.18 0.25 0.92 2.13 0.53 1.88 0.68 2.36 0.36 5.3 1.3 12.5
American Shale Composite (NASC) normalised abundances of REE (normalising values are from Gromet et al. 1984). REE patterns of metapelites at several tens of metres distance from tourmalinite veins (Fig. la) display minor depletions or enrichments of especially the light(L) FEE. REE abundances of tourmalinites from one anastomosing vein system (Fig. lb) show a more intense depletion of the LREE for some samples. In particular, the depletion increases more or less with increasing distance from outcrop A. Figure l c shows one tourmalinite (H5) and three metapelites at increasing distance from the tourmalinite vein where sample H5 was taken. The sample close to the tourmalinite vein (H6) displays a considerable REE depletion which is larger for the lighter elements. This depletional trend is clearly smaller for the host rock at half a metre distance from the vein and diminishes further with increasing distance from the vein (Fig. l ~ ) . NASC normalised spider diagrams are represented in Figure 2. The host rock compositions at several tens of metres distance from the tourmalinites (Fig. 2a) all lie close to the normalising values while all tourmalinites (Fig. 2b,c) show a depletion of the low field strength elements Cs, Rb and Ba. In Figure 2c
the metapelite sample close to the tourmalinite vein (H6) equally displays a depletion of this three elements, while in the host rocks further away from the vein (H7-8) these depletions become smaller with increasing distance from the vein.
Figure 2. NASC normalised SPIDER diagrams of the same samples as in figure 1 (see figure 1 caption for explanations of sample details).
3 DISCUSSION As usual for tourmalinites (e.g. Bajwah et al. 1995; Slack et al. 1993), tourmalinisation transformed the host rock mineralogy nearly completely into tourmaline and quartz. The preservation of the s&mentary bedding indicates a metasomatic replacement origin. The differences in REE patterns in the metapelites (Fig* la) might be by redistributions of these elements due to the contact metmorphism as previously suggested by Cooper et al. (1988).
Figure 1. NASC normalised REE abundances of: a (upper): 4 metapelites at considerable distance of tourmalinites; b (middle): 7 tourmalinites from one anastomosing vein system, distances of samples from outcrop A: B5: 10 m, H2: 200 m, K1: 290 m, L l : 390 m, M1: 580 m, N2: 760 m.; c (lower): H5: tourmalinite at the edge of a vein, other samples are host rocks taken in a straight line at increasing distances from sample H5: H6: 5-10 cm; H7: 0.5 m and H8: 5 m.
699
Depleted values of low field strength elements in tourmalinites relative to unaltered precursors have been interpreted by Bajwah et al. (1995) as reflecting strong metasomatic tourmalinization. Although this conclusion matches our observations for the metasomatic tourmalinites from Fig. 2b, we would rather attribute these depletions for the samples analysed in this study to more intense (longer) hydrothermal fluid migration because the host rock close to the tourmalinite vein (Fig. 2c) equally displays considerable depletion of Cs, Rb and Ba. Fluid migration through this rock was very similar to fluid passage in the metasomatic altered wallrock, but a critical factor, probably the availability of Boron in the fluid, was missing to form a tourmalinite. In general, duration of tourmalinisation is believed to be correlated with the amount of LREE depletion (e.g. Slack et al. 1993), which explains the differences between the REE patterns for the tourmalinites in Figures lb,c. According to the observations in Figure l b these differences (and times of exposure to tourmalinisation) are related to the relative position of the samples, probably reflecting their position on the fluid circulation path along the fractures. The observed patterns in Figure l c could suggest that for the considered samples REE migrated from the rocks bordering the tourmalinites into the veins, an effect quickly decreasing on decimetre scale and most pronounced in the zone where hydrothermal activity occurred without tourmalinisation (sample H6 in Fig. lc).
REFERENCES Bajwah, Z. U., White, A. J. R., Kwak, T. A. P. & R. C. Price 1995. The Renison Granite, Northwestern Tasmania: A petrological, geochemical and fluid inclusion study of hydrothermal alteration. Economic Geology 90: 1663-1675. Cooper, D. C., Lee, M. K., Fortey, A. H., Cooper, A. H., Rundle, C. C., Webb, B. C. & P. M. Allen 1988. The Crummock Water aureole: a zone of metasomatism and source of ore metals in the English Lake District. Journal of the Geological Society, London 145: 523-540. Fortey, N.J. & D. C. Cooper 1986. Tourmalinisation in the Skiddaw Group around Crummock Water, English Lake District. Mineralogical Magazine 50: 17-26. Gromet, L. P., Dymek, R. F., Haskin, L. A. & R. L. Korotev 1984. The “North American shale composite”: Its compilation, major and trace element characteristics. Geochimica et Cosmochimica Acta 48: 2469-2482. Lee, M. K. 1986. A new gravity survey of the Lake District and three-dimensional model of the granite batholith. Journal of the Geological Society, London 143: 425-435 . Slack, J. F., Palmer, M. R., Stevens, B. P. & R. G. Barnes 1993. Origin and significance of tourmaline-rich rocks in the Broken Hill District, Australia. Economic Geology 88: 505-541.
4 CONCLUSIONS Petrography and trace element abundances suggest metasomatic replacement by hydrothermal fluid migration for the genesis of the Crummock Water tourmalinites. According to low field strength element abundances the hydrothermal fluid migration equally occurred in the wallrocks. This fluid activity in the host rocks is strongest close to the zone of tourmalinisation and quickly decreases over several decimetres. REE have been proven to be relatively mobile under the conditions of formation in at least a part of the studied tourmalinites and host rocks. Concerning the host rocks the REE mobility is roughly comparable to that of the low field stength elements. For the metasomatic tourmalinites the degree of mobility might be related to the position of the rock on the fluid circulation path, althoug more analyses, currently running, need to confirm this preliminary conclusions. 700
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Gold ore-system in Archean greenstone structures of Middle-Dnieper Area (Ukrainian Shield) Yu .Fomin, He .Lasarenko , Yu .Dernikhov & V1.Blazhko State Scientific Center of Environmental Radio geochemistry Nat. Acad. of Sci. Kyiv, Ukraine
ABSTRACT: The subject of this paper is gold deposits from Sura and Chertomlyk Archean granitegreenstone structures within central part of the Ukrainian Shield. Five types of ores were recognized: Au-AgBi-Te, Au-bearing sulphide, Au-bearing Cu-MO (Sura) and Au-Fe, Au-Ag-Pb-Zn (Chertomlyk). On the base of geologic, mineralogic, fluid inclusion, and light stable isotope (S34S,S13C, S**O,SD) studies the conditions of gold deposit formation were established. The multi-stage model of ore system evolution, including a primary volcanogenic ore genesis, regional metamorphism, diaphtores (retrogressive metamorphism) and epithermal redistribution is proposed. Gold-ore processes occurred under near neutral or slightly alkaline conditions. The environment was reduced with high sulfur activity. P-T ranges were: 390-160°C and 0.81 . 2 ~ 1 Pa. 0~
and (3) Au-bearing Cu-MO porphyry. The first two types were studied. Ores are controlled by local tectonic zones of cataclasites. Metavolcanic rocks in these zones were transformed into albite-epidote-quartz-carbonatechlorite diaphthoritic schists. Zonal quartzcarbonate-amphibole metasomatites are the important type of gold-associated alteration products. Complex Au-Ag-Bi-Te ores consist of pyrrhotite, pyrite (several generations), arsenopyrite, chalcopyrite, sphalerite, galena, magnetite, scheelite, bismuth, tellurides of Bi, Au, Ag, bismuth-lead sulfosalts and native gold. Isotopic age (lead-isotope ratio in galena) of Sura ore deposits is 2,650-2,550 Ma. Stratiform gold-bearing sulfide ore was also found in the upper parts of the volcano-sedimentary sequence. Gold-bearing pyrite, marcasite, pyrrhotite, chalcopyrite and sphalerite are the main minerals. Bismuth and antimony-lead-bismuth sulfosalts (kobellite-stibio cosalite) and silver tetrahedrite are rare accessory phases. The Balka Shyrokaya deposit is located in the Chertomlyk structure. Tholeitic metabasalts, quartzcarbonate-chlorite schists and banded iron formation (BIF) are the local rocks. Two types of ores were distinguished: (1) Au-Fe
1 INTRODUCTION Recently discovered within central part of the Ukrainian Shield gold deposits are connected with the different structures of the single Middle-Dnieper Archean granite-greenstone region. The comparison of these deposits with the well known gold-bearing greenstone belts of the world (Cameron & Hattori, 1985; Ho and others, 1985; Ronde and others, 1992; Shibetsky and others, 1997) shows that they are identical. The problem of these gold deposit genesis remains still open to discussion.
2 GEOLOGY AND ORE TYPES The Sergeevsk and Balka Shirokaya gold deposits were studied. The Sergeevsk deposit is located in the Sura structure, characterized by the occurrence of metavolcanic rocks. Tholeitic metabasalts are dominant; metadacites and metaultrabasites are subordinated. Metamorphic grade of ore-bearing rocks is epidote amphibolitic facies. Three types of ores were recognized: (1) Au-Ag-BiTe in zones of metasomatites - quartz-carbonateamphibole, chlorite-albite and quartz-carbonatesericite; (2) Au-bearing stratiform sulphide bodies 701
and (2) Au-Ag-Pb-Zn. Both types are in altered zones of quartz-carbonate-sericite metasomatites. Ores are controlled by tectonic zones of schist, cataclasite, and breccia, which were metamorphosed in greenschist fasies. The types of gold deposits include gold-iron ores in BIF and gold-silverpolymetallic ores in iron-rich metavolcanic and metasedimentary rocks. Both are accompanied by zones of beresites, which are altered rocks consisting of quartz, carbonates, white mica, biotite, and pyrite. Gold-iron ores are post-beresite tectonic breccias of BIF. They have a quartz-carbonate-sulfide matrix (Fomin and others, 1994). Gold-silver-polymetallic ores are connected with post-beresite cataclastic zones (Fomin and others, 1996). Both types mentioned above include goldbearing pyrite, arsenopyrite, pyrrhotite, chalcopyrite, sphalerite, native gold. Besides, ore of the first type contain magnetite, ferrotetrahedrite; the second one silver tetrahedrite (freibergite), boulangerite, galena, electrum. Isotopic age (lead-isotope ratio of galena) of ore deposits is about 2,860 5-100 Ma.
Table 2. Isotope composition of fluid inclusion components in hydrothermal quartz of gold ore deposits at Sergeevsk.
Fluid components H20 Quantity, mg/g 6D, %o P O , %o Quantity, mg/g 613C, %o P O . %o
Table I. Isotopic data of gold ore deposits at Sergeevsk.
Sulfides 634S,%O +0.3 to +6.5 Range Mean +2.7 (45)* Magnetite 6I8O, %O Range -1.6 to -0.9 -1.2 (6) Mean Carbonate minerals 613C,%O -1.9 to +2.0 Range Mean -0.5 (22) Carbonate minerals 6l'O. %O Range +9.0 to +16.4 Mean +11.1 (23) Quartz 6 " 0 , %o Range +9.9 to +10.6 Mean 1-10.2 (2)
0.10 to 0.33 -61 t0+38 -2.1 to +1.3
0.03 to 0.18 -75 t0+16 -7.1 to -1.5
0.08 to 0.10 -1.0 to +1.5 +13.6 to 21.8
0,03 to 0.1 1 -5.2 to +4.8 +19.6 to 27.3
The temperature of mineral formation was determined by the homogenization of gas-liquid inclusions, the pressure - in application to standart methods by syngenetic inclusions: gas-liquid (watersalt) - gas-liquid with C02. T-P conditions of Sergeevsk ores were: Au-Ag-Bi-Te type - 360180°C (quartz), 260-130°C (carbonates) and 0.81.2~10'Pa; Au-sulfide type - 350-140°C (quartz). For ores of Balka Shirokaya these conditions were: Au-Fe type - 380-120°C (quartz) and 0.5-0.9 x108 Pa; Au-Ag-Pb-Zn type - 420-1 1O'C (quartz). Ore deposits of all types in both structures were similar with respect to major mineralogical, thermobarogeochemical (fluid inclusions) and isotopic (S,C,O,H) characteristics.
The isotopic composition of sulfide sulfur, carbon, oxygen and hydrogen of carbonates, quartz, magnetite and H20, CO2 of inclusion water was investigated. Isotopic characteristics of minerals and fluid inclusions of all ore types are presented in tables 1 , 2 (Sergeevsk) and 3 , 4 (Balka Shirokaya).
Au-Ag-Bi-Te ore deposits
Au-sulfide ore deposits*
*Range of 3 (Au-Ag-Bi-Te ore) and 6 (Au-sulfide ore) measurements
3 RESULTS
Minerals
Au-Ag-Bi-Te ore deposits*
Table 3. Isotopic data of gold ore deposits at Balka Shirokaya (Fomin and others, 1995).
Au-sulfide ore deposits
Minerals
-0.2 to +5.2 +2.3 (41)
Au-Fe ore deposits
Sulfides 634S.%O Range -4.1 to +6.1 Mean +2.7 (23) Magnetite 6l80, %O Range -1.2 to +4.4 Mean +1.2 (21) Carbonate minerals 613C,%O Range -5.6 to -1.2 Mean -3.8 (17) Carbonate minerals 6"O. %O Range +9.2 to +11.4 Mean +10.8 (6) Quartz 6"0, %O Range +8.7 to +11.6 Mean +10.2 (8)
+2.4
--- (1) -2.2 to 0 -0.9 (7) +10.8 to 15.4 +12.8 (6) +8.7 to +10.6 +9.7 (3)
* Number of measurements 702
Au-Ag-Pb-Zn ore deposits -4.1 to +3.9 +1.4 (26) -2.6 to +2.3 -0.6 (6) -5.2 to -2.2 -3.7 (9) +10.8 to +12.1 +11.7 (6) +8.8 to +10.7 +9.9 (8)
Table 4. Isotope composition of fluid inclusion components in hydrothermal quartz of gold ore deposits at Balka Shirokaya (Fomin and others, 1995).
Fluid components H20 Quantity, mg/g 6D, %o P O . %o
Au-Fe ore deposits*
Au-Ag-Pb-Zn ore deposits*
0.1 1 to 0.50 -82 to -25 -3.5 to +2.6
0.15 to 0.62 -82 to -29 -2.7 to +0.8
Quantity, mg/g 6I3C, %o 6l80. %o
0,Ol to 0,96 -5.8 to +2.5 +17.6 to 31.4
0.04 to 0.16 -6.6 to -2.2 +141 to 21.0
*Range of 1 1 (Au-Fe ore) and 9 (Au-Ag-Pb-Zn measurements.
ore)
4 CONCLUSIONS In the gold-bearing altered zones, sulfides and carbonates inherited sulfur and carbon from the host rocks, with evidence of isotopic fractionation. The country rocks may be also the major gold source. Fluids become gold-bearing by interaction with rocks during their ascent in the crust. Fluid source had a complex nature. The meteoric component was dominant, and fluid was isotopically homogeneous. The wide 6D, 6l80, 6I3C, 634S variations of minerals and H20, CO:! from fluid inclusions are mainly a reflection of processes in the ore zones as a result of fluid-wall rock interaction. These results conform to the concept of multistage ore-system evolution: primary ore mineralization of volcanic genesis, then regional metamorphism of greenschist - epidote-amphibolite grades, and retrogressive metamorphism and epithermal gold redistribution in the altered zones. At the final stage of ore- forming process the gaswater fluids penetrated into rocks at T-P ranges: 390-160" C and 0.8-1.2~10' Pa. They were nearly neutral or slightly alkaline. Reduced conditions and high sulfur activity existed there. The content of H20, N2, CO increased in connection with the increase of gold concentration. Ores of all types were deposited in conditions of declining temperatures (homogenization of fluid inclusions and isotopic): from 390-320°C (quartzmagnetite assemblage) to 280-2OO'C (quartz-pyritearsenopyrite assemblage), to 260- 160°C (carbonatespyrrhotite-chalcopyrite-sphalerite and galenasulfosalts-gold assemblages), and pressures: from 23 (calculated) to 0.8-1.2 xl0' Pa. The gold-ore
mineralization was precipitated from near neutral or slightly alkaline fluids (pH 6-7) during the final stage of process. These fluids had hydrocarbonate potassium - sodium compositions with admixture of chlorine- and sulfate-ions. The condition were reducing (log fo2 from -39 to -38) with high sulfur activity (log as2 from -15.7 to -13.4) (Shibetsky and others, 1996).
REFERENCES Cameron, E.M. & K. Hattori 1985. The Hemlo gold deposit, Ontario: A geochemical and isotopic study. Geochim. et Cosmochim. Acta 49: 204 1-205 1. Fomin, Yu.A., Demikhov, Yu.N., Shibetsky, Yu.A. & N.M. Gostyaeva 1995. Gold-bearing epigenetic systems of the Ukrainian greenstone belt. WRZ-8. Kharaka&Chudaev. Balkema, Rotterdam : 667-669 Fomin, Yu.A., Demikhov, Yu.N., Shibetsky, Yu.A., Lasarenko He. & V.I. Blazhko 1996. Au-Ag-Pb-Zn ore of Balka Shirokaya (Middle Pridneprovie). Mineralogicheski journal. 1 : 74-87. Fomin, Yu.A., Savchenko, L.T., Demikhov, Yu.N., Gostyaeva, N.M. & He. Lasarenko 1994. Gold-jaspilite ore of Balka Shirokaya (Middle Pridneprovie). Geologicheskijournal. 3 : 84-95. Ho S.E., Groves D.I. & N. Phillips 1985. Fluid inclusions as indicators of the nature and source of ore fluids and ore depositional conditions for the Archean gold deposits of the Yilgarn block. Western Australia. Trans. Geol. Soc. S. Afr. 88: 149-158. Ronde S.E.J., Spooner E.T.S., Wit M.J. & C.J. Bray 1992. Shear zone-related, Au-quartz vein deposits in the Barberton greenstone belt, South Africa: field and petrographic characteristics, fluid properties and light stable isotope geochemistryJbid. 87: 366-402. Shibetsky, Yu.A., Fomin, Yu. A. & Yu.N. Demikhov 1997. Isotope-geochemical comparison of gold deposits of Archean greenstone belts. Mineralogicheskijournal. 1 : 3850.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Late Hercynian fluid circulation in the Charroux-Civray plutonic complex, NW Massif Central, France Regina Freiberger Institute of Mineralogy, Petrology and Geochemistry,Albert-Ludwig-University of Freiburg, Albertstr. 23b, 79104 Freiburg, Germany
Marie-Christine Boiron, Michel Cathelineau & Michel Cuney CREGU-UAM, 7566 G2R, BP 23,54501 Vandoxvre-l&-Nancy,France
ABSTRACT: The Hercynian calc-alkaline plutonic complex of Charroux-Civray (3 50-360 Ma) was intruded by a slightly younger (-320 Ma) peraluminous leucogranite. Several episodes of hydrothermal alteration have been recognized in the entire batholith. Based on fluid inclusion studies and mineralogical evidences, the P-T evolution of the early postmagmatic history of the plutonic complex can be redrawn. Primary aqueous FI within subsolidus epidote indicate relatively high pressure up to about 4kbar for the very early postmagmatic alteration, comprising the formation of hrther Ca-A1 silicates formed during plutonite cooling down to about 250°C. Aqueous-carbonic FIP of different densities can be correlated with restricted greisen-like alteration, including white mica, quartz, tourmaline, pyrite and carbonate in the P-T range of about 450°C and 1.5 to 2kbar. This hot fluid circulation causing greisen alteration is most probably induced by intrusion of the younger leucogranite, distinctly after cooling of the calc-alkaline magmatic complex, or by abnormal heat flow in relation to the delamination of the chain. Secondary aqueous FIP as well as nitrogen-bearing liquid and vapor FIP correspond to later hydrothermal alteration with decreasing T (<400"C) in the P-range of -0.5kbarYwhich caused later multiphase pervasive alteration. 1 INTRODUCTION
2 PETROGRAPHY
In the Charroux-Civray area, located at the NW margin of the French Massif Central south of Poitiers, the Hercynian basement is covered by 150m of Mesozoic sediments. The plutonic basement was studied on drill cores performed by ANDRA (France) to explore a potential site for radioactive waste disposal studies in crystalline rocks. An area of about 125 km2 has been explored by mapping, geological and geophysical studies and through the study of 17 continuously cored and oriented boreholes, with depths ranging from 200 m to 1000 m. The plutonic complex consists of a medium-K calcalkaline and a high-K (monzonitic) calcalkaline association (Cuney et al. 1999) and was emplaced during an early stage of the Hercynian orogeny (350-360 May U-Pb, Bertrand et al. 1996). A twomica leucogranite intrusion, found in one drilling, is considered as a small part of a large peraluminous body at depth (-320 Ma, comparable to neighboring leucogranites, Duthou et al. 1984), inferred from major geophysical anomalies (Virlogeux et al. 1999). Calc-alkaline magmatic suites are affected by strong pervasive alteration that display heterogeneous distribution, probably in response to heterogeneous paleofluid flows.
Several episodes of hydrothermal alteration have been recognized in the batholith. Pervasive alteration affects the calcalkaline granitoid rocks of Charroux-Civray with variable intensities, depending on the degree of microfracturing and on the proximity to fractures. Overall, early postmagmatic alteration of the Charroux-Civray plutonic complex is characterized by changing mineralogy (Fig. 1) and can be grouped into four main alteration stages (Fig. 2).
2.1 Ca-AI silicate assemblage Very early postmagmatic alteration (stage 1) related to the course of plutonite cooling, includes pervasive formation of several Ca-A1 silicates in otherwise fresh rocks (Freiberger et al., in review). First subsolidus alteration comprises formation of sub-euhedral epidote within distinct zones (Fig. la), partly associated with slight ductile deformation. Further Ca-A1 silicates include hydrogarnet, post-dated by prehnite and pumpellyite (partly in textural paragenesis), Fe-rich epidote (pistacite) and later laumontite, predominately in small fissures associated with adularia (Fig. 1b). 705
Fig. 1 : Microphotographs of characteristic alteration products. a) early subsolidus epidote, b) Ca-AI silicate assemblage within biotite, c) bulky sheets of white mica (greisen), d) tourmaline associated with quartz (greisen), e) chloritization of biotite, associated with K-feldspar, f ) argilization of feldspar.
2.2 Greisen assemblage
feldspar) (Fig. 1e), frequently associated with anatase or local carbonate, iii) Ca-plagioclase is replaced by phengite and calcite, K-feldspars are better preserved but display phengite patches, iv) titanite is replaced by anatase and carbonate. Additionally the whole mineral association of this stage can be found in thin microfractures (<200 pm).
Some occurrences of greisen-like assemblages (stage 2) within the calc-alkaline rocks were found in the vicinity of quartz fractures. In general the greisen zones are limited to about ten centimeters each side of the associated fracture. The greisen assemblage in the granitoid rocks is characterized by an almost complete pervasive replacement of the main rock forming minerals (feldspars, Fe-Mg silicates) by bulky sheets of white mica (mainly muscovite to phengite composition, Fig. lc). Newly formed quartz crystals mainly occur within fracture zones. Small needles of tourmaline (Fig. Id) are found in some greisen samples, generally occurring in fracture zones. Euhedral pyrite may grow pervasively in zones dominated by white mica. Chlorite frequently fills grain boundaries and vugs within the greisen zones and therefore it may be considered as post-dating muscovite. Anhedral to sub-euhedral carbonate occurs as slightly younger formation, both in fi-acturesand pervasively.
2.4 Illite-chlorite-dolomite assemblage Locally, an illite-chlorite-dolomite~calcitemineral assemblage occurs. Sheet silicates are mainly illite or ordered mixed layered illite/smectite with a high content of illite (Fig. If). This stage 4 presents similarities with that of stage 3, but is locally replacing or invading all earlier alteration minerals, e.g. Ca-A1 silicates of alteration stage 1 are altered mainly to chlorite and illite. It is indicative of lower temperanagmatic stage Hgr, Pr, Pm, Ps:
2.3 Chlorite-phengite-dolomite assemblage Pervasive alteration of stage 3, postdating the two first stages, is by far the most predominant and produces the largest alteration halos. It is characterized by the following association: chlorite-phengite-dolomite*calcite. Alteration minerals are mostly found in close association with the primary minerals: i) hornblende is altered to chlorite, calcite and Fe-oxide, ii) biotite is replaced by chlorite (=tK-
taumontite Adularia
4
4 Phengite
: Chloritei
+
Phengite
Tourmaline4
_____
i
'
:
-Chlorite
~~~~
*
i-
Fig. 2: Relative crystallization sequence of postmagmatic alteration stages 1 to 4 (Fig. 2) affecting the calc-alkaline plutonites.
706
ture (low temperature epithermal type quartz (quartz combs) and illite) as also reflected by fluid inclusion data. Similar mineral assemblages are locally found in macroscopic veins. 2.5 Post-Hercynian alteration The Charroux-Civray area is more recently affected by several fluid percolation stages, which produce specific mineral assemblages. All are characterized by low temperatures and are considered as later than the uplift and Permian-Triassic erosion (Cathelineau et al. 1999). These processes did not affect the studied samples. 3 LOCATION AND TYPOLOGY OF FLUID INCLUSIONS
The calc-alkaline plutonic rocks from CharrouxCivray were repeatedly affected by pervasive fluid percolation. Microfractures in quartz grains from calc-alkaline plutonic rocks appear as healed (fluid inclusion planes - FIP) and may bear all fluid types, although alterations are not developed. In other cases, the abundance of FIP is such in strongly altered zones that the fluid inclusions can be related with a rather good confidence to the alteration facies. Based on microthermometric data and Raman microspectroscopy, the fluids from the CharrouxCivray samples may be classified into three compositional groups: (i) carbonic fluids (Lc(-w), Lc-w, Lw-c), (ii) aqueous fluids (LwI, LWII),and (iii) nitrogen-bearing fluids (Lw-n, Vn). Raman spectroscopy shows that the vapor composition varies between H20 and CO2 and H20 and N2 end-members and therefore it can be divided into two compositional trends.
Fig. 3: P-T diagram showing the isochores of the investigated fluid inclusions with interpretative P-T-ranges (gray fields) for fluids of different composition.
silicates during a very early postmagmatic step (stage 1). This very early subsolidus P-T range is in good accordance with the assumed P-T conditions for the late-magmatic stage (Cuney et al. 1999). The following period of the early postmagmatic stage is characterized by the retrograde crystallization of Ca-AI silicates like hydrogarnet, prehnite, pumpellyite, pistacite and laumontite during successive cooling of the granitoids (Freiberger et al. in press). The development of the common paragenesis prehnite+pumpellyite indicates a rather narrow temperature range of 200 to 280°C and pressures of 2 to 3 kbar (Frey et al. 1991). 4.2 Carbonicjluids Lc(-w), Le-w and Lw-c Postdating the Ca-A1 silicate formation, the greisen alteration within the calcalkaline plutonites is frequently associated with CO2-bearing fluids and can be divided into almost pure CO2 inclusions Lc(-w), 3-phase carbonic inclusions Lc-w with relatively high density and less dense 2-phase aqueous-carbonic inclusions Lw-c. They show homogenization temperature (Th) of more than 250°C up to more than 400°C and show isochores that are concentrated within the 1-2 kbar pressure range (Fig. 3). The temperature was at least around 400°C (around or slightly above the maximal Th) and maximal around 450 to 500"C, as no biotite but chlorite is stable. Petrographic chronology as well as the generally higher Th of Lc-w type inclusions argue for trapping of Lc-w inclusions possibly before Lw-c inclusions. Temperatures around 400°C are in accordance with those generally assumed for greisen mineralization. The greisen assemblage and the fluid inclusion data suggest a hydrothermal system, initiated by the leucogranite intrusion. Starting with circulation of fluids of the Lc(-w) type, with decreasing tempera-
4 P-T-X EVOLUTION A reconstruction of the P-T changes has been carried out based on calculated fluid inclusion isochores (according to Bakker 1999 and Zhang & Frantz 1987) and the consideration of the mineral assemblages. Calculated isochores of the investigated fluids are shown in Fig. 3. 4.1 Aqueousjluids LWI The late-Hercynian postmagmatic evolution starts with subsolidus epidote, whose calculated isochores of primary fluid inclusions (FI) indicate relatively high pressure during the subsolidus stage (-4 kbar at 600°C; Fig. 3). Some aqueous inclusion (FIP) within quartz present similar characteristics as the primary inclusions in epidote, and are therefore presumably also correlated with the formation of Ca-A1 707
to about 250°C, within the calc-alkaline rocks a hot geothermal circulation was induced by the emplacement of the younger leucogranite or by abnormal heat flows in relation to the delamination of the chain. Greisen zones with quartz-phengite- tourmaline assemblages formed locally. This stage 2 is marked by the percolation of carbonic fluids, characterized by rather high temperatures at a high structural level. During this stage, temperatures around 340°C h30"C are still rather high for the inferred depths and imply abnormal heat flows and thermal gradients of around 60-80"Ckm. As most fluids were trapped in thin micro-fractures in rocks, fluids are likely to be in thermal equilibrium with their host rocks. Such conditions are typically those one may observe in the proximity of geothermal fields, related to concealed granite intrusions. Postdating hot geothermal circulation induced by the leucogranite intrusion, decreasingly hot circulations (<400"C) of aqueous, partly N2-bearing fluids lead to multiphase pervasive alterations. This reflects the progressive decrease of the thermal effects linked to the peraluminous magmatism.
ture and decreasing pressure in the course of leucogranite cooling, the fluids get less dense and richer in water. In this later stage Lw-c inclusions are formed. 4.3 Nitrogen and aqueous fluids Lw-n, Vn and LWII Pervasive multiphase hydrothermal alteration of stage 3 and 4 occurred after greisen formation. Nzbearing as well as aqueous fluid inclusions may be correlated with this alteration stage. 4.3.1 Nitrogenfluids Lw-n and Vn Nitrogen fluids occur either as 2-phase aqueous inclusions with relatively low salinity and with high Th around 380°C or as monophase vapor inclusions. Assuming a relationship between vapor and liquid N2-bearing inclusions, the pressure range can be restricted to <1 kbar at 450-500°C (Fig. 3). These Lw-n and Vn inclusions are probably related to high-temperature hydrothermal alteration of primary rock forming minerals as well as the alteration of Ca-A1 silicates (stage 3), postdating the carbonic fluid event. Because of their high Thybut comparably low P, these fluid inclusions indicate a relatively fast pressure decrease after the greisen alteration stage. Results of chlorite thermometry (Cathelineau 1988) show a large range from -400 to 220"C, which may reflect the successive temperature decrease in the course of the pervasive alteration stage.
REFERENCES Bakker R.J. 1999. Adaptation of the Bowers and Helgeson (1983) equation of state to the H20-C02-CH4-N2-NaC1 system. Chem. Geol., 154: 225-236. Bertrand, J.M., J. Leterrier & E. Delaperrikre 1996. GCochronologie U-Pb de granitoides du Confolentais, de VendCe et des forages de Charroux-Civray. 17th. Rkunion des Sciences de la Terre, SOCGkol France: 74. Cathelineau, M. 1988. Cation site occupancy in chlorites and illites as a function of temperature. Clay Minerals, 23: 471485. Cathelineau, M., M. Cuney, M.C. Boiron, Y. Coulibaly & M. Ayt Ougougdal 1999. PalCopercolations et paltointeractions fluideshoches dans les plutonites de CharrouxCivray. Actes des Journkes Scientifiques CNRS/ANDRA: 159-179. Cuney, M., M. Brouand, J.M. Stussi & C. Gagny 1999. Le massif de Charroux-Civray (Vienne): un exemple caractkristique des premikres manifestations plutoniques de la chaine hercynienne. Actes des Journkes ScientiJiques CNRS/ANDRA: 63- 104. Duthou, J.L., J.M. Cantagrel, J. Didier & Y. Vialette 1984. Paleozoic granitoids from the French Massif Central: age and origin studied by 87Rb-87Srsystem. Physic Earth Planet Sci Int, 35 : 131- 144. Freiberger, R., Hecht, L., Cuney, M. & Morteani, M. in press. Secondary Ca-AI silicates in plutonic rocks: Implications for their cooling history. Contrib Mineral Petrol. Frey, M., C. de Capitani & J.G. Liou 1991. A new petrogenetic grid for low-grade metabasites. J. Metamorph. Geol., 9: 497-509. Virlogeux, D., J. Roux & D. Guillemot 1999. Apport de la gCophysique a la connaissance geologique du massif de Charroux-Civray et du socle poitevin. Actes des Journkes ScientiJiques CNRS/ ANDRA: 33-6 1. Zhang, Y.G. & J.D. Frantz 1987. Determination of the homogenization temperatures and densities of supercritical fluids in the system NaCI-KC1-CaCl2-H20 using synthetic fluid inclusions. Chem. Geol., 64: 335-350.
4.3.2 Aqueous fluids LWII Secondary aqueous fluid inclusions within quartz (LWII)present higher Th (around 380°C), but distinctly less steep isochores (Fig. 3) than LWIconnected with Ca-A1 silicate formation. Furthermore, aqueous LWIIinclusions have distinctly lower salinities than LWI inclusions. Both, different Th and salinity indicate the dissimilar origin of LWI and LWII.The inclusions of LWIItype may be correlated with inclusions containing NZ in their vapor phase (Lw-n, Vn), due to petrographic observations. 5 CONCLUSION The preliminary study of fluid inclusions allows a reconstruction of the P-T-X evolution of the Hercynian basement in the Charroux-Civray area since the crystallization of the main calc-alkaline plutonic bodies. Cooling of the calc-alkaline series starts at solidus temperatures at about 4 kbar as indicated by the data of primary aqueous fluid inclusions in subsolidus epidote, which is in accordance with the stability of magmatic epidote and hornblende barometry. The post-magmatic evolution continues under slightly decreasing pressure (uplift of the basement) down to 2-3 kbar at 200-280°C (prehnitepumpellyite paragenesis). Thus, after cooling down 708
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Experimental study on clinoptilolite and rnordenite crystallization M .R.Ghiara Dipartimento di Scienza dellu Terra, Universitd Federico II, Napoli, Italy
C .Petti Centro Musei delle Scienze Naturali Universitd Federico II, Napoli, Italy
R.Lonis Progeinisa S.p.a., Societd Sarda Valorozzuzione Georisorse, Cugliari, Italy
ABSTRACT: Experiments were carried out in a closed hydrothermal system at 200°C using as reactants unaltered rhyolitic glass from ignimbrites and deioniozed water and 0.01, 0.03, 0.05 and 0.1 N KOH and NaOH solutions. Newly formed minerals are: smectite, clinoptilolite and mordenite. The clinoptilolite is Ktype both in KOH and NaOH experimental runs and its chemical composition is comparable to that of clinoptilolite occurring in the pyroclastic flows of Sardinia. Indeed, mordenite occurring only in NaOH experimental runs, shows low Si content and high K amount respect to natural one. It to be noted that the early clinoptilolite laths occur, in the NaOH experimental runs. The formation of different minerals depends on pH-value and Na+/H+and K'/H+ ratios of the reacting solution.
1000 mg of glass and 10 ml of fluid were sealed in a bronze Teflon-lined autoclaves and placed in a thermostat-controlled oven at 200°C and autogenous pressures. At arbitrary intervals, the pressure vessel was quenched to room temperature and the solid and solution was separated by filtration. The newly formed minerals were identified by means of X-ray powder diffraction, scanning electron micrograph (SEM) and microprobe analyses. pH-measurements were carried out at 25°C by means of pHM80 pHmeter (uncertainty lower than +2%).
1 INTRODUCTION The voluminous Oligo-Miocene pyroclastic flows of Sardinia, emplaced in volcano-tectonic depression in subaereal, lacustrine and marine environments, are very useful to investigate the fluidhock interaction processes. Recent studies (Ghiara et al. 1997, 1999, 2000, Morbidelli et al. 1999) on subaerial pyroclastic flows of central-northern Sardinia, showed more or less advanced alteration of the glassy component (shards) mainly forming clinoptilolite and mordenite and, to a minor extent smectite and opal-CT. Minerogenetic models for most zeolites from pyroclastic flows require interaction processes between cognate entrapped fluids and glassy components. The present study was undertaken to define the crystallization relations between clinoptilolite and mordenite during alteration processes of rhyolitic glass.
3 RESULTS Smectite crystallized in experiments with deionized water after 42 days. It was formed also in 0.01N KOH and NaOH runs after 42 and 7 days respectively (Tables 1, 2). Clinoptilolite crystallized in 0.05 and O.1N KOH experimental runs after 7 days, whereas it occurred after 49 days in 0.03 KOH experimental runs. Clinoptilolite and mordenite crystallized in 0.05 and 0.1N NaOH experimental runs after 7 and 14 days whereas in 0.03N NaOH run occurred after 49 and 56 days respectively (Table 2). In Table 3 atomic ratios of clinoptilolite (calculated on the basis of 72 oxygens) and mordenite (on the basis of 96 oxygens) formed in KOH and NaOH
2 MATERIALS AND METHODS The fluid/glass interaction process has been carried out in a closed system using unaltered rhyolitic glass from ignimbrites and deionized water (run l), KOH (runs 2, 3,4, 5) and NaOH solutions (runs 6, 7, 8,9) of different concentrations (0.01 - 0.03 - 0.05 0.1N). The glass was crushed to less then 150 pm; 709
Table 3. Unit-cell composition of synthetic clinoptilolite (cp) and mordenite (mo). The unit formula is calculated on the basis of 72 and 96 oxygens respectively.
Table 1 . Secondary minerals crystallized in KOH runs.
Runs Days 7 14 21 28 35
42 49 56
2
(sm) (sm) (sm)
3
4
5
--
CP CP CP CP CP CP CP cp
CP CP CP CP CP CP CP CP
(CP) CP
r u n s 4 cp Si 29.36 A1 6.48 Fe3+ 0.20 Mg 0.15 Ca 0.46 Na 2.23 K 3.07 Si/Al 4.53 E% 2.20
Run 2: 0.01N; run 3: 0.03N; run 4: 0.05N; run 5 0.1N; sm: smectite; cp: clinoptilolite; (): traces of a mineral; - absent.
runs are given. The chemical composition was obtained by averaging a total of 10 microprobe point-analyses for each sample. In Table 4 the atomic ratios of representative clinoptilolites and mordenites from pyroclastic flows of Sardinia are shown. Clinoptilolites crystallized both from KOH and NaOH runs are K-clinoptilolites and their chemical compositition is comparable to that of natural ones even if these latter have, generally, lower K content. The synthetic mordenites are characterized by low Si and high K content respect to natural ones. Well-formed clinoptilolite laths crystallized in KOH and NaOH runs were observed by means scanning electron microscopy (Fig. 1). They are very similar to natural ones. Rods of mordenite were observed in NaOH runs (Fig. 2), whereas natural mordenite showed generally filiform or spider-web fashion. In KOH runs (4 and 5), after 14 days of interaction SEM images showed the crystallization of a rosette aggregate that took place throughout the reaction time (Fig. 1). Qualitative analyses (EDS) on this unidentified phase indicated a Ca-rich silicatic composition.
7 14 21 28 35 42 49 56
6 (sm) (sm) (sm) (sm) (sm) (sm) (sm) (sm)
7
-
(cp) cp (mo)
8
8 cp 28.21 7.09 0.82 0.91 0.04 0.26 5.29 3.98 6.02
8 mo 38.56 8.94 0.62 0.07 0.11 5.07 3.62 4.31 5.49
9 cp 29.33 6.29 0.45 0.33 0.19 0.74 4.66 4.66 4.56
9 mo 37.79 8.30 0.10 0.52 1.22 3.04 3.32 4.55 9.15
E% total balance error from Passaglia (1970). Run 4: clinoptilolite crystallized after 49 days of reaction; run 5: after 56 days; run 8: after 42 days; run 9: after 49 days.
In Figure 3 the pH trending in glass/NaOH 0.01N and 0.1N runs are shown. The pH-values of 0.03 and 0.05N experimental runs showed similar trends with intermediate pH-value. For all glassKOH and glass/NaOH experimental runs a sudden decrease of pH values was observed in the first hours of reaction and then it stabilized around a constant value (Fig. 3). An opposite trend was showed by pH of glasddeionized water system. The pH-value abruptly increase after 7 days from 6.3 to 7.8, it is likely related to the consumption of H+ (Berger et al. 1988, Gislason & Eugster, 1987). Finally also in KOH - NaOH 0.01N runs the pHvalue stabilizes around 8 value after 7 days (Figure 3).
Table 4. Unit-cell composition of natural clinoptilolite and mordenite from pyroclastic flows of Sardinia, Italy (Ghiara et al. 1997).
Table 2. Secondary minerals crystallized in NaOH runs.
Runs
5 cp 29.20 6.39 0.45 0.37 0.07 1.14 4.68 4.57 2.05
9 Si A1 Fe3+ Mg Ca Na K %/A1 E%
CP (CP) CP, (mo) (CPh (mo) cp,mo cp,mo cp, mo cp,mo cp,mo cp,mo cp, mo cp,mo cp,mo cp,mo cp,mo cp, mo
Run 6: 0.01N; run 7: 0.03N; run 8: 0.05N; run 9: 0.1N; sm: smectite; cp: clinoptilolite; mo: mordenite; (): traces of a mineral; - absent.
IS-cu 29.91 6.16 0.02 0.58 0.78 0.78 2.28 4.85 4.58
Na-cP 30.04 6.06 0.02 0.66 1.03 0.70 1.03 4.96 4.95
Ca-cP 29.67 6.49
0.09 0.96 3.09 0.65 4.57 4.57
mo 40.27 7.70 0.20 0.09 1.87 2.82 0.47 5.23 8.71
mo 40.43 7.75 0.02 0.19 2.28 1.06 0.91 5.22 9.86
K-cp: K-clinoptilolite; Na-cp: Na-clinoptilolite; Ca-Cp: Caclinoptilolite; mo: mordenite.
710
Figure 1. Rosette aggregates and clinoptilolite laths formed in glass/KOH 0.1N run after 28 days of reaction.
Figure 2. Mordenite and clinoptilolite formed after 28 days of reaction in glass/NaOH 0.05N run.
in good agreement with the pH value of deionized water system and 0.01N KOH and NaOH runs (Fig. 3), the [Al(0H)6l3- complex is likely formed, favouring the formation of clay minerals. Instead, the [A1(OH)4]- complex is formed in 0.03, 0.05 and 0.1N alkaline solutions (Fig. 3), favouring the crystallization of zeolites (Merino et al. 1989, Ghiara et al. 1993, Ghiara & Petti, 1996). In these latter the lack of smectite and the early crystallization of clinoptilolite preclude the Si increase into the interacting solutions (Reynold & Anderson 1967). As consequence, the mordenite shows a lower %/A1 ratio then the natural one (Table 4) and silica minerals (opal-CT and/or newly quartz grains) formation is precluded. In closed-system experiments carried out by Barth-Wirshing & Holler 1989, using rhyolitic glass from Lake Tecopa, Inyo County, California (Table 5), phillipsite and analcime crystallized from 0.0 1, 0.05 and 0.1N NaOH runs at 200 "C whereas alkali feldspar and quartz formed with 0.01, 0.05 and 0.1N KOH solutions. The crystallization of different minerals pairs, being starting glass composition very similar (Table 5), is likely due to different experimental conditions (i.e. solid/liquid ratio). Another important factor that influences the crystallization of authigenic minerals is linked to chemical composition of interacting fluids. The experimental results show that different kinds of zeolites are constrained by Na+/K' and K " / P ratios of reacting solution. Finally, it is important to note that the K amount, leached during the interaction process, is sufficient for early c~inoptilo~ite crystallization, whereas the formation of mordenite requires solutions dominated by sodium.
Table 5. Unit-cell composition of rhyolitic glasses. The unit formula is calculated on the basis of 80 oxygens.
Si Ti A1 Fe Mg Ca Na K OH Si/A1
Figure 3. pH-value vs. reaction time in NaOH run.
4 DISCUSSION AND CONCLUSION Experimental data indicate that the pH of interacting solutions plays an important role in the crystallization of secondary minerals. In particular,
thiswork Barth-Wirshing
711
1 32.78 0.02 6.98 0.52 0.18 0.5 1 2.16 2.59 18.21 4.69 Holler 989.
2 33.14 0.04 6.38 0.49 0.15 0.27 2.67 3.08 11.18 5.19
REFERENCES Barth-Wirsching, U. & H. Holler 1989. Experimental studies on zeolite formation conditions. Eur. J Mineral. 1: 489506. Berger, G., Schott, J. & M. Loubet 1988. Fundamental processes controlling the first stage alteration of basaltic glass by seawater: An experimental study between 200300°C. Earth Planet. Sci. Lett. 84: 43 1-445. Ghiara, M.R., Franco, E., Petti, C., Stanzione D. & G.M. Valentino 1993. Hydrothermal interaction between basaltic glass, deionized water and seawater. Chemical Geology 104: 125-138. 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 1-229. Ghiara, M.R. & C. Petti, 1996. Chemical alteration of volcanic glasses and related control by secondary minerals: experimental studies. Aquatic Geochemistry, 1: 329-354 Ghiara, M.R., Petti, C., Franco, E. & R. Lonis 2000. Distribution and genesis of zeolites in Tertiary ignimbrites of Sardinia: evidence of superimposed mineralogenic processes. In: C. Colella & F.A. Mumpton (eds), Natural Zeolite for the third Millenium: 177 - 192. Ghiara, M.R., Petti, C., Franco, E., Lonis, R., Luxoro, S. & L. Gnazzo 1999. Occurrence of clinoptilolite and mordenite in Tertiary calc-alkaline pyroclastites from Sardinia (Italy). Clays and Clay Minerals. 47(3): 3 19-328. Gislason, S.R. & H.P. Eugster 1987. Meteoric water-basalt interactions, I. A laboratory study. Geochim. Cosmochim. Acta. 5 1 : 2827-2840 Merino, E., Colin, H. & H. H. Murray 1989: Aqueous-chemical control of the tetrahedral-aluminium content of quartz, halloysite and other low-temperature silicates. Clays and Clay Minerals., 37: 135-142. Morbidelli, P., Ghiara, M.R., Lonis, R. & A Sau 1999. Zeolitic occurrence from Tertiary pyroclastic flows and related epiclastic deposits outcropping in northern Sardinia (Italy). Per. Minerl., 68 (3): 287-313. Passaglia E. 1970. The crystal chemistry of chabazites. Amer. Miner. 55: 1278-1301. Reynolds, R. C. Jr. & D. M. Anderson 1967. Cristobalite and clinoptilolite in bentonite beds of the Colville Group, northern Alaska. J Sed Petrol. 37: 966-969.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger,Lisse, ISBN 90 2651 824 2
Ore fluid, of late Mesozoic porphyry-epithermal gold-copper system i e a s t China Renmin Hua, Xiaofeng Li, Jianjun Lu, Peirong Chen & Xiaodong Liu State Key Lib. of Mineral Deposit Research, Dept. of Earth Sciences, Nanjing University, Nanjing, China
ABSTRACT: Extensive acidic-intermediate magmatism took place in East China in late Mesozoic to form a belt of volcanic-intrusive rocks. This resulted in well-developed gold-copper deposits of epithermal, mesothermal and porphyry types, constituting a Mesozoic porphyry-epithermal metallogenic system. Ore fluids of several typical deposits are studied and reviewed in present paper. Epithermal gold (copper) deposits of Tuanjiegou and Zijinshan, representing low- and high-sulphidation types respectively, are characterized by meteoric-source fluid circulation. Ore fluid of mesothermal gold-copper deposit in Yinshan is also dominated by meteoric water. Even in Tongchang porphyry copper system, meteoric water played an important role during the main stage of copper deposition. deposit in Jiangxi province, Zijinshan gold-copper deposit in Fujian province, and many others in Jilin, Shandong and Zhejiang provinces. In the present paper, three typical deposits (Tuanjiegou, Yinshan, and Zijinshan) are studied and reviewed with respects to their major characteristics, especially the ore-forming fluids, based on many previous researches.
1 INTRODUCTION Cenozoic gold andor copper deposits of porphyry, mesothermal and epithermal types are well developed in southwestern Pacific volcanic islandarc belt, which are genetically related to the subduction of Pacific and Philippine plates, and are generally defined as the porphyry-epithermal goldcopper system (Corbett & Leach 1995). In East China, acidic to intermediate volcanic, sub-volcanic and intrusive rocks of late Mesozoic ages are widely distributed (Wang et al. 1996), and are also closely related with extensive gold and copper mineralizations. This late Mesozoic volcanicintrusive event was generally explained as a result of subduction of paleo-Pacific plate towards the Eurasia plate (Ren 1997). In fact, the East China, together with the Russian Fareast, is considered as the outer zone of western part of circum-Pacific magmatic-metallogenic belt, while the Cenozoic island-arc chain as its inner zone. It is recently summarized that the extensive gold-copper mineralization was a product of so-called "Yanshanian metallogenic explosion" in East China (Hua 1999). Gold andor copper deposits in late Mesozoic volcanic-intrusive belt of East China can also be attributed to the porphyry-epithermal system. Many deposits have been well studied, such as Tuanjiegou gold deposit in Heilongjiang province, Tongchang porphyry copper deposit and Yinshan polymetallic
2 TUANJIEGOU GOLD DEPOSIT Tuanjiegou gold deposit is located in the margin of a Proterozoic basement near a Mesozoic depression in eastern Heilongjiang province. Spatially the deposit is closely related to Cretaceous granodiorite porphyry. There are two main models concerning the ore genesis: one is the porphyry type (Wu 1984, Chen & Zhu 1993), and the other is the epithermal type (Chen et al. 1992, Ren et al. 1993). Several important aspects can be recognized as the main factors affecting the ore genesis and ore fluid origin: 1. The gold ore bodies occurred in the structural fracture zone within or close to the contact between porphyry and basal metamorphic rock, but not within the porphyry. After condensation of granodiorite porphyry the ore-hosting fracture zone was evolved from early compression-shearing to late tension. 2. Hydrothermal breccia is very common in this 713
some authors preferred a magmatic source for the ore fluid, (Shen et al. 1991, He et al. 1992, Wu 1998). While some others suggested a meteoric source (Hua et al. 1995, Zhang et al. 1996). The fissure-filling subvolcanic dikes are of very small size both at surface and depth. The volatile system seems too limited for supporting such a large-scale mineralization. As a matter of fact, all ore veins occur in the wall rocks but not in the dikes. Geochemical investigation by the present authors showed that a gold depleted zone (from background value of about 4 ppb to less than 1 ppb) existed in an area larger than 220 km2 around the Yinshan deposit. This indicates the extensive fluid movement during mineralizing process in this area, which was not possible by magma derived fluid. To explain the oxygen isotope data (mostly in area between meteoric line and magmatic domain), Hua et al. (1995) studied the process of water-rock interaction in Yinshan deposit, and proved that metaoric water was the major source of ore fluid. Although ore fluid of Yinshan deposit is mainly of meteoric source, which is similar to Tuanjiegou gold deposit, some distinct differences between these two deposits occur. Tuanjiegou deposit originated at a very shallow depth, similar to the hot-spring type deposit, such as the McLaughlin gold deposit in California (Lehrman 1986). The depth of mineralization of Yinshan, however, was obviously much greater. The drill holes show that copper mineralization still exist at about 1500 m below surface. On the other hand, the temperature and salinity of ore fluid in Yinshan is also higher than that in Tuanjiegou. Substantially, mineralization in Yinshan has more close relation with magmatism than Tuanjiegou, and magmatic water did play some role in the early stage of alteration-mineralizing process. It is suggested that Yinshan deposit can be referred to mesothermal type. About 20 km to the northeast of Yinshan is the Tongchang porphyry copper deposit. The host rock granodiorite porphyry is considered as the deeper equivalent of dacite porphyry in Yinshan, owing to the similarity in geology, mineralogy, geochemistry and ages (Hua & Dong 1984), hence, the two deposits can be attributed to the same metallogenic system. However, no epithermal deposit was found in that area.
deposit, revealing the existence of a hydrothermal explosion at shallow depth. 3. Quartz and adularia are the main alteration minerals of the deposit, and the ore-body is also located in this alteration zone. Low-temperature chalcedony occurs in association with the gold ore. 4. The main sulfide mineral associated with gold is marcasite rather than pyrite. Colloidal and framboidal pyrite, and pyritized fossil algae are also widely developed. 5 . Mineralizing processes include two stages, the most important being the low-temperature stage. Fluid inclusion studies show that the homogenization temperatures of this stage range from 230' to 9OoC. 6. Data from different sources show relatively low salinity of the ore fluid (
3 YINSHAN POLYMETALLIC DEPOSIT The Yinshan deposit is a large deposit of gold, copper, silver, lead and zinc. It is genetically related with a Mesozoic dacitic-rhyolitic caldera and subvolcanic dikes. Therefore, it has been clarified as a subvolcanic hydrothermal deposit (Ye 1983, Hua 1987, Hao 1988). According to the close spatial relationship between ore and subvolcanic system, as well as the data of oxygen isotope compositions,
4 ZIJINSHAN GOLD-COPPER DEPOSIT
Because of the development of typical quartz-alunite alteration and other characteristics, Zijinshan was considered as the first example of quartz-alunite 714
type epithermal copper-gold deposit in mainland China (Zhang et al.1991). In fact, it is characterized by the zoning of gold at top and copper below. Gold ore occur only in the shallowest zone associated with strong silicification, whereas the underneath copper ore bodies occur as veins in the quartzalunite alteration zone. Besides, there are Zhongliao porphyry copper deposit and Wuzhiqilong mesothermal deposit in the vicinity area, to form a porphyry-epithermal system. Zijinshan deposit belongs to the high sulfidation type of epithermal series. Ore fluid of this type includes two portions: magma-derived gases (Cl, C02, H2S, SO2, etc) and meteoric waters where these gases dissolved when they rose to shallow depth. The input of gases makes the meteoric water more acidic and causes strong acidic alteration of the wall rock, resulting in the formation of a special mineral assemblage (alunite, dickite, pyrophyllite etc), commonly called advanced argillic alteration. This alteration occurs in the Zijinshan deposit, as well as in the Tertiary Chinkuashih deposit in Taiwan (Tan 1991) and many other high-sufidation epithermal gold-copper deposits in the southwest Pacific region. Although a magmatic input has been recognized, the most important source of the ore fluid is not magmatic. Studies by Zhang et al. (1992) and Hua & Hu (1998) show that the circulation of heated meteoric water dominated the ore-forming process. Fluid inclusion studies show that although the temperature and salinity variation of ore fluid has wider range as compared to the Tuanjiegou gold deposit, the fluid carrying the shallow gold mineralization is of low temperature and salinity. For instance, Zhang et al. (1992) measured the salinity of fluid inclusion in the range from 0% to 2l.6%, mostly 4% to 8% (NaC1 equiv.). Data from Wei & Tao (1998) show fluid salinity for gold mineralization is 0.1% to 4.1%. The present authors determined the salinity of fluid inclusion between 0.2% and 12.8%. All these data basically reflect the feature of a meteoric-sourced ore fluid.
Yinshan polymetallic deposit is of mesothermal type, a transitional type of epithermal and porphyry deposits. The Naozhi gold (copper) deposit in Jilin and Wuzhiqilong copper deposit in Fujian can also be attributed to this group. The typical porphyry copper deposits include Tongchang near Yinshan in Jiangxi, Zhongliao near Zijinshan in Fujian, and Xiaoxinancha near Wufeng in Jilin. All these deposits developed in the Mesozoic volcanicintrusive belt of East China. They form a late Mesozoic porphyry-epithermal gold-copper system. 2. Ore fluids for deposits of Mesozoic porphyryepithermal gold-copper system in East China were dominated by meteoric waters, and magma-derived gases to variable extent. Ore fluid of the Tuanjiegou deposit is substantially meteoric water. In the Zijinshan deposit, meteoric water caused the silicification and gold mineralization in a relatively late stage and at near-surface zones, while some amount of magmatic volatile, was dissolved in down-going meteoric water. As to the porphyry endmember of a porphyry-epithermal system, although magmatic water played an important role in early mineralizing period, e.g. potassic alteration, copper deposition associated with sericite and argillic alteration was in a relatively lower temperature (<4OO0C) and meteoric-dominated environment (Sheppard et al. 1971, Beane 1983). This feature is also quite distinct in Tongchang porphyry copper deposit (Jin et al. 1999, Hua et al. 2000). Ore fluid of mesothermal mineralizations, exemplified by the Yinshan deposit, also exhibit meteoric characteristics after water-rock interaction was considered. ACKNOWLEDGEMENTS This study was supported by grant 49733 120 of China National Natural Science Foundation and a Major State Research Program (No. G 1999043209)of PRC.
REFERENCES Beane, R.E. 1983. The magmatic-meteoric transition. Geothermal Resources Council, Special Report. 1 3: 145-
5 CONCLUSIONS
253 Chen, Y. & Y. Zhu (eds), 1993. Mineral Deposit Models of China. Beijing: Geol. Pub. House. Chen, Z., Q. Ren, C. Zhang, J. Qiu. & W. Zhang 1992. Discovery of pyritized fossil algae in ores from Tuanjiegou gold deposit and its geologic significance. ,I. Nanjing Univ. (Earth Sci). 4(2): 77-80. Chen, Z., Zhang, L., Wang, K., & J. Liu 1993. Oxygen (hydrogen) isotope compositions of the Tuanj iegou gold deposit in relation to the ore-forming process. Mineral
1. The main features, ore geneses, and ore fluids of several gold-copper deposits are listed in this paper. It is concluded that the Tuanjiegou gold deposit is a typical low-sulfidation epithermal deposit; similar deposits include Wufeng and Ciweigou in the Yanbian area.The Zij inshan gold-copper deposit is a high-sulfidation epithermal deposit, very similar to the Chinkuashih gold-copper deposit in Taiwan. The
Deposits.l2:174-181.
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Corbet, G.E. & T.M. Leach 1995. SW Pacific gold-copper systems. A workshop presented at the Pacrim Conference Auckland, NZ. Hao, Z. 1988. Metallogenic zoning of the Yinshan polymetallic deposit, Jiangxi province. Mineral Deposits. 7(3): 3-1 4. Hayba, D.O., Bethke, P.M., Heald, P. & N. K. Foley 1985 Geologic, mineralogic and geochemical charateristics of volcanic-hosted epithermal precious metal deposits. In: Berger & Bethke (eds), Geology and Geochemistry of Epithermal Systems, Reviews in Econ. Geol., 2: 129-162. He, G. & D. Lin 1992. Study on stable isotope composition of Yinshan deposit at Dexing, Jiangxi province. Mineral Resource and Geol. 6(5): 406-41 1. Hedenquist, J. W. 1987. Mineralisation associated with volcanic related hydrothermal systems in the Circum Pacific Basin. In: Horn (ed), Transactions 4'h Circum Pacific Conference, Singapore, AAPG. 513-523. Hua, R. & Z. Dong 1984. The characteristics and origins of granitic rocks of two genetic series in Dexing, Jiangxi. In: Xu & Tu (eds), Geology of Granites and Their Metallogenetic Relations. Beijing: Science Press. 347-366. Hua, R. 1987. A discussion on the mechanism of lead, zinc and copper metallogeneses in Yinshan, Jiangxi province. Mineral Deposits. 6(2): 90-96. Hua, R., Wu P. & K. Chen 1995. Two-stage water-rock interaction indicated by variation of oxygen isotope compositions of altered rocks in the Yinshan polymetallic deposit, Jiangxi, China. In: Kharaka & Chudaev (eds), Water-Rock Interaction: 187-190. Rotterdam: Balkema. Hua, R. & J. Hu 1998. Fluid migration-reaction model of Zijinshan epithermal deposit as traced by variation of oxygen isotope compositions of altered wall rocks. In: Arehart & Hulston (eds), Water-Rock Interaction: 549-552. Rotterdam: Balkema. Hua, R. 1999. The Yanshanian metallogenic explosion in East China. J. Geosci. China. l(1): 4-12. Hua, R., Li, X., Lu, J., Chen, P., Qiu, D. & G. Wang 2000. Study on the tectonic setting and ore-forming fluids of Dexing large ore-concentrating area, northeast Jiangxi province. Advance in Earth Sci. 15: 525-533. Jin, Z. & J. Zhu 1998. The source of ore-forming material in the Dexing porphyry copper deposits, Jiangxi. Geol. Review. 44(5): 464-469. Lehrman, N.J. 1986. The McLaughlin mine, Napa and Yolo Counties, California. In: Tingley & Bonham (eds), Precious Metal Mineralization in Hot Spring Systems, NevadaCalifornia. NBMG Report 41: 85-89. Ren, J. 1997. Tectonic framework and geodynamic evolution of eastern China. Geoscience Research. 29: 43-55. Ren, Q., Qiu, J., Wang, D., Zhang, C., Xie, X., Yang, R., Xu, Z. & Z. Chen 1993. Four types epithermal gold deposits in Mesozoic volcanic areas of eastern China. Resource Geol. Special Issue. 16: 307-3 1 3. Rui, Z., Zhang, H., Wang, L., Chen, R., Jin, B., Jin, F., Wan, Y., Zhou, Y. & Q. Meng 1995. Porphyry-epithermal coppergold deposits in Yanbian area, Jilin province. Mineral Deposits. 14: 99-126. Shen, W., Chen, F. & C. Liu 1991. A study on stable isotope feature of Yinshan multi-metallic deposit, Jiangxi Province. J. Nanjing Univ. (Earth Science). 2: 186-194 Sheppard, S.M.F., Nielson, R. & H.P. Taylor 1971. Hydrogen and oxygen isotope ratios in minerals from porphyry copper deposits. Econ. Geol. 6 6 5 15-542 Tan, L.P. 1991. The Chinkuashih gold-copper deposits, Taiwan. Soc. Econ. Geol. Newslett. 7:22-24
Wang, D., Ren, Q., Qiu, J., Chen, K., Xu, Z. & J. Zen 1996. Characteristics of volcanic rocks in the shoshonite province, Eastern China, and their metallogenesis. Acta Geol. Sinica. 70(1):23-34 Wei, J. & G. Tao 1998. Ore fluid and metallogenic model of Zijinshan copper-gold deposit. Mineral Deposits. 1 SUP.): 1095-1098. Wu, S. 1984. The framboidal marcasites in Tuanjiegou porphyry gold deposit, Heilongliang province. Minerals & Rocks. 4(4): 20-25. Ye. Q. 1983. A preliminary study on hypogene zoning of the Yinshan copper-lead-zinc deposit. Bull. Nanjing Inst. Geol. M. R. Chinese Acad. Geol. Sci.4(1): 1-16 Zhang, D., Li, D. & Y. Zhao 1991. Zijinshan deposit: the first example of quartz-alunite type epithermal Cu-Au deposit in mainland China. Geol. Review. 37:48 1-491. Zhang, D., Li, D. & Y. Zhao 1992. Alteration and Mineralization Zoning of Zijinshan Cu-Au Deposit. Beijing: Geol. Pub. House. Zhang, L., Liu, J., Chen, Z. & G. Yu 1996. Hydrogen and oxygen evolution for water-rock system in super-huge Tongchang copper deposit, Jiangxi province. Sci. Geol. Sinica. 3 1: 250-263.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
The characteristics and genesis of the kaolinite-bearing gold-rich Nurukawa Kuroko deposit, Aomori Prefecture, Japan D. Ishiyama Akita University, Akita, Japan
K.Hirose Earth Remote Sensing Data Analysis Center, Tokyo, Japan
T.Mizuta, 0 .Matsubaya & Y.Ishikawa Akita University, Akita, Japan
ABSTRACT: The Nurukawa deposit is a kaolinite-bearing gold-rich Kuroko deposit. The formation of Aubearing ores of the deposit is earlier than the formation of black ores. The fluid forming Au-bearing ores was acidic fluid, high saline, and depleted in deuterium. These characteristics suggest that the Au mineralization was caused by a system in which magmatic water derived from dacitic volcanism underlying the deposit mixed with a circulating hydrothermal solution of seawater origin. On the other hand, the fluid responsible for the formation of black ores is dominantly of seawater origin.
1 INTRODUCTION The Nurukawa Kuroko deposit is located in the northeast margin of the Hokuroku district, northeastern Japan, an area in which many Kuroko deposits are distributed (Fig. 1). The average Au content of typical Kuroko deposits in the Hokuroku district is 1.3 ghon (Tanimura et al. 1983). The average Au content of the Nurukawa deposit is 6.8 g/ton (Yamada et al. 1988) that is much higher than typical Kuroko deposits. In this paper, we describe and compare the mode of occurrence of ores and the geochemical characteristics of Au and Pb-Zn mineralization of the Nurukawa deposit. 2 OUTLINE OF GEOLOGY AND ORE DEPOSITS The Nurukawa Kuroko deposit was formed in Middle Miocene time (12.5-10.7 Ma; MMAJ 1988). The Nurukawa deposit consists of five orebodies, Nos. 1 to 5 (Fig. 1). The largest of them is No. 5 orebody. It lies in the uppermost part of acidic tuff breccia of the Lower-Hayasemori Formation. It is covered at the top by pumice tuff of the Upper-Hayasemori Formation (Nishitani et al. 1986). Some dacitic crypto-lavadomes, intruded into the acidic tuff breccia of the Lower-Hayasemori Formation, can be found under No. 5 orebody (Figs. 2 & 3 ; Yamada 1988). No. 5 orebody is composed of an Au-bearing stockwork siliceous orebody, an Au-bearing bedded siliceous orebody, and a massive base-metal orebody in ascending order. The Au-bearing stockwork sili-
Figure 1. Location and geologic maps of the Nurukawa mining area (Nishitani et al. 1986). Location of orebodies is pro.jetted to the surface. I : Volcanic ash CTowada Volcanics), 2: basalt tuff (l’owada Volcanics), 3: ‘ruff breccia (Tobe F.), 4: Pumice tuff (U. Hayasemori F.), 5: Mucktone (IJ. Habasemori F.), 6: lapilli tuff (L. Hayasemori F.), 7: Mudstone (L. Hayasemori F.), 8: Andesitc tuff (Nabekurdsawa I-.), 9: PreTertiary sdmentao rocks, 10: Dacitic intrusive rochs, I I : Ore deposi 1s
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Figure 2. North (N) - South (S) cross section through No. 5 orebody of the Nurukawa &posit (Yanach 1988). N and S correspond to N and S in Fig. I . The enlargement of the redangular in Fig. 2 is shown in Fig. 3. I : Volcanic ash (Towada Volcanics), 2: basalt tuff Cl’owach Volcanics), 3: Mudstone & sandstone (Ikarigaseki F.), 4: ‘I’iiff breccia (Tobe F.), 5: Pumice tuff & mudstone ([I. Hayaseniori F.), 6: ’I‘uff brcccia emori F.)>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).
719
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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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 (<3OO0C) hydrothermal system, where mixing of moderately saline fluids (of inferred magmatic nature) and meteoric fluids occurred. Ar/Ar data suggest a cooling age of the unaltered protolith of 8 Ma ca., consistent with ages from the literature; the age of sericitization is tentatively indicated at about 7 Ma. These radiometric ages, and reconstruction of the pre-detachment geometry, are compatible with a source of the sericitization fluids from the mafic (monzogranodioritic) Orano Porphyries (Ar/Ar age 6.8 Ma). The CEF acted as the main channelway for mineralizing fluids.
4 BOTRO AI MARMI 2 MARCIANA At Marciana (western Elba Island), exploitation is focused on a lenticular body of porphyritic aplite belonging to the Cap0 Bianco Unit. The unit originally consisted of two main horizons, and represents the oldest known (about 8 Ma) magmatic event at Elba (Dini 1997). The Marciana body belongs to the lower horizon, that was subsequently intruded by granitic poprhyries (about 7.2 Ma), and affected by slight thermal metamorphism associated with the Monte Capanne intrusion (about 6.8 Ma). The mined rock is composed of a very fine grained aggregate of albite, orthoclase, quartz and muscovite. It is apparently fresh, even at the microscopic scale. Only small portions of the mined body show scarce pyrite-sericite-calcite veinlets and disseminations. Hence, the economic qualities of the rock (high alkali and low iron and calcium contents; Table 1) are the result of primary magmatic processes. Their exact nature, whether fractional crystallisation and/or pure crustal melting, is still a matter of debate (Dini 1997).
3 LACROCETI'A The La Crocetta mine (Maineri et al. 1999, 2000) is located in central eastern Elba, and represents the currently exploited portion of a mineralized area including the old mine of Buraccio to the north. The ore is mainly comprised of a pervasively sericitized porphyritic aplite of the Cap0 Bianco Unit (upper horizon), i.e. similar to that mined at Marciana. Outside of the La Crocetta area, other outcrops of the Cap0 Bianco Unit consist of a tourmaline-rich facies, that has no economic value. At La Crocetta, sericitization is apparently spatially related to a lowangle detachment fault (Central Elba Fault, CEF),
The Botro ai Marmi (BM) feldspar deposit is associated with the apical part of an epizonal stock (WAr age 5.7 Ma, Ferrara & Tonarini 1985). The Campiglia Marittima area has long been known for Cu-PbZn skarn deposits (Corsini et al. 1980). These deposits lie 1-2 km E and NE of the stock, in strict spatial association with (4-5 Ma?) porphyry dikes; only minor bodies occur in the immediate proximity of the intrusion. Most chemical data for the BM stock plot in the sienogranite field; some rocks encountered in drill holes show granodioritic compositions (Raveggi 1998, and references therein; Ricci 2000). The economic value of the mined rocks lies in their high K, low Ca-Fe-S contents (Table 1). Mineralogically, this high K content is expressed by the occurrence of two generations of K-feldspar; the first one is in equilibrium with other magmatic minerals, whereas the second shows replacement textures, e.g. after plagioclase. Explanations for the high K content of BM stock include 1) displacement toward the Kfeldspar apex of the melt composition because of assimilation of carbonate country rocks (Poli et al. 1989) and 2) late- to post-magmatic Kmetasomatism. Recent studies (Raveggi 1998, Fulignati et al. 1999, Rossato 1999, Ricci 2000) seem to favor the second mechanism, although the first one may have contributed as well. Specifically, fluids inferred to be in equilibrium with the second generation of K-feldspar show temperatures as high as 500" C, with salinities up to 38 % wt. NaCl and 19 % wt. KCl. These features are typical of porphyry-related metalliferous deposits, also characterized by potassic alteration; however, in the BM system the potassic alteration is metal-poor: metalliferous mineralization occurs in a later skarn to vein stage, and is mostly localized away from the intrusion. 726
5 TORNIELLA The Piloni di Torniella deposit, north of Roccastrada, is the most important of a group of kaolinitealunite deposits of southern Tuscany (Lombardi & Mattias 1987). The deposit results from hydrothermal alteration of Quaternary (2.4 Ma) rhyolitic lavas, described by Pinarelli et al. (1989). Mazzuoli (1967) suggests that the hydrothermal alteration was mainly controlled by the fault system (apenninic, i.e. NWSE, and anti-apenninic, i.e. SW-NE). Hydrothermal kaolinite + alunite replace magmatic feldspar and glass (Bertolani & Loschi Ghittoni 1989). However, both Lombardi & Mattias (1987) and Bertolani & Loschi Ghittoni (1989) describe minor occurrences of sedimentary (re-sedimented?) kaolinite in small lacustrine post-orogenic basins; in these occurrences, of better economic quality, alunite is scarce or absent. In comparison with other deposits of this study, the chemical composition of mined rocks is widely variable (Table 2), and S contents are locally high, because of the presence of alunite; in fact, alunite was also mined in the past, but nowadays it severely detracts from the quality of the ore. Because of the highly variable compositions, no systematic trends of chemical variations between the original rhyolites and the hydrothermally altered material can be recognized. Broadly speaking, there is a tendency toward a decrease of silica, alkalis and iron, and increase of alumina and LOI, with alteration. Cortecci et al. (1981) report for alunites from Torniella 6”s values ranging from +6.7 to +9.6 per mil; they consider these values compatible with an origin from oxidation of sulfides. However, there is no evidence that sulfides were ever present at Torniella; an alternative explanation is oxidation of Table 2. Representative chemical analyses of Roccastrada rhyolites and of mined rocks at Torniella Rhyolites (avg. of 22 data) SiO, Ti02 Fe203 Fe0 MgO CaO K20 Na20 p205 S0,O
LOI Total
73.35 0.24 13.47 0.95 0.91 0.29 0.90 5.08 2.47 0.14 not anal. 1.57 99.38
Min
Torn i ell a A* Max
H,S upon reaction of ascending H,S-bearing steam with cool groundwater (cf. Rye et al. 1992) Even if there is no recent detailed research on the locality, the available information may suggest that the Torniella deposit represents an advanced argillic alteration assemblage formed in a shallow environment by meteoric fluids interacting with (magmatic?) gases (see e.g. Sillitoe 1993). Similar, smaller systems occur elsewhere in southern Tuscany (e.g., Tanelli & Scarsella 1990). 6 CONCLUSIONS
The main deposits of ceramic raw material in Tuscany are associated with acid magmatic rocks of the Tuscan Magmatic Province. Development of oregrade material is the result of a combination to various degrees of magmatic and hydrothermal processes. Primary magmatic rocks are already characterized by comparatively high alkali, and low iron and calcium, contents. Therefore, in some cases (notably Marciana) they may represent ores in themselves. In other localities, the commercial qualities of the currently mined rocks arise from hydrothermal alteration, resulting in either K-enriched (Crocetta, Botro ai Marmi) or kaolinitek a1unite)-rich (Torniella) material. The nature of the fluids involved range from high-temperature, high-salinity, presumably magmatic, fluids (Botro ai Marmi), to moderate temperature fluids of mixed magmatic and meteoric origin (La Crocetta), to presumably steamheated meteoric fluids (Torniella). ACKNOWLEDGEMENTS Financial support was provided by MURST and CNR funds. REFERENCES
B**
15.72 79.13 67.44 46.91 0.19 0.61 0.31 0.35 12.64 38.61 21.67 30.39 0.08 1.97 0.89 1.38 not anal. Not anal. 0.04 1.10 0.51 0.08 0.02 4.40 0.26 0.06 0.95 9.21 3.43 2.93 0.16 2.15 1.22 0.65 0.02 0.99 0.06 0.16 < d.1. 25.86 < d.1 6.48 0.88 11.25 4.17 11.25 99.99 100.66
*A = representative analysis of low-sulfur material * *B = representative analysis of high-sulfur material ‘total sulfur as SO,; < d.1. = below detection limit
Bertolani, M. & A. Loschi Ghittoni 1989. Kaolin of Piloni di Torniella (Tuscany, Italy). L ’industria mineraria 10(2): 1927. Cipriani, C. & G. Tanelli 1983. Le risorse minerarie della Toscana: note storiche ed economiche. Accad. Toscana Sci. Lett. La Colombaria 48: 141-183. Corsini, F., Cortecci, G., Leone, G. & G. Tanelli 1980. Sulfur isotope study of skarn (Cu-Pb-Zn) sulfide deposit of Valle del Temperino, Campiglia Marittima, Tuscany, Italy. Econ. Geol. 75: 83-96. Cortecci, G., Lombardi, G., Reyes, E. & B. Turi 1981. A sulphur isotopic study of alunites from Latium and Tuscany, Central Italy. Mineral Deposita 16: 147-156. Dini, A. 1997. Le rocce porfiriche dell’isola d’Elba: geologia, geocronologia e geochimica. Tesi di dottorato, Universith di Pisa. Ferrara, G. 1962. Nuovi dati sull’intrusione terziaria del Campigliese. Atti Soc. Toscana Sci. Nut. (A) 69: 559-583.
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Ferrara, G. & S. Tonarini 1985. Radiometric geochronology in Tuscany: results and problems. Rendiconti Soc. Ital. Mineral. Petrol. 40: 73-106. Fulignati, P., Marianelli, P. & A. Sbrana 1999. Fluid evolution and minerogenetic implications in the Botro ai Marmi monzogranitic intrusion (Campiglia Marittima, Italy). Terra Nostra, ECROFI XV - Potsdam, 21-24 June 1999,106-108. Lattanzi, P., Benvenuti, M.., Costagliola, P. & G. Tanelli 1994. A n overview of recent research on the metallogeny of Tuscany, with special reference to the Apuane Alps. Mem. Soc. Geol. Ital. 48: 613-625. Lombardi, G. & P.P. Mattias 1987. The kaolin deposits of Italy. L 'industria mineraria 8(6): 1-34. Maineri, C., Costagliola, P., Benvenuti, M., Tanelli, G., Dini, A,, Lattanzi, P., Ruggieri, G. & I.M. Villa 1999. The deposit of raw ceramic material at La Crocetta, Elba Island, Italy. In: C. J. Stanley et al. (eds.), Mineral deposits: processes to processing, Balkema, Rotterdam, 1121-1124. Maineri, C., Benvenuti, M., Costagliola, P., Dini, A., Lattanzi P., Ruggieri, G. & I.M. Villa 2000. Alkali - metasomatic processes at La Crocetta raw ceramic material mine (Elba Island, Italy): interplay between magmatism, tectonics and mineralization. Mineral. Deposita (submitted). Marinelli, G. 1959. Le intrusioni terziarie dell'Isola d'Elba. Atti Soc. Tosc. Sc Nut A 66: 50-253. Marinelli, G. 1961. Genesi e classificazione delle vulcaniti recenti toscane. Atti Soc. Tosc. Sc Nat .Mem. A 68: 74-116. Mazzuoli, R. 1967. Le vulcaniti di Roccastrada (Grosseto). Studio chimico-petrografico e geologico. Atti Soc. Tosc. Sc Nut. A 74, 2: 315-373. Pinarelli, L., Poli, G. & A.P. Santo 1989. Geochemical characterization of recent volcanism from the Tuscan magmatic province (central Italy): the Roccastrada and San Vincenzo centers. Per. Mineral. 58: 67-96. Poli, G., Manetti, P. & S. Tommasini 1989. A petrological review on Miocene-Pliocene intrusive rocks from southern Tuscany and Tyrrhenian sea (Italy). Period. Mineral. 58: 109-126. Raveggi, M. 1998. Studio giacimentologico della miniera di materiali feldspatici di Botro ai Marmi (Toscana meridionale). Tesi di laurea, Universith di Firenze. Ricci, F. 2000. Alterazione idrotermale in campioni profondi dell'intrusione di Botro ai Marmi. Tesi di laurea, Universita di Firenze. Rossato, L. 1999. Caratterizzazione dei fluidi idrotermali associati alla mineralizzazione a feldspati di Campiglia Marittima. Tesi di laurea, Universith di Firenze. Rye, R. O., Bethke, P. M. & M. D. Wasserman 1992. The stable isotope geochemistry of acid sulfate alteration. Econ. Geol. 87: 225-262. Sillitoe, R.H. 1993. Epithermal models: genetic types, geometrical controls and shallow features. In: R.V. Kirkham, W.D. Sinclair, R.I. Thorpe & J.M. Duke (eds.), Mineral deposit modeling, Geol. Assoc. Canada Spec. Pap. 40: 403-417. Tanelli, G. 1983. Mineralizzazioni metallifere e minerogenesi della Toscana. Mem. Soc. Geol. Ital. 25: 91-109. Tanelli, G. 1993. I minerali e la geologia del Campligliese. Un contributo alla cultura dell'arnbiente ed alla gestione del territorio. Accad. Toscana Sci. Lett. La Colombaria 58: 1-47. Tanelli, G. & P. Lattanzi 1986. Metallogeny and mineral exploration in Tuscany: state of the art. Mem. Soc. Geol. Ital. 31: 299-304. Tanelli, G. & A. Scarsella 1990. Tipologia e modellizzazione genetica delle mineralizzazioni aurifere della Toscana meridionale. L'industria mineraria, 3rd series, 1l(2): 1-9.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger,Lisse, ISBN 90 2651 824 2
Occurrence of halite in kaolin of NW Sardinia: genetic implications P.Mameli Istituto di Scienze Geologico-Mineralogiche,Universita degli Studi di Sassari, Italy
ABSTRACT: Field and analytical work, basically SEM-EDS analyses, on the North-Western Sardinia kaolin deposits allow to ascertain high contents of chlorides in the form of NaCl micro-crystals. The micro-textures showed that halite mainly occurs within bleached kaolin and is enveloped by opal-CT which prevented its dissolution by late circulation of groundwater. In six kaolin samples deriving from two different deposits (Locchera and Badu e' Ludu) the C1- content, as dosed by direct calibration, ranges between 2 112 and 20,152 ppm. Literature data do not report such high chloride contents in kaolin deposits. Possibly, this new occurrence seems to be a confirmation that chlorine rich fluids can be the acidifying agent necessary to remove iron, as proposed by Berger & Velde (1992). 1 INTRODUCTION NW Sardinia is characterized by the presence of different kaolin mineralizations exploited since long time. We studied the deposits in the "district of Romana" with the aim of testing the possible use of this kaolin as raw material in zeolite synthesis. Researches have been focused on the structural, compositional, and petrologic features of the different deposits. These data allowed us to characterize the kaolin mineralization and to hypothesize a genetic path. Previous works suggested for these deposits a mild, CO2 controlled, acid environment, with temperatures at about 100°C. Our researches ascertained the occurrence of mineral assemblages that point to different conditions of temperature and pH.
2 GEOLOGICAL OUTLINES Kaolin deposits of NW Sardinia are hosted within a thick volcanic succession pertaining to the Tertiary calkalcaline suite. This volcanic suite ranges from tholeiitic (Coulon 1977, Beccaluva et al. 1987) to shoshonitic (Beccaluva et al. 1987, Mameli 1997) series, and spans between 28 and 15 my. In the studied area the succession is made up by lavas alternating pyroclastic flows with minor amounts of epiclastites. 729
From the bottom the following volcanic units have been distinguished: 0 M.te Seda Oro basalts and basaltic andesites; 0 M.te Tiloromo andesites; 0 M.te Pizzinnu dacitic lava flows. On these basal lava flows different volcanic and pyroclastic flows rest. We have recognised: 0 Dacitic to rhyodacitic Monte Frusciu and rhyodacitic Cossoine cumulo-domes A "pyroclastic complex'' characterised by watersettled falls, pyroclastic flows, welded ignimbrites; 0 The rhyolitic complex of Monte Traessu. Limestones of lower Burdigalian age rest on the previous volcanic sequence. 0 Andesitic and dacitic late effusions "M.te Larenta" type. The area is characterised by two different faults trends that predate and control the Burdigalian transgression. The first one is characterised by ENE and/or EW trending faults and shows the same trend of the Aquitanian strike slip faults. They are sealed by younger lava flows and marine sediments of upper Burdigalian age. The second one is characterised by NNW trending faults and is comparable to the normal faults which border the north Sardinia halfgraben system, and controlled the upper Burdigalian transgression (Funedda et al. 2000). Both fault systems were reactivated during Pliocene uplift of Sardinia.
In all deposits the kaolin shows massive aspect and different degrees of bleaching. Relationships with NNW trending faults are evident in the Monte Tiloromo, Badu e' Ludu and Donigazza deposits; Locchera, which is the largest deposit, and Cuguruntis occur at the intersection between NNW and ENE faults. A rough zoning of alteration facies is still detectable moving away from the main faults: the outer alteration zone, characterised by low fluid\rock ratio, shows assemblages typical of propylitic alteration, whether closer to the main faults zones advanced argillic alteration occurred, leading to widespread kaolinization. The parent rocks of kaolin are different volcanic rocks. Pyroclastites represent the most abundant ones, due to high porosity and low crystal content which enhances reactivity. 3 ANALYTICAL RESULTS AND PROCEDURES Samples were taken at regular intervals both in Locchera and in Badu e' Ludu quarries, in the former at increasing distance fiom a main fault zone, and at increasing depth from the field level in the latter. 6 samples deriving from two different outcrops of the Romana district (Locchera and Badu e' Ludu) were studied in detail. The altered bulk rock was crushed and finely ground in an agate mortar and XRD, DTA/TG/DTG, FTIR, SEM-EDS and conductimetric analyses have been carried out. 3.1 X-Ray Diffractometry The mineralogical composition of the kaolin samples was determinated by XRD analyses using a Philips PW 1730 diffractometer with CuKa radiation at 35 KV and 40 mA. Whole rocks and the fraction smaller than 2 pm were analysed. The latter was separated by dispersion in water and NH3 and sedimentation. Diffractograms have shown the presence of kaolinite, cristobalite or opal-CT, scarce quartz, alunite and pyrite.
STA 409 EP with a Temperature Programmer 410 controlled by computer. Most of the diagrams obtained show only a dehydroxylation peak around 600°C corresponding to kaolinite; some DTG diagrams showed a deflection around 650°C corresponding to dickite. 3.3 Infrared spectrophotometry In order to confirm the occurrence of dickite IR analyses were carried out with a NICOLET 20SXB spectrometer at 2 cm-' work resolution. The sample were diluted 1:lOO in KBr. These analyses have clearly detected the dickite among the mineral phases of the kaolin (Fig. I ) as evidenced by peaks at about 3620, 3655 and 3700 cm-' and by relative intensity of the OH-stretching bands (Farmer 1974). Electron 3.4 Scanning Microanalysis EDS
Microscopy
SEM analyses, with a Zeiss 962 microscope, were carried in order to establish relationships between the morphology of kaolin crystals and temperature, as well as to verify the distribution of some elements within microstructures by elements mapping. From EDS maps, the occurrence of NaCl crystals rimmed by opal-CT has been ascertained (Figs 2-5). 3.5 Conductimetry Conductimetric analyses were carried out to check the chlorine content in bulk rocks with the use of an ORION 94-17B chloride electrode. Direct Calibration was the used analytical technique. This procedure permitted us to determine the Cl- concentration of the samples by comparison to the standards. ISA (ionic strength adjuster, NaN03 5M) was added to all solutions to ensure that samples and standard had similar ionic strength. The chloride ise (ion-
3.2 Thermal methods DTA, TG and DTG curves, carried out on whole rocks and <2 pm fractions, were simultaneously obtained at a heating rate of 10"C/min in air, with alumina as reference in alumina sample holders. Samples, whose weights were between 50-55 mg, were run in a NETZSCH Simultaneous Thermal Analysis
and
Figure 1. FTIR spectrum of Romana kaolin.
730
Table 1. C1- content in the analysed samples
Sample 1 2 Clppm 1666 2112
3 4 5 6 20152 3429 4370 2486
Figure 2. BSE image. Crystals of NaCl rimmed by opal-CT.
Figure 5. Si map.
solutions were measured and plotted on the linear axis versus their concentrations on the log axis. Results obtained are not homogeneous. In fact this kaolin is characterised by a chlorine content generally higher than 2000 ppm and sometimes more than 20.000 ppm (Table 1).
4 DISCUSSION AND CONCLUSIONS
Figure 3. C1 map.
Relationships between kaolin deposits and fault zones allow to envisage in this area a fracturecontrolled, fluid dominated system. A double control on the geometry of kaolin bodies had been exerted: the first by tectonic structures and the second by porosity and glass content of the parent rock. The Locchera deposit is more complex: here protoliths of kaolin are both pyroclastites and hydrothermal breccias which were generated as consequence of sub-surface boiling (Hulen & Nielson 1988, Mameli 2000). This evidence, together with the widespread occurrence of the quartz-dickitekaolinite assemblage point out temperature conditions of fluids between 170 and 200°C (Utada 1980). A subsequent lowering of temperature due to the interaction between juvenile fluids and groundwater should have enhanced Eh rise and also might have permitted the following reactions:
Figure 4. Na map.
selective electrode) used was Ag/AgCl. Four standards were prepared, ranging from 1 to 1000 ppm Cl- content, at the same temperature as the samples (1 7,4"C). The calibration curve was constructed on semilogarithmic paper. Electrode potentials of standard
C12 + H20 = HCl + HClO 73
(2)
In this way an aqueous phase dominated by H+, Cl', and SO-: generates. Such acidic fluid easily evolutes to a saline one, due to the reaction with volcanic rocks such as andesite, dacite, rhyodacite, giving rise Na' and to an assemblage dominated by Cl-, SO,:
and M. Pi10 who provided the conductimetric analyses and G. Oggiano who introduced me to the geology of the studied structures. The referee is kindly acknowledged for helpful suggestions and careful revision of the manuscript.
K'. The critical role of chloride in pH lowering and iron leaching was established by Berger & Velde simulations (1 992). These authors evidenced that a pH value of z4-4.5 does not justify Fe leaching, particularly if kaolin parent rock is basic or intermediate. They concluded that bleaching of kaolin in advanced argillic alteration needs pH values lower than 3, that cannot be accounted only by carbonic leaching, so that they invoked the reaction (2) as responsible of pH lowering even if chlorides do not leave any trace of their presence in the residual mineral assemblage due to their high solubility. In any case a mix of sulphate and chloride salts would be expected in the more or less neutralized fluid. Similar acid neutralization by wall-rock reaction involving exchange of hydrogen ion for base cations was proposed by Reed (1997). SEM and EDS analyses, particularly on those samples coming from contact areas between strongly bleached and still "red", Fe-rich, kaolin evidenced the presence of halite. NaCl crystals are still abundant because of the favourable occurrence of opalCT which rimmed halite crystals preventing them from further dissolution. So chlorine presence, both as volcanic gas and dissociated Clz according to the reaction (2), could have been the most efficient agent responsible of kaolin bleaching. On the light of these new evidences we consider unlikely that pH values compatible with kaolinization of Romana deposits, can be connected with carbonate anion (Garbarino et al. 1994). At low pressure and subsurface boiling conditions, bicarbonate will be decomposed as follows: HC03-= CO2 + (OH)-
REFERENCES Beccaluva, L., Brotzu, P., Morbidelli, L., Serri, G. & G. Traversa 1987. Cainozoic tectono-magmatic evolution and inferred mantle in the Sardo-Tyrrhenian area The lithosphere in Italy Advances in Earth Science Research. Accad. Naz. Lincei: 229-248. Berger, G. & B. Velde 1992. Chemical parameters controlling the propylitic and argillic alteration process. Eur. Jour. Mineral. 4: 1439-1454. Coulon, C. 1977. Le volcanisme calco-alcalin cenozoique de Sardaigne (Italie). Petrographie, geochimie et genese des laves andesitiques et des ignimbrites- signification geodynamiques. These Doct. 3" cycle Univ. Aix-Marseille III, 288 pp. Farmer, V.C. 1974. The layer silicates. In The infared spectra of minerals (Farmer V.C., ed.), Mineralogical Society London, 331-365. Funedda, A., Oggiano, G. & S. Pasci 2000. The Logudoro basin: a key area for the tertiary tectono-sedimentary evolution of North Sardinia Boll. Soc. Geol. It. 119: 3 1-38. Garbarino, C., Masi, U., Padalino, G. & M. Palomba 1994. Geochemical features of the kaolin deposits from Sardinia (Italy) and genetic implications. Chem. Erde 54: 213-233. Hulen J.B. & Nielson D.L. 1988. Hydrothermal brecciation in the Jemez fault zone, Valles caldera, New Mexico: results from continental scientific drilling program core hole VC- 1. Jour. Geophys. Res. 93: 6077-6090. Mameli, P. 1997. I1 vulcanismo oligo-miocenico dell'Anglona: caratteri petrochimici. Plinius 18: 137-139. Mameli, P. 2000. Rilevamento e caratterizzazione mineralogica del caolino della Sardegna settentrionale e proposta di impiego in settori non convenzionali. Tesi di Dottorato XI1 cicl0 Univ. Sassari, 125 pp. Reed, M.H. 1997. Hydrothermal alteration and its relationship to ore fluid composition. In Geochemistry of hydrothermal ore deposits (Barnes H.L., ed.), Wiley & Sons NY, 303365. Utada, M. 1980. Hydrothermal alteration related to igneous acidity in Cretaceous and Neogene formations of Japan. Mining Geol. Jpn. Spec. Issue 8: 67-83.
(3)
Then as CO2 leaves the solution, its pH increases leading to weakly alkaline conditions. The issue of field and analytical work on the Romana kaolin, particularly the occurrence of meaningful amounts of NaC1, represents a good print of chloride rich alterating fluids. Such a circumstance can support the simulation of Berger & Velde (1992). ACKNOWLEDGEMENTS I greatly acknowledge J. Linares and his collaborators for the support at XRD, DTA, TG, DTG and IR analyses, G. Sanna
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Interaction of twinning structure of the feldspars with water fluid - the most significant geological process in the Earth’s crust V.S.Melnikov Inst. of Geochemistry Mineralogy and Ore Formation Nat. Acad Sci. Ukraine
ABSTRACT: Sanidine-microcline phase transition is the result of Al/Si ordering of K-feldspar at cooling of igneous and metamorphic rocks. As a consequence of this transition a twinning domain structure is developed. Originally a size of domains is very small and such cryptotwinning structure characterizes tweed ortoclase. Under certain conditions an enlargement of the scale twinning happens. A study of orthoclase-microcline interrelation in natural samples of alkali feldspar shows that coarsening of twin domains takes place under the influence of water fluid. The deuteric water was localized in K-feldspar structure in magmatic rocks, but for granulites water infiltration in intragranular space is very likely. A water fluid-tweed orthoclase interaction was displayed in huge massifs of granitic rocks.
1 INTRODUCTION Alkali feldspars (AFS) are main rock forming minerals of the sialic Earth’s crust. The real structure of AFS is a result of ordering, twinning and exolution processes (Smith 1974). AFS are monoclinic at high temperatures and triclinic at low temperatures. The symmetry change process is the result of Al/Si ordering in K-rich feldspars. During phase transformation of a K-feldspar triclinic domains are formed and relate to each other as twins. If twin-domains have fine size, then cross hatched twinning is not visible on optical light scale. The formation of a tweed-orthoclase occurs, that it is a monoclinic K-rich feldspar with cryptotwinning structure. Its ordering corresponds to low or intermediate microcline. The domain structure is observed only under transmission electron microscope (TEM). The system which consists of the finest domains is unstable because in consequence of fully coherent interfaces between twin domains in crystal lattice a strong strain arises Under some conditions enlargement and coarsening of twin-domains happens and cross-hatching microcline is developed. Many studies have shown that interaction with deuteric fluid is a main factor of the orthoclase-microcline transformation (Parsons 1978, Walker et al. 1995). However, wide-spreading of the rocks which contain orthoclase or mixture of orthoclase and microcline show that conditions for orthoclase-microcline transition may not always be
realised. At last one question connected with waterfeldspar interaction exists. It is not clear where disposition of fluid in rocks and how it moves through crystal.
2 EVOLUTION STAGES OF TWIN STRUCTURE TRANSFORMATION Using results of optical, TEM and x-ray studies of the orthoclase from different rocks, four stages of domain coarsening may be distinguished. 1. At the sanidine-microcline inversion homogeneous tweed structure on the scale < 5 - 10 nm appears. Its feature is a full coherence of domain boundaries and as a result a strong strain appears and twodimensional modulations in twin structure are formed. Diffraction symmetry of tweed orthoclase is monoclinic.
2. A regular twin structure is developed and visible on a scale of light microscope as a cross-hatching microcline. When the size of twin domain is smaller then above mentioned scale two dimensional large scale modulations are developed and pseudo-twins with albite and/or pericline orientation are observed (Fitz Gerald & Harrison 1993). 3. The following stage of coarsening of twin structure is a development of patch structure which represents a mosaic of large twin domains of low microcline. Twins by albite-law are dominant. 733
perthites, subgrains and micropores are microfeatures of deuteric alteration of AFS in plutonic rocks. New microtexture formation occurs at 400-450°C (Walker et al. 1995). This temperature is close to the temperature point of orthoclasemicrocline transformation, in some cases, a good correlation between development of the perthites coarsening and amount of microcline phase in orthoclase crystal appeared. The chamber pegmatites and diaphthoresis zones in charnockites are two different geological situations for which orthoclasemicrocline transformation is well known. However, water fluid which affects orthoclase may be of different origin. In pegmatites it may be localised in feldspar crystals. An orthoclase-microcline transformation in charnocites occurs through interaction with fluid flows that have been introduced in rocks on diaphthoresis stage.
4. The forming action of non-twinning microcline with small relicts of patch structure is a final stage of orthoclase transformation. Undoubtedly, this scheme is a highly idealized model of the real process. K-feldspar does not reveal simple niicrotextures. It shows their mixture and even such a complicated twin construction as an “irregular” niicrocline (Banibauer et al. 1989). Domain structure of orthoclase is not stable, because total energy of twin boundaries is large. The decrease of crystal fiee energy may be obtained by increase of the domain sizes. Distinct structure of the domain boundary prevents it from domain growth since rnonodomainization process is a transformation of the right domain into the left one or vice versa. It is supposed that a degree of Al/Si order of the domain boundary is similar to low sanidine and nionoclinic layers stabilise orthoclase tweed structure. Al/Si ordering is necessary for their destruction, but at low temperature %/A1 ordering is a slow process. It was shown that transformation of monoclinic-triclinic K-feldspar at 500°C is very sharp (Brown & Parsons 1989). Near inversion point the domain structure is unstable and thermal fluctuations may be the reason of the change in domain size. Slow cooling of rock favours the development of a large-scale domain structure. The detailed mechanism of the catalytical influence of water on twin transformation in the initial stage is also not known, but it “works” on the atomic level. The presence of water in reforming the crosshatching twins to patch twins is more evident. This process has analogy with deuteric coarsening of perthites (Parsons 1978, Waldron & Parsons 1992). Both these processes often happen at the same time. It is not known what twin domain size scale may be achieved at only physical enlargement, because there are also may be other factors.
3.1 Chamber pegmatites The Volyn (Ukraine) chamber pegmatites, that located in rapakivilike granites are crystallised in relatively closed system. Its degree of differentiation is different and relation with pegmatite size and initial water content in silica melt has been established. Forming contrasted zones during crystallization is a characteristic of a highly differentiated pegmatites. K-feldspar is orthoclase, microcline or its mixture. Some common features of its relationship in different zones are inherited in great number of pegmatites (Melnikov et al. 1991). 1. Orthoclase is a prevailing phase in all zones of slowly differentiated pegmatites. In highly differentiated pegmatites the microcline abounded but in outer zones it is often found together with orthoclase. 2. For every distinct pegmatite a microcline content is increased from graphic zone to quartz core. Orthoclase/microcline ratio is variable in intermediate zones, but tendency for orthoclase decrease remains.
3 PETROLOGICAL DATA Orthoclase is a metastable phase in igneous rocks (subsolvus granites and syenites, some pegmatites) and metamorphic rocks (granulitic charnockites). It was established that these rocks are crystallized under relatively “dry” conditions. If a coexistence of the orthoclase and microcline is observed the indications of the rock interaction with water may also be observed. The turbidity of alkali feldspar is a common macro feature of interaction with water fluid (Parsons 1978). Coarsening of perthites and the formation of a new microtexture with patch
3. In surrounding granites the orthoclase is always more abundant. During crystallisation of pegmatitic zones water content in silica melt is increased. At the growth an alkali feldspar is in equilibrium with silica melt. When water dissolves in melt some portion of it enter into the solid phase (AFS). One can see that fraction of water included in crystal has been proportionate to the water content in silica melt. Hence, K-feldspar that crystallised on later stage will 734
500
microcline. Since, a diaphtoresis is result of water entrance into the system, it is more probably that a water movement in rock occurred in intragranular space and microcracks. Probably time is a main factor in this process, but it is also known that the deformation of orthoclase crystal under stress involves its transition to microcline (Smith 1974).
0 0
20
40
60
80
100
miaodine, % Figure 1. A plot of KiRb ratio versus microcline content in Kfeldspars from chamber pegmatites of Volyn (Ukraine).
be enriched by water to a greater degree. It is supposed, that water affected the structure of tweedorthoclase and conversed it to cross-hatching microcline. That supposition well corresponds to the observed data on distribution of the K-feldspar varieties in different zones. It is clear that the process is possible if water fluid has been included in a crystal. As it is known the relic melt enriched by water accumulates some rare metals (Li, Be, Rb, Cs, and others). A size of Rb’-ions is favourable for substitution of K’ in the structure of K- feldspar. It is established for a chamber pegmatites that a Rb content increases from graphic zone to inner zones. A number of microcline phase in AFS also increases. Figure 1 shows that WRb and orthoclase/microcline ratio reflects a connection of the activity of the detwinning agent (H20) in feldspar‘s structure and water fluid concentration in silica melt. The correlation between coarsening of perthites and development of a cross-hatched microcline also exists. There is no evidence that a later hydrothermal fluid (< 450°C) which has been localised on subgrain‘s interface and in microcracs affects the tweed-orthoclase. In the graphic zone, where metasomatic mica was developed, a tweedorthoclase has not been changed. Perhaps effect of the water fluid lasted for a short time or conditions for infiltration solution were not sufficient.
4 DISCUSSION AND CONCLUSION Orthoclase appears as a result of a phase transformation of low sanidine during rock cooling. Cryptotwinning structure of orthoclase is unstable, but the microcline is the stable phase. At orthoclasemicrocline transformation an increase and coarsening of twin domains is performed. The mechanism of this transformation is not sufficiently clear. It is supposed that two main stages of twin domains coarsening exist. At first, domains increase near point of orthoclase-microcline transformation. When size of twin domains approaches light microscope scale then cross-hatched twin structure is observed. Water participation in this process is problematic. Coarsening of twin domains under influence of water fluid occurs in the second stage. As a result of enlargement of twin domains is a formation of monodomain regions. The mechanism of this stage looks like recrystallisation. Two modes of fluid localization are possible: a) in single crystals of feldspar as a submicroinclusions, lattice defects or as a chemically bonded water; b) in intragranular space and microcracks as capillary water. Deuteric water is localised mainly in single crystals. In spite of that many features of mechanism of water-twins interaction is unknown, this process is widespread in natural samples of AFS. The large granitic plutons of the Earth’s crust contain a microcline as a main mineral, but initial phase has been an orthoclase.
3.2 Charnockites The main fluidic component during crystallisation of granulitic rocks is COl(Touret 1971). This means that a granulitic system had a “dry” regime of formation. Orthoclase development is a main peculiarity of granulites and charnockites. AFS in rocks of the amphibolite facies represents a crosshatching microcline. However, some charnockites with features of diaphtoritic alteration show the microcline development as a result of tweedorthoclase changing. Initial tweed-orthoclase often coexists with microtwinning and nontwinning
REFERENCES Brown, W.L & J. Parsons 1989. Alkali feldspars: ordering rates, phase transformation and behaviour diagrams for igneous rocks. Mineral. Mag.53: 25-42. Bambauer, H.U., Krause, C. & H. Kroll 1989. TEM investigation of the sanidineimicrocline transition across metamorphic zones: the K-feldspar varieties. Eur.J. Mineral. 1 : 47-58. Fitz Gerald, J.O. & T. Harrison 1993. Argon diffusion domains in K-feldspar I : microstructures in MH- 10. Contrib. Miner. Petrol. 113: 367-380.
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Melnikov, V.S.. Pavlishin, V.I., Pshentsova N.P. & 0.V.Usenko 199 I . Structural features and symmetry of alkali feldspars from Chamber pegmatites of Volyn (in Russian). Mineral.Journ. 13(4): 12-24. Parsons, 1. 1978. Feldspars and fluids in cooling plutons. Mineral.Mag.42: 1- 17. Smith J.V. 1974. Feldspars minerals.V.2.Chemical and textural properties. Springer - Verlag Berlin Heilderber New Yurk. Touret J. 197 1. Le facies granulite en Norvege meridionale, 2. Les inclusions fluides. Lithos 4: 423-436. Waldron, K.A. & I. Parsons 1992. Feldspars microtexture and multistage thermal history of syenites from the Coldwell Complex, Ontario. Contrib.Petrol.1 1 1: 222-234. Walker F.D.L., Lee K.M. & I. Parsons 1995. Micropores and micropermiable texture in alkaly feldspars: geochemical and geophysical implications. Mineral.Mag. 59: 505-534.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Fossil geothennal systems in the continental rift zone of the Ku@ik Menderes within the Menderes Massif, Western Anatolia, Turkey Nevzat Ozgiir Siileynzan Demirel Cniversitesi, Miihendislik-Mimarhk Fukiiltesi, Jeoloji Miihendisligi’ Boliimii, 32260 Ispurtu, Turkey
ABSTRACT: The Hg, Sb and Au deposits of Halikoy, Emirli and Kiire can be considered as fossil equivalents of formerly active geothermal systems due to various similarities of parameters. The metals originated in the host rocks were transported in a hydrothermal system as (HS); complexes. The precipitation of metals depended upon T, P, pH and redox potential as well as total sulfur concentration. 1 INTRODUCTION The Anatolian and Aegean micro plates control the plate tectonic position of the Eastern Mediterranean area between the Eurasian and African plates. This plate tectonic development results in the lifting of the Menderes Massif in Western Anatolia, Turkey showing a dome structure due to compressional tectonic features from Oligocene to Miocene (Ozgiir 1998, 1999). From Early Miocene to Middle Miocene, the continental rift zones of the Buyuk Menderes, Kuqiik Menderes, and Gediz were formed by extensional tectonic features, which strike E-W generally and are represented by a great number of Hg, Sb, and Au mineralizations and thermal waters in connection with volcanic rocks from Middle Miocene to recent (Fig. 1; Ozgur 1999). The Hg, Sb and Au mineralizations and thermal waters are related to faults, which strike preferentially NW-SE and NE-SW, diagonal to the general strike of the continental rift zones. These faults, representing a multitude of Hg, Sb and Au mineralizations and thermal waters, are probably generated by compressional tectonic stress, which leads to the deformation of uplift between two extensional continental rift zones. The investigated Hg, Sb and Au deposits of Halikoy, Emirli and Kure (Fig. 1) represent typical examples of epithermal mineralizations. The aim of this paper is to introduce Hg, Sb and Au deposits in the continental rift zone of the Kuquk Menderes within the Menderes Massif and to compare them genetically to each other with regard to (i) relationship between Hg, Sb and Au deposits as fossil geothermal systems, active geothermal systems, tectonic features and volcanism, (ii) fluidrock interactions, (iii) geochemical,
hydrogeochemical and isotope geochemical features, (iv) fluid inclusion studies in ore and gangue minerals and (v) source, transport and deposition of Hg, Sb and Au deposits.
2 GEOLOGIC SETTING The Hg, Sb and Au mineralizations of Halikoy, Emjdi and Kure occur 25 km S and SE of the town of Odemiq in SE part of the rift zone of the Kuquk Menderes. The metamorphic rocks of the Massif are (i) Precambrian to Cambrian core series consisting of high-grade schists, gneisses, granites and metagabbros and (ii) Ordovician to Paleocene cover series composed of mica schists, phyllites, metaquartzites and metagabbros (Fig. 1; Dora et al.
Figure 1. Epithermal ore fields in the rift zone of the Kuquk Menderes within the Menderes Massif. 1: Hg deposit of Halikoy; 2: Sb deposit of Emirli; 3 : Au deposit of Kiire.
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1995). In Cambrian, the first metamorphism in the massif took place under amphibolite to granulite facies conditions affecting the core series. The late phase of the Variscan orogenesis affected the massif for a second time, which may be postdated by the Early to Middle Triassic granites. In Eocene, highpressure metamorphism was generated under epidote-blue schist to eclogite facies conditions, and the late main metamorphism developed in Late Eocene to Early Oligocene (Dora et al. 1995) and is overprinted by the Barrowian-type metamorphism. The metamorphic rocks in the massif are overlain by Neogene to Quaternary thick sedimentary rocks. From Middle Miocene to recent, an intense volcanism was generated in connection with the evolution of the continental rift zones in the Menderes Massif (Ercan & Gunay 1981; Ercan et al. 1983, 1992; Ozgur 1998). The volcanic rocks in the NE part of the Hg, Sb and Au deposits of Halikoy, Emirli and Kure in the rift zone of the Kuqiik Menderes are distinguished by Rb/Sr age of 15,O k 0,2 Ma in KaraburC and a K/Ar age of 16,7 k 0,5 Ma in Yenigehj- and can be classified into Middle Miocene (Ozgur et al. 1997; Ozgur 1998). These volcanic rocks may be considered as products of continental crust due to isotope analyses of s7Sr/86Sr and 143Nd/'44Nd.The youngest volcanism in the Menderes Massif is characterized by the Late Pliocene volcanic rocks of Denizli (Ercan et al. 1983) and the volcanics of Kula with an age ranging from 7,5 Ma to 18.000a (Ercan et al. 1992). Finally, the volcanic rocks can be considered as heat source for the heating of thermal fluids in the continental rift zones in addition to earthquake activity and heat flow anomalies.
3 MERCURY, ANTIMONY AND GOLD DEPOSITS The strongly altered mica schists and quartzites form the host rocks of Hg, Sb and Au deposits of Halikoy, Emirli and Kure. The Hg deposit of Halikoy consists of ore formations in terms of veins and veinlets, which contain cinnabar, metacinnabarite, pyrite, marcasite, chalcopyrite and quartz and calcite as gangue minerals (Fig. 2). The Emirli Sb deposit is distinguished by veins and veinlets and consists of pyrite, arsenopyrite, stibnite, sphalerite, chalcopyrite, tetraedrite, marcasite, orpiment, realgar, cinnabar and gangue minerals of quartz, adularia and calcite. As mineral assemblage in the Au deposit of Kure representing ore types of veins and veinlets, there are arsenopyrite, gold, pyrite, marcasite, fahlore and gangue minerals of quartz and calcite. In the abandoned Halikoy Hg mine, there are measured ore reserves of 56.000 metric tonnes with
Figure 2. Ore mineral assemblage of epithermal Hg, Sb, and Au deposits of Hahkoy, Emirli and Kure in the rift zone of the KuGuk Menderes.
a Hg mean value of 0,30 percent, indicated ore reserves of 210.000 metric tonnes with a Hg mean value of 0,25 percent and possible ore reserves of 210.000 petric tonnes with Hg mean value of 0,23 percent (Ozgiir, 1998). The recoverable ore reserves of the Sb deposit of Emirli are estimated at 400.000 tonnes with an average content of 5,5 percent Sb. The Au deposit of Kure consists of three occurrences and shows Au contents up to 30 ppm.
4 FLUID INCLUSION STUDY AND FLUID GEOCHEMISTRY For the fluid inclusion studies of the Hg, Sb and Au deposits of Halikoy, Emirli and Kure, we have collected about 50 quartz samples (stage 2), which were compared with quartz samples of the crystalline massif (stage 1) and stibnite crystals. The fluid inclusion measurements were made in the Freie Universitat Berlin, Germany using a Linkam THMSG 600 programmable freezing-heating stage attached to a Leitz Ortholux transmitted-light microscope. This used stage has a dynamic range from -180 to 600 "C. The measurements in stibnite crystals were used in GeoForschungsZentrum Potsdam, Germany by a FLUID.INC SYSTEM gasflow freezing-heating stage attached to an Olympus IR transmitted-light microscope. Three types of primary fluid inclusions were recognized in quartz and stibnite crystals. The type (i) is an inclusion (H2O-NaCI), which can be observed in quartz stage (1) as well as in stibnite crystals and shows two phases (liquid and vapor) at room temperature (Ozgur 1998). The type (ii) is a system of H20NaCl-CO2 f CH4 of which inclusions show three phases of CO2 (CH&iquid, CO2 (CH4)-gas and aqueous solutions at room temperature and an increase in salinity. The type (iii) is a system of
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aqueous vapor, which is associated with Sb deposit of Emirli and indicates boiling conditions in the system. The quartz samples of stage (2) which are associated with ore deposits show homogenization temperatures from 180 to 300 "C and salinity ranging from 3,75 to 9,O NaCl (eq wt %), which correspond with each other (Ozgur et al. 1997). The quartz samples of the Halikoy indicate average homogenization temperature of 128 "C and salinity of 3,4 NaCl (eq wt %) (Gokqe & Spiro 1995). The quartz samples of the Au deposit of Kure is distinguished by the homogenization temperatures from 210 to 300 "C and a salinity between 1,60 and 10,O NaCl (eg wt Yo). In order to understand the composition of hydrothermal ore-forming fluids we have analyzed the fluid inclusions in the quartz samples to 6l80and tj2H by mass spectrometer, K+, Na', Ca2', Mg2+, Fe3+, A13+ by ICP emission spectrometer, F- by anion-selective electrode, C1' by- mercury thiocyanade colorimetric analysis, SO4 by barium sulfate turbimetric analysis, and HCO3- by acid-base capacity titration in the State Key Laboratory for Research of Mineral Deposits, Nanj ing University, PR China. The stable isotopes of 6l80 and 2i2H in fluid inclusions of quartz samples from the investigated Hg, Sb and Au deposits show a connection with them of active geothermal systems (Ozgur et al. 1997; Ozgur 1998). The strongly deviation of the 6l80 from the global meteoric water line proves the intensively fluid-rock interaction in the hydrothermally environment during the differences in the degree of isotope shift from the active geothermal systems to investigated Hg, Sb and Au deposits indicate a relationship between the.. both systems in meteoric origin (Ozgur et al. 1997; Ozgur 1998). The fluid inclusion studies show that the fluids in quartz and stibnite crystals from the ore fields of Halikoy, Emirli and Kure are of Na-HC03 type similar to those of active geothermal systems in the rift zones of the Menderes Massif, whereas the analyses of anions and cations in fluid inclusions of quartz crystals show Ca-(Cl)-(S04)-HC03 type exchange fluids which form a sharp contrast with them of active geothermal systems and fluid inclusion studies. In one respect, this may relate to the process of albitization by Na metasomatism in the investigated area and its environs in which Ca contents dominate during Na contents decrease due to the same process (Nebert & Ronner, 1956).
NE-SW and NW-SE and are located diagonal to the general strike of the rift zone of the Kuquk Menderes (Fig. 1). The calcalkaline volcanics of Middle Miocene age, which occur in the E and NE part of the investigated ore fields, indicate basic towards acidic features and crustal origin (Ozgur, 1998) and seem to be closely related with tectonic feat-pes, ore deposits and active geothermal systems (Ozgur & Pekdeger 1995). The intensively altered mica schists and quartzites, which form the host rocks for Hg, Sb and Au deposits, can be considered as source rocks for the metals, as supported by leaching tests (Ozgur 1998). The hydrothermal alteration is distinguished by phyllic, argillic and silicic k hematitization alteration zones which are comparable with those of active geothermal systems in the continental rift zones of the Menderes Massif. This type of alteration is comparable with adularia-sericite-typemineralization due to presence of adularia and bladed calcite crystals. The homogenization temperatures of quartz and stibnite crystals range from 150 to 300 "C (Ozgur et al. 1997), which can be compared with geochemical temperatures of thermal water reservoirs from 220 to 260 (Ozgur & Pekdeger 1995). The ore-forming fluids of Halikoy, Emirli and Kure show a mean value of 6 NaCl eq wt %; it is comparable with the salinity of active geothermal fluids. The isotope ratios of 6l80 and 62H in fluid inclusions of quartz crystals of Halikoy, Emirli and Kiire show a similarity with those of active geothermal fluids. The strong deviation of the 6l80 values from the meteoric water line shows the intensive fluid-rock interaction in the hydrothermal environment. The trend of deviation increases linearly from the active geothermal field to the epithermal ore fields indicating a relationship between the two. Finally, it might be concluded that (i) the fluids of Hg, Sb and Au mineralizations of Halikoy, Emirli and Kiire can be attributed to a meteoric origin due to stable isotopes of 6l80 and ?i2H, (ii) the metamorphic rocks act as the source of metals of Hg, Sb, As and Au which are forming ore deposits in the rift zones and leached from the metamorphic rocks by fluid-rock interaction and transported as bisulfide complexes with circulating geothermal fluids to the subsurface environment between 500 to 1500 m in depth at temperature below 350 "C. Ultimately, with the cooling of magma chamber, the Au mineralization of Kure was formed at temperatures below 300 "C. Arsenopyrite is present in the ore mineral assemblages of investigated Hg, Sb and Au mineralizations. In connection with further cooling of magma chamber, the Sb mineralization of Emirli took place at temperatures from 180 to 250 "C. The Hg deposit of Halikoy is
5 DISCUSSION The Hg, Sb and Au deposits of Halikoy, Emirli and Kure are related to faults which strike preferentially 739
Ercan, T., Dinqel, A., Tiirkecan, A. & G. Erdogdu 1992. Petrochemical characteristics and genetic interpretation of the basaltic volcanism of Kula (Manisa), Turkey. Geologica Balcanica 22: 5 1-73. Gokqe, A. & B. Spiro 1995. S u l h isotope study of the antimony and mercury deposits in Beydagi (Izmir, Western Turkey) area and the origin of the sulfur in stibnite and cinnabar. Tr. J. Earth Sciences 4: 23-28. Nebert, K. & F. Ronner 1956. Alpidic albitization evidences in the Menderes Massif and its environs. Maden Tetkik ve Arama Enstitiisii Dergisi 48: 83-96. Ozgiir, N. & A. Pekdeger 1995. Active geothermal systems in the rift zones of the Menderes Massif, Western Anatolia, Turkey. in: Kharaka, Y.K. & Chudaev, O.V. (eds.): Proc. Internat. 8th Symp. on Water-Rock Interaction, VladivostokBussia: 529-532. Ozgiir, N., Halbach, P., Pekdeger, A., Sommer-v. Jarmersted, C., Sonmez, N., Dora, O.O., Ma, D.-S., Wolf, M. & W. Stichler 1997. Epithermal antimony, mercury, and gold deposits in the continental rift zone of the Kiiqiik Menderes, western Anatolia, Turkey. in: Papunen, H. (ed.): Proc. 4th Biennial SGA Meeting, Turku, Finland, 1 1- 13 August 1997: 269-272. Ozgiir, N. 1998. Aktive und fossile Geothermalsysteme in den kontinentalen Riftzonen des Menderes-Massives, WAnatolien, Tiirkei. Habilitationsschrift, Freie Universitat Berlin, 171 p. Ozgiir, N. 1999. Active and extinct geothermal systems in the continental rift zones of the Menderes Massif, Western Anatolia, Turkey. in: Stanley et al., 1999 (eds.): Proc. 5th Biennial SGA Meeting and 10th Quadrennial IAGOD Meeting, London, UK, 22-25 August 1999: 559-562.
generated as the last mineralization at temperatures from 128 to 200 "C. The meteoric fluids percolate above permeable clastic sediments in the reaction zone of the roof area of magma chamber in depth of 2-3 km where the fluids are heated and ascend to the surface because low density. The volatile components of C02, S02, H2S and HC1 from magma reached the geothermal water reservoir as ascending gas phases where an equilibrium reactions between altered rocks, gas components and fluids took place. The ascending fluids contain C02, H2S and HCI particularly. Hydrothermal convection cells press the heated fluids toward the surface because of their lower density. Thus, geothermal waters ascend in tectonic zones of weakness. As geochemical pHneutral fluids, the waters outlet at the surface as hot springs, gas and steam. The fluids indicate a reduced pH-neutral sphere in the reaction zone after equilibrium with host rocks. At the subsurface spheres, the ore deposits are generated in terms of stockwork mineralizations (veins, veinlets) and gangue mAnerals represented by quartz, calcite and adularia (Ozgiir 1998). The Hg, Sb and Au deposits of Hahkoy, Emirli and Kiire can be assigned to an epithermal type in connection with an calkalkaline volcanism in Middle Miocene age, comparable to other Hg, Sb and mineralizations in the rift zones of the Menderes Massif, similar to the epithermal Sb and Au deposits in the metallotect of Jiangnan, PR China and active and extinct geothermal systems of New Zealand, and considered as fossil equivalents of active geothermal systems. ACKNOWLEDGEMENTS The project has been supported by the Commission for Research and Scientific Training for New Recruits, Freie Universitat Berlin. We would like to thank Mrs. S . Altinkale, Mrs. D. Yaman and Mrs. M. Zerener for electronic preparation of figures with their indefatigable patience and feeling.
REFERENCES Dora, O.O., Candan, O., Diirr, S. & R. Oberhansli 1995. New evidence on the geotectonic evolution of the Menderes Massif: in: Piskin, O., Ergiin. M., Savascm, M.Y. & Tarcan, G. (eds.): Proc. Internat. Earth Sci. Colloqium on the Aegean Region, 9-14 October 1995, Izmir-Gull&, Turkey, 1: 53-72. Ercan, T. & E. Giinay 1981. Tertiary volcanism in Soke area and Its regional distribution. Jeomorfoloji Bull. 10: 1 17137. Ercan, T., Erdogdu, G. & H. Bas 1983. Petrology and plate tectonic implications of Denizli volcanics. Geol. Soc. Bull. Turk. 26: 153-159.
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Water-Rock Interaction 2001, Cidu (ed.), 02001 Swefs & Zeiflinger, Lisse, ISBN 90 2651 824 2
Gold in Sardinia: recent developments in exploration and exploitation J .Rayner & D .Manis Sardinia Gold Mining S.PA., Sardinia, Italy
ABSTRACT: Sardinia was a major producer of lead and zinc in Europe in the last 150 years. It is now emerging as a significant gold province. The discovery of epithermal gold mineralisation associated with Tertiary aged volcanics was made in the late 1980’s. Gold is currently being mined at Furtei and resources at Osilo and Furtei total almost 1M oz Au. In addition, there is also potential for Hercynian aged arsenic - stibnite veins to host significant gold deposits.
1 INTRODUCTION
Until recently, Sardinia was not regarded as a gold province. Minor amounts of this metal were noted during the processing of As, Sb 2 Cu t Zn t Pb ores of the historic Villasalto, Baccu Locci and Genna Ureu mines in the eastern Paleozoics and in the Pb, Zn and Ag ores of Monte Vecchio (Fig. 1). These styles of mineralisation are hosted in mesothermal veins related to granitic intrusions and deformation during the Hercynian. Epithermal gold mineralisation, was recently discovered in Sardinia as a result of research studies by Cagliari University in the late 1980’s (Fiori et al. 1990). Both high and low sulphidation epithermal gold and porphyry Cu-Au mineralization are hosted in Oligo - Miocene calc-alkaline volcanics that developed in a back arc rift environment in response to a northwestward - dipping subduction zone of oceanic crust located east of present day Sardinia (Coulon 1977, Lattanzi 1999). Open-cut mining of oxide ore is being undertaken at the Furtei mine and resource drilling is being completed on a number of veins at the Osilo Project by Sardinia Gold Mining S.p.A, a joint venture between Gold Mines of Sardinia and Progemisa S.p.A. Mining at Furtei commenced in mid 1997 and as at October 2000, 100,000 oz of gold have been produced from CIL and dump leach operations. Mining and flotation of sulphides is scheduled to commence in early 2001. Resources and reserves at Furtei currently total 9.3Mt at 2.08 g/t Au (620,000 oz Au) and 15000t
Figure 1.Location map showing gold exploration projects 74 1
Cu. At Osilo, resources total 1.66Mt at 7.1 g/t Au and 30 g/t Ag (376,000 oz Au and 1.6 M oz Ag).
2 FURTEI The Furtei Volcanic Complex (FVC) is a major Tertiary age sub - aerial calc - alkaline volcanic centre 5 km in diameter. It is located on the eastern side of the Campidano Graben (Fig. 1). The volcano - sedimentary sequence is dominated by andesitic lavas and domes, block and ash deposits, interlayered pyroclastics and sedimentary rocks (Ruggieri et al. 1997). A central diatreme of dimensions 1000 m north south by 600 m east west crosscuts the volcano sedimentary sequence. Miocene limestone overlies the FVC to the east and west. The earliest phase of gold mineralization is hosted in sheeted microfractures developed in flat lying sediments and ignimbrite sandwiched between andesite flows. Gold is associated with silica, pyrite and lesser sphalerite. Copper and tellurium content are low. This style of mineralization is termed “stratabound” and has been intersected over a large area, 2 km north south and 500m east west in the Sa Perrima - Coronas Arrubias (SP-CA) area, and at the Bruncu Murdegu prospect. The average thickness of the mineralization is 12-15m and average grades are 1.5 - 2.5 g/t Au. Feeder structures to the SP-CA mineralization have not been located, but higher grade structures have been intersected at Bruncu Murdegu, (54 m at 5.4 g/t Au in diamond drill hole, Fig. 2) The stratabound mineralisation is crosscut by the diatreme breccia, and the latter is the predominant host to sub - vertical vuggy silica structures containing pyrite, enargite and gold mineralization. The structures trend N70 and N20, and are irregular in continuity and shape. They vary from a few centimeters to 50 m in width and up to 100 m in length. They can occur as vertical shoots (e.g. Est, Cima) or as discrete pods (Is Concas, Su Coru). Mineral zoning is from luzonite, tetrahedrite, pyrite, coloradoite at higher levels to enargite, tennanite, calaverite and hessite at deeper levels at Cima and East (Maxwell 1995). The location of intersecting faults and mega blocks of andesite within the diatreme are the dominant controls on the higher grade zones. Late and distal low sulphidation epithermal quartz veins occur to the south west at Bruncu de Didus and Amigu Furoni. They are generally less than lm wide and carry 2 -6 g/t Au. The source intrusion to the mineralization is postulated to be at 1-1.5 km depth beneath the FVC (Meloni 1994). 742
Figure 2. Geological map, Furtei
3 OSLO Gold and silver mineralization is hosted in a number of low sulphidation epithermal quartz veins that predominantly trend N70 in the Osilo region (Fig. 1). The veins are hosted in a thick pile of andesites, sediments, dacites and dacitic domes. The veins are generally 2 - 4 m wide and up to 3 km long. Economic mineralization is known to occur over a 250 - 300 m vertical range. Quartz textures in the higher grade zones (10 -50 g/t Au) consist of colloform and crustiform banding, platy calcite replacement and sulphidic breccia with minor adularia (Simeone & Simmons 1999). Silica sinter is preserved at Pedra Canarza and dense chalcedonic quartz with geyserites are present along the Sa Makhesa vein in the north (Fig. 3). Elevated amounts of Hg, As, Sb and Ag occurs at the higher levels of the veins with increasing Pb, Zn, and Cu at depth. Ag to Au ratios decrease from 10:1 to 2: 1 from the outer zones of the veins towards the higher grade shoots.
Figure 4. Prospect location and geology, Eastern Palaeozoics
Figure 3 . Osilo low sulphidation epithermal vein field
gold is associated with low temperature silica-pyrite, carbonate replacement and skarn at the contact of dacitic porphyry and limestone. High grades (27 g/t Au) from rock chip channel samples and drilling have been obtained over a 5 km east west trend of the stock. The alteration consists of a cylindrical core of potassic alteration surrounded by propylitic alteration (Baker 1999, Stefanini & William - Jones 1996). In the Bosa to Scano Montiferro and Romana areas, high level exposures of low sulphidation veins and fossiliferous sinter crop out over a large area. Anomalous As (2000 ppm), Hg (5000 ppm) and up to 0.5 g/t Au have been returned from silica sinter. Kaolinite is currently mined at Romana, and is a product of supergene processes on argillic alteration and silicified hydrothermal breccias.
The Osilo vein field is interpreted to be controlled by dextral strike slip faulting. The higher grade shoots occur along the east-west flexures (dilational jogs) of the veins, or at intersections of veins, or at a change in the competency of lithologies that hosts the vein. Soil geochemistry and interpretation of quartz textures to determine the level of exposure of the vein are important exploration tools. Strong demagnetisation of the country rocks is associated with epithermal alteration immediately around the veins. The demagnetised structures continue along strike to join different vein sets up to 12 km apart (Sa Pala to Pireddu, Bunnari to Pala Edra). In excess of 50 linear km of potential mineralized structure can be delineated in this way. SGM is completing the drilling out of 5 of the 20 known veins in preparation for underground mining to initially commence at Pala Edra and Bunnari.
A large 1 km diameter diatreme breccia occurs in the southern Narbolia region. Low sulphidation quartz hematite veins crosscut the breccia and carry up to 6 g/t Au at the Marragattu sheeted vein prospect.
4 OTHER EPITHERMAL GOLD PROSPECTS Porhyry copper-gold styles of mineralisation occur at Calabona and Siliqua (Sillitoe, 1991 a,b). Low sulphidation veins and sinter predominates at Scano Montiferro, Narbolia, Bosa and Romana areas (Dessi et al. 1996) (Fig. 1). The Calabona area was previously mined for smithsonite (3,000 tons), manganese, copper (15,000 t at 8.2%) and pyrite, and explored by Noranda for copper in the 1970’s. Gold was first noted in Progemisa drill cuttings (12g/t Au) in 1991. The
5 PALEOZOIC GOLD MINERALISATION The Eastern Paleozoic (Sarrabus) area has been mined over the past century for Ag, As, Pb and Zn. Gold was noticed in the dressing plant at Baccu Locci and assayed up to 12 g/t Au in the Sb ores at Genna Ureu. Gold is also present in the Sb ores of 743
J. Ravner. D. Manis
Villasalto. Traces of gold were noted in the Ag rich Sarrabus Lode (Testa & Sartori, 1918). The As-Sb-Au mineralisation is hosted in shear zones and quartz veins, primarily at the contacts of shales with limestones or with meta-rhyolite. The veins crosscut the main deformation fabrics, and are probably related to late - to post - tectonic stage of the Hercynian deformation and magmatism. Metamorphic grade is lower greenschist. There is no obvious spatial relationship of mineralisation with granites, but felsic and lamprophyric dykes are common (Baker 1999). Recent mapping and sampling in the areas of old workings show potential for economic gold to be hosted in shoots along narrow veins. A new discovery at Monte Ollasteddu indicates potential for low grade, large tonnage disseminated and stockwork vein hosted As-Au mineralisation (Fig. 4).
6 CONCLUSIONS Both low sulphidation and high sulphidation epithermal gold and porphyry copper gold mineralization have only been recently discovered in Sardinia. In a short period of time numerous gold prospects have been identified and mining has commenced at Furtei. The Osilo vein field is in an advanced exploration stage. New areas of epithermal alteration are being recognised during the progression of exploration. Hercynian aged quartz veins mined in the past for As and Sb are also potential hosts for significant gold mineralisation. ACKNOWLEDGEMENTS The authors wish to acknowledge SGM for their permission to publish this paper. Thanks also to the many contributions of past and present members of the SGM geology team.
REFERENCES Baker, M. 1999. A review of the gold potential of Sardinia. Unpublished SGM S.p.A Report. Coulon, C. 1977. Le volcanisme calco - alcalin cenozoique de Sardaigne (Italie). Petrographie, geochimique et genkse des lavas andesitiques et des ignimbrites - signification geodynamique: Unpublished Ph.D.thesis, University of Aix - Marseille 111, 288 p. Dessi, R., Fiori, M., Garbarino, C., Grillo, S. M., Marcello, A. & S. Pretti 1996. Risorse minerarie metallifere associate alle vulcaniti Terziarie della Sardegna. Atti del Congress0 Internazionale per il Centenario dell 'Associazione Mineraria Sarda 3: 115 - 135.
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Fiori, M., Grillo, S.M., Marcello, A. & S. Pretti 1990. Depositi epitermali a metalli preziosi in Sardegna. Prime notizie sul sistema filoniano di S.Martino (Osilo - Codrongianus). L 'industria Mineraria 2:11- 14. Lattanzi, P. 1999. Epithermal precious metal deposits of Italy an overview. Mineralium Deposita 34: 630 - 638. Maxwell, M.G. 1995. Furtei Deposits Sardinia. Mineragraphic and geochemical characterizations of gold mineralization. Unpublished SGM S.p.A Report. Meloni, P. 1994. Metodologie di prospezione geomineraria e geofisica di mineralizzazioni sepolte tip0 porphyry copper. Studio del settore di Serrenti - Furtei (Sardegna meridionale): Unpublished Ph.D. thesis, University of Cagliari (English abstract in Plinius 13: 145-148. Ruggieri, G., Lattanzi, P., Luxoro, S., Dessi, R., Benvenuti, M. & G. Tanelli 1997. Geology, Mineralogy and Fluid Inclusion data of the Furtei High - Sulfidation Gold Deposit, Sardinia, Italy: Economic Geology 92: 1 - 19. Sillitoe, R.H. 1991 (a). Further comments on gold exploration at Furtei. Sa Pala de Sa Fae & Calabona, Sardinia. Unpublished AGIP Miniere Report. Sillitoe, R.H. 1991 (b). Exploration potential of the Siliqua porphyry copper prospects, Sardinia. Unpublished Progemisa Report. Simeone, R. & S. Simmons 1999. Mineralogical and fluid inclusion studies of low sulfidation epithermal veins at Osilo (Sardinia), Italy: Mineralium Deposita 34: 705 - 717. Stefanini, B. & A. William-Jones 1996. Hydrothermal evolution in the Calabona porphyry copper system (Sardinia, Italy): The path to an uneconomic deposit. Economic Geology 91: 774 - 791. Testa, L. & F. Sartori 1918. Tenori d'oro nei minerali di Sardegna. Bollettino dell 'Associazione Mineraria Sarda 9: 304 - 305.
Water-Rock Interaction 2001, Cidu (ed.), 02001 Swets & Zeitlinger, Lisse, ISBN 90 2651 824 2
Hydrothermal mineralization of Zr and other “immobile elements”: field evidence and experimental constraints S .Salvi. B .Tagirov & B .Moine Laboratoire de Mkanismes de Transferts en Gkologie, UMR-5563, Toulouse, France
ABSTRACT: It is becoming increasingly evident that the commonly immobile elements Zr, Ti, Th, A1 and REE do not always behave as such during water-rock interaction. Natural observations and mineralogical evidences suggest that fluorine play a fundamental role in this process. This is consistent with electrolytic solution chemical predictions that most of these elements form very strong aqueous fluoride complexes. In this contribution we present examples from alkaline igneous settings and metasomatic systems, where F-bearing hydrothermal fluids are suspected to have concentrated these elements to ore or potential ore grades. Preliminary results of Zr02 solubility experiments at 25 and 200°C show that in HF-bearing solutions aqueous Zr concentration can reach ppm values. Extrapolation of these data permits to model the above field occurrences and to show the importance of fluorine on “immobile element” transport. 1 INTRODUCTION High field-strength elements (HFSE) such as Zr, Ti, Th, REE as well as Al, have long been considered to be virtually inert during fluid-rock interaction, hence their label of “immobile elements” and their wide utilization as tracers of geological processes (e.g. Lipin & McKey 1989). Recently however, field observations gathered from a variety of environments, including REE deposits (Chao et al. 1997), HFSE, Th, A1 mineralizations in skams (Gierk, 1989, Moine et al. 1998), and HFSE deposits in alkaline magmatic complexes (Salvi & Williams-Jones 1996, Olivo & Williams-]ones 1999, Salvi et al. 2000), indicate that hydrothermal solutions can indeed remobilize and concentrate these metals. In addition, in most occurrences fluorine is found in the mineralized rocks. This evidence, together with the tendency of HFSE to form very stable fluoride aqueous complexes, suggests that fluorine may play a fundamental role on metal transport. Thus far however, lack of data on the stoechiometry and stability of F-HFSE complexes has prevented significant progress toward quantitative understanding of these processes. In this contribution we discuss a number of field occurrences from different settings where geological evidence indicates that HFSE were concentrated by F-bearing hydrothermal fluids. In addition, we present preliminary results on Zr solubility experiments, and use these data together with information from field occurrences where mobility of these elements and the role of fluorine could be estimated, to
roughly constrain the conditions that may promote hydrothermal mineralization of these metals. 2 GEOLOGICAL OCCURRENCES 2.1 Alkaline Igneous systems Of all geological settings, alkaline igneous rocks are among the most enriched in Zr and other HFSE, and commonly host exploitable resources of these metals. The Zr-Y-REE-Nb-Be mineralization in the lower-Proterozoic Strange Lake peralkaline complex (northern Labrador) is such an example. The ore zone occurs in granites and pegmatites that underwent extensive wall-rock alteration, characterized by Ca metasomatism as well as HFSE and F enrichment. A study of the post-magmatic history of this pluton showed that HFSE were concentrated hydrothermally (cf. Salvi & Williams-Jones 1996). The principal Zr ore mineral, gittinsite (CaZrSi207) occurs almost exclusively as a pseudomorph after primary elpidite. Fluid inclusions defining boundaries of these pseudomorphs trapped a moderately saline, Ca-rich meteoric fluid, circulating at 150-200°C. Bastnasite occurs as daughter mineral in these inclusions (Fig. la). The model proposed for HFSE mineralization involved exsolution of an orthomagmatic brine (-400°C) which leached the metals from p i mary minerals and transported them as fluoride complexes. Mixing of the orthomagmatic fluid with the cooler meteoric fluid responsible for calcium metasomatism, promoted HFSE precipitation. 745
contribution of fluorine in their transport. In this area outcrop a large number of diopside-phlogopite skams formed at 800°C and 5-10 kb (Moine et al. 1985). To the east (Tranomaro area), the skams host several Th deposits (mainly thorianite) and are enriched in Zr (baddeleyite, zircon, zirconolite, Zr-rich diopside) and REE (in hibonite, CaA112019). These mineralizations contrast sharply with similar skams further west (Ampananihy area) which contain only background HFSE. Microprobe analyses of phlogopite and biotite from the Tranomaro area show extremely high fluorine contents (XI: = 0.50.8), corresponding to LIHFICLI-I~O in the fluid of 10-O.' to 1O-O.'. In contrast, micas in Ampananihy skarns have F-contents ranging within normal crust values (ie., about 1 log unit below QWF, cf. Fig. 2). In the Tranomaro region also occur the huge Andranondambo sapphire deposits, formed during a lower-temperature metasomatic event (500"C, 2 kb). In these rocks, phlogopite compositions (X,= 0.40.5) again indicate high fluorination of the fluid (S in Fig. 2).
Figure 1. SEM images of bastnasite (A) and parisite (B) in opened fluid inclusions from the Strange Lake and Tamazeght complexes, respectively.
A similar occurrence is the sub-economical Zr mineralization hosted in the Eocene Tamazeght alkaline complex (Morocco). Here, primary zirconosilicates are replaced by secondary Ca- and F-rich phases in agpaitic syenites and pegmatites. Nepheline crystals adjacent to these mineral replacements trapped saline Ca-rich fluid inclusions at -3OO"C, which contain several HFSE daughter minerals, including parisite (Fig. lb), zircon, and a Ti silicate. This fluid is interpreted to have evolved from pegmatite crystallization ( 5 500°C) and probably transported HFSE as fluoride complexes (Salvi et al. 2000). Metal deposition took place upon interaction of the fluid with surrounding limestone and marble.
3 EXPERIMENTAL STUDY The principles of solution chemistry suggest that hard Pearson's metals (cf. Pearson, 1963) like HFSE favor strongly electronegative anions (e.g. OH-, F-, P043- vs. Cl-). The paucity of thermodynamic data available for Zr plus the evidence obtained in the field suggesting a role of fluorine, encouraged us to investigate aqueous Zr-F complexation. Solubility of ZrOz (baddeleyite) was studied in HCl and HF solutions at 25 and 200°C and P,,,. Ex-
2.2 S k i m system The metamorphic terrane of southern Madagascar (granulite facies) contains fine examples of HFSE mobility in metasomatic rocks, and evidence for the T ("C)I P (kbar)
2 .o
1.o
1.5
1000/T (K) Figure 2. XF isopleths in biotite (Xkig = 0.75), QWF ( c i H 2 0 = l), and fluorite stability (in COz-free fluids) along a 150°C/kb thermal gradient (calculated using the model of Zhu & Sverjensky 1992). Also shown are mica compositions for the mineralized and barren skarns of Madagascar. The diagram illustrates how the presence of F-bearing minerals can limit nF in a fluid.
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250
200
4 GEOCHEMICAL CONSIDERATIONS -
T = 200°C pH=4to5
cr, 150
E
2
100 50
m
L
1000
2000
3000
4000
PPb Zr Figure 3. Equilibrium aqueous Zr contents of experimental runs at 200°C as a function of F content of solution.
periments were carried out using 3 80-cm3 pure-Ti (UT40) Prolabo rocking autoclaves, fitted with a water-cooled Ti valve for sampling during runs. A number of experiments were performed in 20-cm3 batch Ti autoclaves that were water quenched. Teflon containers were used for room-temperature runs. All autoclaves were passivated with a 0.1 m HF solution at 300°C for a few weeks. Experimental solutions and pH standards were prepared from doublydeionised H20 and Pro Analysi' Merck reagents. Background Zr concentrations were checked before each run. Experiments were replicated varying the proportions of solution and ZrO2 powder. During sampling at 200°C experimental solutions were filtered in situ through a 10-pm Ti filter and again through a 0.05-pm cellulose nitrate filter at 25°C. Filtration for solutions from 25°C and batchautoclave experiments were carried out only with the 0.05-pm cellulose nitrate filter. Zr and F were analyzed by ICP-MS and HPLC, respectively, using matrix-matched standards, and pH was measured with an Orion combined electrode. Preliminary results show a strong increase in aqueous Zr with fluorine concentration. In the absence of fluorine, Zr concentration is below the detection limit (0.01 ppb) for each temperature even in 0.1 m HCl. In contrast, Zr concentration increases to 10-3m at 25'C and 10-4m Zr at 200°C in 10-2m HF. Figure 3 illustrates dependence of Zr concentration on HF content for 200°C and P,,,. As aqueous Zr increases proportionally to the square of HF concentration, Zr(OH)2F20dominates Zr speciation in the experimental solutions. Results obtained for 25 and 200°C show a strong increase in aqueous Zr with fluorine concentration from nil in the absence of F (even at pH = I ) to 10-3m Zr in 10-2m HF at 25'C. Figure 3 illustrates this effect for 200°C and P,,,.
Based on mass balance computations, Salvi & Williams-Jones (1996) calculated that some 5500 ppm Zr were added to the rock during alteration at Strange Lake. This implies a maximum solubility of -0.05 m Zr in the fluid (for waterhock 11). Mobility of REE was indicated by the presence of bastnasite daughter minerals in fluid inclusions. At Tamazeght, volumetric proportions of HFSE-bearing daughter minerals provide estimates for Zr and REE in solution ranging from 0.01 to 0.1 m. The above values are extremely elevated, but only about 0.5 log units higher than those calculated by Vard & Williams-Jones (1993) for Zr in an alkaline fluid at 400"C, or than REE contents measured in leachates from orthomagmatic fluid inclusions (Banks et al. 1994). Extrapolation of the experimental data presented above indicates that aqueous Zr concentrations such as observed in the fluids at Strange Lake and Tamazeght can occur in the presence of about 10-2to 10-' m F in the form of Zr(0H)zF;. However, it can be seen from examination of Figure 2 that at 350 to 500°C a fluid in equilibrium with fluorite could never contain such amounts of aqueous fluorine. It follows that the ore-leaching fluids at Strange Lake and Tamazeght did not contain Ca (to form fluorite), which is consistent with the composition of evolved alkaline magmas, and with the mode of metal deposition envisaged for the two complexes. In the mineralized skarns of Madagascar, Th solubility was about eighty times higher than in the barren skams (Moine et al, 1985). Formation of ThF aqueous species could account for such a solubility increase and, by analogy with Zr (cf. Pearson 1963), Th(OH)2F2' was likely the main Th complex in solution. Fluorine contents of micas in these skams are consistent with high a~ in the fluid. A pH increase during reaction of the fluid with calcite in marble, to produce diopside and other skarn minerals, may have triggered thorianite precipitation, by destabilizing the Th hydroxy-fluoro complex. Mobility of aluminium and the origin of the Andranondambo sapphire deposits may be explained b the formation of mixed species like Al(OH)2F(7, AlOHF: (?sky 1994), NaA1(OH)3Fo and NaAl(OH)2F2 (Tagirov et al., 2001). Presence of these complexes leads up to 1000-fold higher solubility of Al-bearing minerals in the metasomatic fluid than in pure water. 5 CONCLUSIONS It is conventionally accepted that refractory elements like Zr, Th, Ti, REE and A1 are not effected during wall-rock alteration or water-rock interaction in general. However, field evidence from diverse geologi747
cal settings show that these metals can be remobilized and even concentrated to ore grade by hydrofluids. themodPamic data and the preliminary Zr solubility data presented above indicate that fluoride complexing enhances the solubility of these metals by several orders of magnitude. Metal mobility is expected in fluids circulating in Ca-poor systems where F contents are not buffered by mineral-fluid exchange reactions, or at high temperatures (> 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
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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.
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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
,
-
'
'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.
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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
