Speleogenesis

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Speleogenesis

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Speleogenesis
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Speleogenesis
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Klimchouk, Alexander B. (Aleksandr Borisovich)
Ukrainian Institute of Speleology and Karstology
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No. 8 (2005)

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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info O htin Aragonite Cave (Slovakia): morphology, minera logy and genesis Pavel Bosk1, Pavel Bella2, Vclav Cilek1, Derek C. Ford3, Helena Hercman4, Jaroslav Kadlec1, Armstrong Osborne5 and Petr Pruner1 1 Institute of Geology, Academy of Sciences of the Czech Republic, Rozvojov 135, 165 02 Praha 6-Lysolaje, Czech Republic; E-mail: inst@gli.cas.cz 2 Slovak Caves Administration, Hodova 11, 031 01 Liptovsk Mikulš, Slovak Republic; E-mail: bella@ssj.sk 3 School of Geography and Ge ology, McMaster University, 1280, Main Street West, Hamilton, Ontario L8S 4K1, Canada; E-mail: dford@macmaster.ca 4 Institute of Geological Sciences, Polish Academy of Sciences, Twarda 51/55, 00-818 Warszawa, Poland; E-mail: hhercman@twarda.pan.pl 5 Faculty of Education, A35, University of Sydney, N.S.W. 2006, Australia; E-mail: a.osborne@edfac.usyd.edu.au Re-published from: Geologica Carpathica 2002, 53 (6), 399-410 Abstract Ochtin Aragonite Cave is a 300 m long cryp tokarstic cavity with simple linear secti ons linked to a geom etrically irregular spongework labyrinth. The limestone s, partly metasoma tically altered to ankerite and siderite, occur as lenses in insoluble roc ks. Oxygen-enriched meteoric water seeping along the faults caused siderite/ankerite weathe ring and transformation to ochres that w ere later removed by mechanical erosion. Corro sion was enhanced by sulphi de weathering of gangue mine rals and by carbon dioxide released from decomposition of siderite/ankerite. The initial phreat ic speleogens, older than 780 ka, were created by dissoluti on in density-derived convectiona l cellular circulation condi tions of very slow flow. Thermohali ne convection cells operating in the flooded cave might also have infl uenced its morphology. Later vadose corrosional ev ents have altered the original form to a lar ge extent. Water levels have fluctuated many times during its hist ory as the cave filled during wet periods and then slowly draine d. Mn-rich loams with Ni-bearing asbolane and birnessite were formed by microbial precipitation in the ponds remaining after the floods. Allophane was produced in the acidic environment of sulphide weathering. La -Nd-phosphate and REE enriched Mn-oxide precipitated on geochemical barriers in the asbolane layers. Oc hres containing about 50 wt.% of water influence the cave microclimate and the precip itation of secondary aragoni te. An oldest aragonite generation is preserved as corroded relics in ce iling niches truncated by corrosional bevels. TIMS and alpha countin g U series dating has yielded ages of about 500-450 and 138-121 k a, indicating that there have been several episodes of deposition, occurring during Quaternary warm periods (Elsterian 1/2, Eemian ). Spiral and acicular forms representing a s econd generation began to be deposited in Late Glacial (14 ka – Allerd) times. The youngest aragonite, frostwork, continues to be deposited today. Both of the younger generations have simi lar isotopic compositi ons, indicating that they originated in conditions very simila r, or identical, to those found at present in the cave. Keywords: Slovensk rudohorie Mts., Ochtin Aragonite Cave, cav e morphology, speleogenesis, mi neralogy, aragonite, U-series dating Introduction Ochtin Aragonite Cave is unique among the 4,250 known caves in Slovakia, although with only 300 m of passages it is relatively small (Fig. 1). The cave was discovered in 1954 during the excavation of an adit (Kapusta Gallery) for iron ore exploration. The mine workings also intersected other, smaller caves but none were so interesting or significant. The cave was opened to the public in 1972 and in 1995 was included in the UNESCO World Heritage List as a component of the Caves of the Slovak and Aggtelek Karsts.

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 Fig. 1. Geomorphological map of the Ochtin Aragonite Cave, showing typical cross-sections (after Bella, 1998, modified) and the sediment section in Ovlna Passage with positions of the palaeomagnetic samples (black squares) and the magnetostratigraphic results (black – normal polarised magnetozone, white – reverse polarised magnetozone; for explanations see the text). Corrosion-denudation forms: Planar speleogens: 1 – horizontal solutional ceilings (Laugdecken); 2 – inclined planar walls of passages and halls descending to the floor (planes of repose, Facetten); 3 – inclined, more or less planar walls of passages and halls with smaller corrosion convex and concave forms; Concave speleogens: 4 – shallow oval irregular spoon-like depressions on roofs and walls; 5 – deeper distinct oval irregular depressions on roofs and walls; 6 – distinct, mostly horizontal niches; 7 – cupola-shaped depressions in roofs; 8 – shallow elongated channelshaped forms in roof; 9 – horizontal elongated notches on walls; 10 – blind lateral tube-like holes; 11 – rocky windows in bedrock; 12 – narrow steep corrosion cavities developed along prominent fissures; 13 – horizontal shallow tr ough-like depressions; 14 – tubular karren; 15 – fissure karren on collapsed blocks; 16 – shallow drip-holes on collapsed blocks; Convex speleogens : 17 – large irregular bedrock protrusions in roofs; 18 – structurallycontrolled large elongated roof bedrock juts on roofs; 19 – less pronounced elongated roof bedrock juts on roofs controlled by bedding; 20 – bedrock pendants; 21 – bedrock blades; 22 – elongate, indistinct and irregular bedrock protrusions along walls; 23 – elongated bedrock protrusions above horizontal corrosion notches; 24 – oval bedrock protrusions in floors; Structural-tectonic forms: 25 – smooth breakdown surfaces without corrosional relief; Depositional forms: 26 – sediment sequences; 27 – cones and banks of sediments at the foot of walls; 28 – planar accumulation surface; 29 – piles of collapsed blocks; Erosion forms: 30 – meandering channel on flat accumulation surface; 31 – dripholes; Other: 32 – trail (in plan); 33 – lake; 34 – planes of repose with thin cover of ochres; 35 – ochres; 36 – aragonite; 37 – stalagmite; 38 – trail (in profile). The cave is located in the NW shoulder of Hrdok Hill (809 m a.s.l.) in the Revcka vrchovina Highlands, a part of the Slovensk rudohorie Mts., Ro ava District. Caves there are developed in steeply dipping metalimestone lenses of variable size surrounded by phyllites of the Drnava Formation (Late Silurian to Early Devonian; Gelnica Group; Bajank and Vozrov et al., 1983; Ivani ka et al., 1989). They were folded and metamorphosed during the Variscan Orogeny. Some of the faults and fissures were rejuvenated during the Alpine Orogeny. Portions of the limestone have been metasomatically altered to ankerites and siderites by Mg and Fe-bearing

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 hydrothermal solutions (Mišk, 1953) ascending particular fissures (Droppa, 1957). The hydrothermal activity was associated with the emplacement of the Gemericum granites (Andrusov, 1958), which have been dated to 96 Ma (Kantor in Homza; Rajman and Roda, 1970). No younger hydrothermal activity has been recognised in this region (Gal, 1996). The cave is structurally guided, with N-S, W-E and SW-NE trends (Rajman et al., 1990; Gal, 1996; see Fig. 1). About 15 other caves of the Ochtin cryptokarst type have been intercepted by the Kapusta Gallery (Gal, 1996), some of them containing aragonite speleothems similar to those in Ochtin Aragonite Cave. Other caves are found in the vicinity, as well, but they differ substantially from the Ochtin cryptokarst in form (Gal, 1998). Previous work Droppa (1957) compared the tube-like cave passages to the erosion forms produced by typical flowing streams underground. Aggressive corrosion by meteoric waters percolating along tectonic fissures was the main agent in the development of the cave. Eroded products from the chemical weathering of the ankerit es were deposited in the lower parts of cave, obstructing drainage outlets there. Gal and Ženiš (1986) argued that percolating meteoric waters first oxidised the ankerite to create iron hydroxides – ochres; mechanical erosion of the ochres then produced the larger voids. The general shape of the cave thus is that of the original metasomatic ankerite bodies in the limestone, with later modifications resulting from some subsequent dissolution of the limestone, partly under phreatic conditions. In addition to the oxidati on of siderite/ankerite, Rajman et al. (1990, 1993) stressed the contribution of other mineralisation to the development of the karst. Oxidation of the a bundant gangue minerals in the surrounding rocks (chiefly pyrite) increased corrosional aggressivity of percolating waters by producing H2SO4. Gal (1996) also supposed that limestone corrosion and ankerite oxidation and mechanical washout were the main speleogenetic agents. He considered that during periods of higher rainfall the limestone lenses, hydrogeologically isolated by the phyllites, could become temporarily flooded with water. Despite the abundance of distinctive corrosion forms in the cave, little has been written about their genesis and the hydraulic conditions under which they may have formed. Droppa (1957) mentioned effects of hydrostatic pressure when the open cavities were completely flooded. Gal and Ženiš (1986) and Gal (1996, 1998) argued that the cave formed under phreatic conditions. Bella (1997, 1998) was the first to recognise that the planar solution roofs (bevels, Laugdecken), planes of repose (Facetten) and longitudinal wall notches are particularly important. Cupola-shaped depressions in the roof originated by convective processes in the water. He also recognised the dominant role of phreatic and stagnant vado se waters in the caves evolution at times when the carbonate lens was water-saturated ( cf. Ford and Williams, 1989, pp. 294-308). Due to the difficulty of explaining the origin of concave corrosion forms and the development of passages with oval cross-sections, Choppy (1994, following ideas of Nicod, 1974), suggested that Ochtin Aragonite Cave evolved as a result of hydrothermal processes. Gal (1996) contended that hydrothermal processes during the Upper Cretaceous operated at much greater depths, however, and that the accelerated Tertiary and Quaternary meteoric karst corrosion completely overprinted traces of any earlier hydrothermal activity. Results of detailed geomorphological research do not support a hydrothermal genesis for the surviving initial forms (Bella, 1998). Clek et al. (1998) stressed the nothephreatic origin of some of these morphologies ( cf Jennings, 1985). No hydrothermal minerals have been detected in the cave and the aragonite deposition was not related to hydrothermal conditions (see Clek and Šmejkal, 1986; Rajman et al., 1990, 1993). The presence of fresh, unweathered corrosion forms led Droppa (1957) to propose that the caves were relatively young, with speleogenesis occurring at the beginning of the Holocene. Kubny (1959) suggested that the caves originated during Quaternary glacials. Weathering of the ankerites and successive exhumation and erosion of the ochres began in the Tertiary and has been active ever since in the view of Homza, Rajman and Roda (1970), Rajman et al. (1990). The first U series

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 dating of samples of the aragonite (Ford, unpublished, cited in Rajman et al., 1990, 1993) gave two ages, one of 138-121 and the other of 14 ka B.P. that rule out a Holocene or latest glacial origin. Morphology The morphology of the cave was described by Šev k and Kantor (1956) and Droppa (1957). It consists of simple linear sections linked to a geometrically irregular sponge-work labyrinth (Bella, 1995). Detailed geomorphological mapping has defined the principal and smaller morphological forms (for location and list of forms, see Fig. 1). The forms described hereafter are given non-genetic descriptive names owing to the fact that some them cannot be correlated with any commonly applied terms (e.g., those of Slabe, 1995). The cave consists of two genetically different types of voids or principal speleogens: (1) high and narrow linear fissures (e.g., from Vstupn Hall to Mramorov Hall; Fig. 1), and (2) a labyrinth of passages and chambers with oval cross-sections. Bedrock corrosion forms are the most abundant type of speleogens. Structural-tectonic forms, clastic sedimentary de positional and erosional forms are less frequent (Fig. 1). The corrosional speleogens are products of the enlargement of the caves, occurring on the floors, walls and roofs of all passages and chambers. They can be classified by their geometry into: planar, concave and convex types. The principal planar speleogens are horizontal solutional ceilings (Laugdecken sensu Kempe et al., 1975 or bevels sensu Ford and Williams, 1989) and inclined planar walls descending to the floors of passages and halls (planes of repose sensu Lange, 1963; Goodman, 1964 or Facetten sensu Kempe et al., 1975; Fig. 2/3). The predominant concave speleogens are pronounced, more or less closed oval cupolashaped depressions in roofs (Fig. 2/1). Horizontal concave notches extending along walls and convex bedrock prominences just above them indicate positions of long-lasting paleo-water levels. In addition there are elongated shallow trough-like depressions and tubular karren produced by flowing water e.g., in Vstupn and Mramorov Halls. Corrosion bevels are developed at three different levels within the cave. The highest is preserved in Ovlna Passage. Lower bevels occur in Jeovit Passage and Aragonitov zhrada and 2 m below the roof in Ovlna Passage. Near the junction of Jeovit Passage and Hlbok Hall, there is also a bedrock pendant truncated by bevelling 0.4 m below the roof level. The lowest bevels correspond with the low roof level in Aragonitov zhrada near Hlbok Hall. This indicates that the retention level of stagnant water in the cave has fallen over time and/or oscillated substantia lly at different times. Flat surfaces (smooth joint planes without any corrosional relief) produced by breakdown along structural discontinuities (Mramorov Hall) represent structural-tectonic forms. Depositional forms (sediment sequences, piles of collapsed blocks, etc.) developed when clastic sediments were deposited or removed from the caves: e.g., horizontal accumulations with desiccation cracks, deposited by periodic floods; alluvial cones of infiltration sediments fr om percolating water (Vstupn Hall). Small depressions (small meandering channels, drip-holes) resulting from the erosion of clayey sediments by flowing and dripping water in Vstupn Hall represent clastic erosional speleogens, which are less important in the cave. Mineralogy Methodology Twenty three samples of ochres, clays, broken aragonite speleothems and neomorphic aragonite were collected in the cave. Twelve typical samples were studied by SEM and analysed on 60 points by EDAX (LINK connected to a JEOL-JXA-50A Microprobe). A total of 34 X-ray diffraction analyses were made (Philips Diffractometer PW 3710). Powder produced for the X-ray work was also analysed by microprobe. Mn oxides were separated by sieving, and in a settling column using deposition times ranging from 2 hours to 8 days. Individual portions were structurally analysed. Mn oxides, goethite and allophane were analysed by DTA and TG (TG-750 Stanton-Reford, University of Chemical Technology, Prague). Carbon and oxygen stable isotope ratios were measured with a Finnigan MAT 251 Mass Spectrometer (Czech Geological Institute, Prague). Water content in ochres was calculated from weight loss at 70 oC. All analyses, except where otherwise mentioned, were carried out at the Institute of Geology, AS CR Prague. Other speleothems were visually examined in the cave with a portable UV-lamp (253 and 360 nm).

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 Fig. 2. Photographs of typical forms and speleothems in the cave. 1 – cupola-shaped depressions in the roof of Hviezdna Hall; 2 – aragonite of the oldest generation truncated by bevels in Hlbok Hall; 3 – cross-section of the passage between Jeovit and Aragonitov Passages, showing planes of repose; 4 – the sedimentary profile in Ovlna Passage; 5 – aragonite of the second generation in Ovlna Passage; 6 – the youngest aragonite on ochres in Jeovit Passage (photos 1 to 4 and 6 by P. Bella, photo 5 by A. Lucinkiewicz). Goethite The ochres are soft and moist, containing 47 to 56 % of water by weight. They formed from weathered ankeritic and si deritic metasomatites and also cover cave walls as irregular crusts deposited from waters. Goethite is present as an extremely fine-grained to cryptocrystalline, although not amorphous, matrix in the ochres and as fine acicular forms (several m) „floating“ in finegrained ochre. The moisture content of the ochres distinctly influences the humidity of the air analysed in the cave, as the ochres function as a

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 humidity exchanger, able to adsorb and release water vapour. The ochres contain irregular laminae of birnessite, a Mn oxide. In addition to inclusions liberated by weathering (quartz), the ochres contain clay minerals (muscovite 2M1, illite, probable chlorite and 14.8 smectite) that were deposited on the rock surface in flooded cave conditions. Some of the ochres contain greater concentrations of P2O5 (0.3 to 1.0 %; Tab. 1). Asbolane Black Mn ochres occur as an admixture in the Fe ochres and other cave fills. They are derived from the ankeritic metasomatites, which contain about 2 % MnO. The sequence in Ovlna Passage (Fig. 2/4) is composed of very fine-grained massive brown clay (a mixture of goethite and clay minerals) and includes a layer about 300 mm thick that is composed of several bands of Mn ochres with abundant intercalations of white allophane and redeposited Fe ochres. As bolane is abundant here as soft, black, earthy material with a clayey appearance. Complete samples and various grainsize fractions were analysed. There were problems of exact identification due to structural disordering, the almost amorphous nature of the mineral, and from coalescing diffraction lines. Submicron-sized plates of muscovite 2M1 remained in the sample even after extended sedimentation, masking other diffuse diffraction lines. Asbolane comprise about 30 to 40 % of the black fills. It is usually accompanied by muscovite, and also by quartz, goethite, allophane, birnessite, apatite, anatase and, more rarely, by rutile and authigenic La-Nd-bearing phosphate. Nickel contents can reach 1.9 to 3.9 % (Table 1), while the magnesium content is relatively stable but can be locally enriched (2.4 to nearly 10 %). Similar variability was detected in P, Ba (0.4 to 1.4 %) and the rare earth elements (REE). Sometimes the asbolane consists of microscopic globules of Mn-oxide covered by fine fossilised organic filaments, indicating the microbial conversion of Mn2+ to Mn4+. The asbolane layers probably result from bacterial precipitation in shallow residual pools as the cave is in the late, very slow, stages of draining episodic flood waters (see also Andrejchuk and Klimchouk, 2001). Fig. 3. X-ray diffraction of the asbolan layer. A – asbolan, Q – quartz, M – muscovite, G – goethite. Birnessite Birnessite occurs as a soft black substance. It cannot be distinguished optically from the asbolane. Birnessite was identified both in Fe ochres (as fine darker coloured and irregular bands) and in the asbolane layers where it is probably a product of maturation of asbolane. Allophane Allophane was found only in the asbolane deposits, as separate white, fine-grained earthy layers 30 to 80 mm thick disintegrating into cubes, or as admixtures within the asbolane. Allophane was identified by chemical analyses, X-ray diffraction and particularly by DTA and TG analyses (Fig.4). Allophane is an uncom mon mineral in karst caves (Hill and Forti, 1997, p. 179-181), but nevertheless, is relatively abundant in speleothems growing in abandoned mines. It has also been found in pseudokarst fissure caves (Clek, Langrov and Melka, 1998). Allophane commonly forms in the acidic environment produced by weathering of sulphides in the surrounding rocks. Its occurrence in limestone environments that are usually associated with high pH may therefore appear somewhat surprising. However, its presence is a strong indication that sulphide weathering played a role during speleogenesis of the cave. Halloysite A mineral of the kaolinite group, structurally similar to halloysite, occurs as an indistinct admixture in the allophane. It was detected by Xray diffraction. We presume that it formed either by maturation of the allophane or that the allophane was formed by the transformation of weathering products containing minerals of the kaolinite group.

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 TABLE 1 Representative chemical analyses of selected minerals Sample/Oxide ochre asbolane allophane muscovite La-Nd-bearing phosphate SiO2 4.31 23.88 51.14 57.67 14.09 TiO2 0.05 nd 0.19 0.40 nd Al2O3 2.48 24.24 41.16 33.45 13.71 Fe2O3/FeO* 84.01 *2.14 1.29 0.75 *5.94 Na2O 0.16 1.03 0.08 0.11 1.04 K2O 0.27 0.10 0.20 11.86 0.20 CaO 0.57 1.22 3.29 0.08 4.01 MgO 1.80 13.28 2.27 1.91 2.37 MnO 6.34 29.65 0.37 0.23 27.8 BaO nd nd nd nd 1.40 NiO 3.94 3.94 nd nd 2.87 P2O5 0.52 0.52 nd nd 5.46 La2O3 nd nd nd nd 12.21 Nd2O5 nd nd nd nd 8.91 nd not determined Fig. 4. DTA/GTA curves of asbolane (a) and allophone (b) from the Ovlna Passage (velocity of heat 10 oC min1, air 10 ml.min-1, sensitivity DTA 10V(10 mV). Anatase Authigenic cryptocrystalline anatase forms amoeba-like patches of cement in small fragments of brown ferruginous clayey siltstones/sandstones. The fragments represent relics of kaolinitisation products washed down into the cave. Anatase was detected by chemical analyses and on the basis of morphological comparison with samples from the Czech Karst (Clek, 1989; Clek and Bedn # ov, 1993). Apatite Authigenic apatite forms irregular thin laminae, less than 1 mm thick, and irregular amoeba-like impregnations in the asbolane profiles. It was detected by chemical analysis and X-ray diffraction. Migration of Ca-phosphate requires an acidic environment. Phosphates, other than those derived from guano, precipitate in limestone from relatively acidic surficial P-enriched solutions, which have been leached from soil and weathering profiles. Accessory minerals Muscovite 2M1 is a very common accessory mineral, detected in samples by X-ray diffraction, then separated and chemically analysed. The typical chemical analysis is given in Table 1. Muscovite occurs as very fine-grained plates. It is probably derived from phyllites. Quartz occurs as angular and corroded silt-sized grains. Its character

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 indicates that it was derived from dissolution of limestones and comes from the immediate surroundings rather than being transported over a long distance. Acicular rutile was found in some places associated with anatase cement. La-Nd-bearing phosphate Detailed sampling of the asbolane layer revealed places with increased P content in the form of apatite (X-ray identification). In other places the P content was slightly greater than Ca content. Nevertheless, high La (La2O3 up to 12.21 %) and Nd contents (Nd2O5 up to 8.91 %) occurred in similar positions repeatedly. The La-Nd-bearing phases are very fine-grained (tens of m) and cannot be macroscopically distinguished within the black asbolane. The chemical composition of the phosphate-asbolane layer with high REE and Ba contents is listed in Table 1. The REE regularly occur in association with high P concentrations. In some places the REE concentration is higher than the P contents, so the relationship between REE and Mn oxides has to be taken into account. A number of REE released by weathering (presumably of volcanics) can migrate efficiently within carbonate sequences. They form authigenic minerals only with difficulty, but they can be fixed in finely dispersed phosphate or in Mn oxides. Due to local permeability and diversified sources, apatite has formed only in some parts of the karst fills. In other places the possible effect of phosphate molecular sieving led to formation of REE-bearing phosphates. The excess of the REE concentration became bound to Mn oxide. Aragonite Aragonite speleothems are the outstanding feature of this cave. According to Rajman et al. (1990, 1993), aragonite speleothems have been traditionally classified according to their morphology into: kidney-shaped, acicular and spiral (i.e. flos ferri ) forms. Clek et al. (1998) identified three generations of aragonite speleothems according to their age and/or relationship to the speleogens. The oldest aragonite generation occurs as massive, whitish milky-coloured, kidney-shaped forms and irregularly corroded relics with polyhedral appearance, rarely more than 300 mm thick (Fig. 2/2). The aragonite is highly and irregularly recrystallised and corroded. Finegrained parts are still composed of aragonite (radial-fibrous aggregates), with some calcite (blocky mosaic). Recrystallised patches consist of calcite, with some aragonite, mica and quartz (Xray analysis). Fine box-work structures or very fine dogtooth-like crystals cover walls of corrosion voids. Some voids are filled with younger milkywhite finely radial-fibrous aggregates of aragonite or by mica-rich sediment (X-ray detection). Long duration phosphorescence (up to 5 s.) after illumination by UV-lamp differentiates the oldest generation from the younger one. Plane solution roofs (bevels) commonly truncate the oldest aragonite. The second generation of aragonite occurs as long needles and helictites, so-called acicular and spiral forms, up to several hundred millimetres long (Fig. 2/5). It displays fluorescence, but no phosphorescence in UV-light. The aragonite needles are sometimes associated with globular opal. Crystal faces do not display any corrosion, even under high magnification (x30 to 500). Microscope and field observations indicate that this generation of aragonite has been growing continuously up to the present time, explaining its bright white colour and fresh appearance. The youngest aragonite generation has not previously been detected due to its tiny size (Fig. 2/6). It occurs as fine fan-like forms with diameter of 2 to 4 mm (sometimes more) and as miniature helictites with lengths not exceeding 40 mm. The helictites usually grow from centres of radial aggregates. These forms – “frostwork” – grow on soil and Fe ochres, typically above the lake in Hlbok Hall. Here, aggregates cover thin coatings of loam and ochre deposited from stagnant water. There was inhomogeneous glowing, with greenish and bluish points and phosphorescence of 1 to 2 seconds duration appeared when the aragonite was illuminated by the UV-lamp. Aragonite genesis Two principal factors caused deposition of aragonite in Ochtin Aragonite Cave: (1) high concentrations of Mg, Fe and Mn ions in the karst solutions, and (2) a closed and deeply-seated, partly flooded cave environment with a high proportion of the walls covered by moist Fe ochres.

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 The ochres act as a humidity exchanger between the walls and the cave atmosphere. Ochtin aragonite occurs most frequently on substrates with water rising by capillary action or with very slowly percolating water on moist sediments, which slowly release water vapour into the cave atmosphere. A similar situation is also observed in Zbrašovsk Aragonite Caves (Czech Republic). The isotopic ratio of carbon in the aragonite, "13C, varies between -7.4 and -6.0 ‰ (PDB). Oxygen isotopic ratio ( "18O) was found to range between -7.0 and -6.3 ‰ (PDB). The C and O isotopes thus are within the range typical for the slow, equilibrium release of CO2 from solution. The graph in Fig. 5 compares aragonites from Ochtin Aragonite Cave with calcites and aragonites from Star hrad Cave (Nzke Tatry Mts., Slovakia). Ochtin aragonites plot within the field of the lowest values of the calcite spelothems from Star hrad Cave, but distinctly away from the Star hrad aragonites, indicating different depositional conditions. While aragonite from the wellventilated Star hrad Cave shows isotopic equilibrium with common atmospheric carbon dioxide, the Ochtin aragonite, from a closed environment, shows an anomalous isotopic composition probably caused by a different cave CO2 composition. Rapid kinematic processes and evaporation can be excluded from any role in the deposition of the aragonite with a high degree of certainty. The genesis of the aragonite in Ochtin Aragonite Cave cannot be completely explained until data from direct measurement of CO2 concentrations and isotopic composition in the cave air are available. Nevertheless, the isotopic data show one important result the neomorphic aragonite of the youngest generation has an identical isotopic composition that is nearly identical to as the acicular aragonite of the second generation both types thus were deposite under the same conditions, which are very similar to the modern ones. U series ages of the aragonite deposits Samples of aragonite and calcite have been dated in two laboratories – at McMaster University, Hamilton, Canada (Ford) and at the Institute of Geological Sciences, Polish Academy of Sciences, Warsaw, Poland (Hercman). Two applications of the 230Th/234U method (Ivanovich and Harmon, 1992) were adopted, the first using the older standard means of estimating the ratio of the two isotopes by counting radioactive disintegrations (alpha particles) by scintillometry; the second using the modern method of direct isotope counting in a mass spectrometer (TIMS – thermal ionisation mass spectrometry; Li et al., 1989). Mr. Štefan Roda sen. collected samples of the aragonites in Ovlna Passage and Hlbok Hall at 1989 or 1990 (locations on the map in Rajman et al., 1990). Ford dated two of them in 1990 (90/Och1 and 90/Och2). All samples had a high uranium content (as is typical in aragonite) and negligible amounts of detrital thorium, and thus yielded precise and unambiguous ages. Results were partly published by Rajman et al. (1993). Fig. 5. The isotopic composition of calcite and aragonites from the well ventilated Star hrad Cave in the Nzke Tatry Mts. compared with the closed deeper system of Ochtin Aragonite Cave. Very slow evaporation and isotopic equilibrium fractionation is proposed for the samples from Ochtin (AII and AIII – aragonite generations).

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.10 Sample 90/Och2 was a portion of aragonite flowstone broken during tr ail construction. It was 20 mm in thickness, clean and opaque white, with a coconut meat texture. For alpha dating it was cut into three slices of ~7 mm thickness each (samples Och2A, Och2B and Och2C) representing the top, middle and base of the deposit respectively. All three analyses yielded U contents of 8-10 ppm. Thorium was lost from Och2C during the extraction, with the result that no date could be obtained. Sample Och2B from the middle third of the flowstone gave an age of 138,000 # 7,000 years BP, where # 7,000 is the one standard deviation statistical counting error. Sample Och2A from the top one third of the flowstone gave an age of 121,000 # 6,500 years. If sample Och2 grew at a constant rate between Och2B and Och2A, then the accumulation rate was ~0.41 mm/1,000 years. If it is further assumed that all of the deposit grew at this constant rate, then its growth commenced about 162,000 years ago and ceased at approximately 115,000 years. Sample 90/Och1 consisted of three broken aragonite spiral helictites (“needles” or “whiskers”), all measuring about 60 mm in length and tapering from ~3 mm external diameter at the base to ~2 mm at the tip. They contained central canals for water flow but the tips were sealed. The needles appear to be extending by fluid permeating out and precipitating in the region of the tips and it was supposed that the latter were modern. The observed sealing of the tips might possibly be a consequence of the artificial opening of the cave changing the microclimate. The basal 15 mm of one needle were analysed by the alpha method in 1990 (sample 90/Och1), yielding the remarkable U content of 15 ppm and an age of 13,600 # 500 years (one standard deviation). As a check, a second needle was analysed by mass spectro metry in 1995. The basal 15 mm of growth contained 16 ppm U and yielded an age of 13,30068 years (two standard deviation error). No age could be derived for the top 15 mm (18 ppm U) because it had insufficient thorium; this implies that it is both very clean (no detrital contamination problems) and young, supporting the assumption that the very tip is modern. From these measurements we establish that the needle grew ~52.5 mm in 13,300 years, giving a mean extension rate of about 4 mm per 1,000 years. Two pieces of the oldest aragonite generation truncated by bevels from the junction of the Ovlna Passage and Hlbok Hall (samples JOA 1 and JOA 2) were corroded and partially recrystallised to calcite. Measurements were done with alpha spectrometry (OCTET PC, EG $ G ORTEC; by Hercman in 1999). Both analyses yielded low uranium contents (0.6 and 3 ppm) and a 230Th/234U ratio significantly higher than 1 (about 1.4). The high ratio, unusual in the nature, suggests that there was preferential leaching of uranium from the samples during recrystallisation and/or corrosion, rendering computed age unreliable. Therefore, TIMS U series analyses were made by Ford in 2001 on similar eroded old aragonite and calcite flowstone of the oldest generation that were found as fragments within the cave. The mineralogy was confirmed by X-ray diffraction. The extractions were carried out in a clean room with laminar flow hoods, and two analyses were made of each sample. Results are summarised in Table 2. TABLE 2 TIMS U series analyses Sample No. U content [ppm] 234U/238Th 230Th/234U 230Th/232Th Age [y] 01 Och 1A aragonite 5.5 0.9934 1.046 # 0.003 63,000 Age cannot be calculated 01 Och 1 B aragonite 5.6 0.9975 1.057 # 0.003 49,000 Age cannot be calculated 01 Och 2 A calcite 0.76 1.0268 0.9938 7,065 449,000 +69,000 -42,000 01 Och 2 B calcite 0.67 0.9999 1.0101 # 0.007 7,064 Age cannot be calculated N.B. Error margins quoted are two standard deviations

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.11 Two analyses of aragonite show that the sample is very clean and has the high U content typical of aragonite. The 230Th/234U ratio is just in secular equilibrium, so that an age cannot be obtained by this method. 234U/238U on the other hand is not in equilibrium, implying that the sample is certainly younger than 1,250,000 years. It is probably a little older than the calcite sample. Two analyses of calcite are very similar. The U content is satisfactory for calcite (most speleothems have between 0.05 and 1.0 ppm). The sample is clean (230Th/232Th >>20). In the second analysis the 230Th/234U ratio just exceeded 1.0000, possibly from detritus. But it is very similar to the first analysis in all other respects. This can be taken as confirmation that the age estimate of approximately 450 ka is acceptable. Palaeomagnetism A clastic sediment section well exposed on the northern side of the Ovlna Passage is about 0.7 m high (Fig. 1). From the top, it is composed of the following layers: 1 – white flowstone with stalagmite; 2 – clay, reddish brown, with greyish black schlieren enriched in Mn-compounds, massive, laminated in places, disintegrated into irregular polyhedral fragmen ts (samples OCH 1 and OCH 2); 3 – alternation of reddish brown clay (thickness max. 1.5 cm; samples OCH 3 and OCH 4) with layers blackened by Mn-rich minerals (thickness from 4 to 6 cm), and white bands with allophane crystals (thickness of 1 to 5 cm; Fig. 2/4); 4 clay, reddish brown, massive, disintegrates into irregular polyhedral fragments (sample OCH 5). The flowstone, about 1-2 cm thick (layer No. 1 on Fig.1) was dated by the U series alpha counting method (Hercman). The U content (about 6 ppm) was similar to aragonite samples (see above). The content of detrital Th was negligible. The analysis gave an age of 164,0007,500 at one standard deviation. Methods Samples were demagnetised by the alternating field procedures up to 1,000 Oe with a Schonstedt GSD-1 machine. The remanent magnetisation Jn was measured on a JR-5 spinner magnetometer (Jelnek, 1966). Values of volume magnetic susceptibility kn were measured on a kappa-bridge KLY-2 (Jelnek, 1973). Separation of the respective remanent magnetisation components was carried out by multi-component analysis (Kirschvink, 1980). Results The magnetic properties, both Jn and kn values, of samples from layer No. 2 are distinctly different from those of layers Nos. 3 to 5 (Tab. 2). Sample OCH 1 (layer No. 2) displays normal remanent magnetisation. All underlying samples are magnetically reversed (Fig. 1). This polarity change can be correlated with the Brunhes/Matuyama reversal of 780 ka B.P. (Pruner et al. 2000), because the speleothem date of 164 ka establishes that it must be older than any of the short-lived reverse magnetic excursions within Matuyama chron ( cf. Zhu and Tschu, Eds. 2001). TABLE 3 Principal magnetic properties of samples Sample No. Jn [pT] Kn [10-6SI] Polarity OCH 1 30,375 714 Normal OCH 2 35,651 1,999 Reverse OCH 3 11,879 347 Reverse OCH 4 283 176 Reverse OCH 5 1,715 161 Reverse Mean value 15,981 679 Standard deviation 16,285 771

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.12 Discussion At several places in the cave it can be clearly seen that the voids were filled by ochres which were later removed. Floors of passages that are partly filled at present have an oval shape. They were formed either before the ochres were produced by ankerite oxidation or before the eroded ochre residuum was deposited in water-filled passages. The first subsurface cavities were formed by corrosion of the limestone and oxidation of the ankerite. These cavities were flooded by meteoric water infiltrating along the major fault line in Vstupn and Mramorov Ha lls and also along the lesser fissures in Hlbok Hall and Sie mlie nej cesty Hall. Continuous and dominantly horizontal cavities formed along parallel fissures. Irregular corrosion features developed on the bedrock surfaces, these were cons iderably enlarged and modified by later corrosional events. The only original forms still preser ved are irregular niches and cupolas found above the younger corrosion bevels. The source of the carbon dioxide for intensive corrosion can be found in the ankerite weathering with an end product of goethite, i.e. by a process similar to that described by Kempe (1998) from Harz (Germany), according to following equations: 2FeCO3 + 2CO2 + 2H2O = 2Fe(HCO3)2` [1] then 2Fe(HCO3)2 + 1/2O2 + H2O = 2Fe(OH)3 + 4CO2 [2] or 4FeCO3 + O2 + H2O = 2Fe2O3.nH2O + 4CO2 [3] where 2Fe2O3.nH2O is limonite. The frequent presence of pyrite inclusions in the limestones and of allophane, a typical product of acid decomposition of clay minerals, suggests that the corrosion might have been enhanced sulphide weathering and oxidation to H2SO4. However, the absence of any gypsum replacing limestone or of native sulphur in the cave indicates this effect was probably minor. The origin of the niches and cupolas was linked by some authors (Nicod, 1974; Choppy, 1994) with hydrothermal processes. However, both forms can originate from convection induced by gravitational settling of water enriched with solute ions, without any hydrothermal influence ( cf. Curl, 1966; Cordingley, 1991; Klimchouk, 1997a). The density of water in the phreatic zone of a karst system will increase as it dissolves the enclosing rock. During continuous or periodic infiltration of “fresh” water into a water-saturated environment, a density gradient forms. This can generate convection cells in the water body, which may produce corrosion forms in the enclosing limestone. The effect is essentially limited to conditions of static or very slow water movement. In rapidly moving water this effect will be negligible. Convection is a differential process, the dissolution producing roof cupolaand tube-like depressions in roofs that are below the waterline, horizontal ceilings (bevels) or corrosion notches in the walls at the waterline, and inclined planar walls beneath it. Convection circulation and solution cannot only modify morphological forms but it can also influence the entir e pattern of such cave systems during their initial phases where the waters are predominantly static or semi-static. This feature is particularly common where there is artesian speleogenesis (Klimchouk, 1997b). Ochtin Aragonite Cave is developed in an isolated lens of limestone surrounded by insoluble rocks. Such lenses readily fill with water during floods and drain only slowly afterwards. Corrosion notches along the walls are produced in very acidic stagnant water conditions and where roofs dip down they will eventually be planed off as bevels at the waterline. Stagnant water still forms a lake in the deepest part of Hlbok Hall. The difference of elevation between the highest bevel in the cave and the present lake level is 12 m. Nearly horizontal bedding favoured bevelling in Jeovit Passage and Aragonitov zhrada, while corrosional ledges that have developed in steeply dipping beds indicate the corrosional origin of the be vels. Notches and bevels commonly intersect the older speleothems, such as those discussed above. Planes of repose (Lange, 1963; Goodman, 1964) are found in many parts of the cave (Fig. 2/3). These are the inclined bedrock surfaces developed in the lower portions of cavern walls. They too formed during periods of very slow water circulation when accumulated insolubles blocked solution enlargement at the base of a wall in a flooded section of cave. In passages where bevels are developed, planes of repose are similar to Facetten ( sensu Kempe et al. (1975).

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.13 Iron ochres were formed by the weathering of ankerite/siderite metasomatites. The ochres are composed principally of goethite and a variety of autochthonous minerals deposited in flooded cave conditions or by the dripping water. The ochres contain on average 50 % water. They cover a substantial area of the cave and act as an important humidity exchanger stabilising the microclimate. Black Mn-bearing loams contain Ni-bearing asbolane and birnessite, which developed from asbolane. Mn-oxides were most probably formed by microbial precipitation at the bottom of water bodies, as recently described by Andrejchuk and Klimchouk (2001), i.e. just after the cave was drained and fresh air entered it. Beds of allophane occur within the asbolane layers (Fig. 2/4). Because it forms in a low pH environment, allophane is not a common mineral in karst caves. We suggest that the allophane could have formed in the acidic conditions produced by weathering of sulphide minerals. A kaolinitic mineral similar to halloysite was formed by maturation of allophane. The asbolane layer formed an important geochemical barrier, which caused the concentration of the REE, the growth of La-Ndbearing phosphate and eventually the formation of the REE-enriched Mn oxide. The allogenic minerals, which have entered the cave are extremely fine-grained and partly weathered and abraded. They indicate that the cave was poorly connected with the surface, allowing only slow infiltration through narrow or choked fissures, rather than direct, open communication with flowing streams. There are three different generations of aragonite. The oldest speleothems are preserved as corroded relics truncated by bevels (Fig. 2/2). The aragonite in them is partly recrystallised to calcite. There were at least two separate periods of the growth. TIMS U series ages for recrystallised calcite of the older sample indicate an age of about 450 ka. TIMS U series ages for the aragonite cannot be calculated as they are at the limit of the method. 234U/238U on the other hand is not in equilibrium, implying that the sample is certainly younger than 1.25 Ma. It is probably a little older than the calcite sample and related to a warm episode of Elster 1/2. The aragonite in the younger recrystallised speleothems yielded U series dates indicating an Eemian age (138-121 ka). The pre-recrystallisation age may be greater. The second aragonite generation, of spiral and acicular aggregates (Fig. 2/5), began to be depos ited during Late Glacial (Allerod, 14 ka). Growth has continued to the present day. The youngest generation, of fine acicular aggregates of aragonite and miniature helictites, is also actively growing (Fig. 2/6). Both of the younger generations have similar isotopic compositions, indicating that they originated in conditions very similar, or identical, to those found at present in the cave. TABLE 4 Succession of processes during the origin of Ochtin Aragonite Cave Age (ka) Process Water regime Notes Upper Cretaceous Hydrothermal activity Thermal Initial speleogenesis Phreatic Late Tertiary Pleistocene Cave enlargement Epiphreatic, vadose >780 <780 Corrosion, bevels and deposition of sediments (Ovlna Passage) Highstand Lowstand Asbolane as a product of sulphide weathering by oxidising waters Glacial? Erosion/redepositon of cave fill, possible bevels Fluctuations Periods of water highstands not excluded, CO2 released from ankerite decomposition in oxidising waters 500 450 Speleothems, the oldest generation I Lowstand Calcite recrystallised from aragonite prevails and is somewhat younger than aragonite (ca 50 ka) Glacial? Corrosion, bevels, cut of rocky pendants, erosion/redeposition of cave fill Highstand and fluctuations CO2 released from ankerite decomposition in oxidising waters

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.14 Conclusions The modern morphology of the cave reflects a comparatively complex evolution under particular lithological and hydrogeological conditions within an isolated lens of karst rock surrounded by insoluble rocks. Such lenses can become filled with water, often with artesian confinement. Primary phreatic subsurface cavities were formed by the corrosion of the limestone and oxidation/erosion of the ankerite. Elongated, chiefly horizontal cavities formed along parallel fissures. Irregular corrosion forms developed on the bedrock surfaces. The niches and cupolas are relics of phreatic speleogens created by convection induced by gravitational, density-derived circulation of water in a regime of very slow flow. Hydrothermal effects are not necessary. The abundant pyrite together with a common allophane indicates the carbonic acid corrosion was most probably enhanced by sulphide weathering producing diluted brines. Thermohaline convection cells operating in the flooded cave might also have influenced the wall morphology. Younger corrosional events under vadose conditions changed the original forms to a large extent. The intensity of corrosion was enhanced by carbon dioxide from ankerite weathering in the oxidising meteoric waters. The water-level fluctuations were repeated several times as indicated by several levels of flat roofs (bevels), wall niches and planes of repose. Bevels form by corrosion in stagnant water conditions. Roof planation was influenced both by limestone bedding and by the duration and intensity of water convection. Bevels intersected older speleothems. Corrosion notches along the walls indicate that the levels of stagnant water were stable for long periods, representing significant phases of cave enlargement. Planes of repose also indicate slow water circulation following floods; accumulated insolubles blocked solution enlargement at the base of a cave wall. Water-level oscillations and water flow have to be very slow, as indicated by the fact that the sediment section studied in Ovlna Passage survived several submerge nces. Nevertheless, the velocity of flow during the early phases of the cave evolution had to be sufficient to transport the clastic products of the ankerite disintegration into lower levels of the cave. Dating this sequence of processes is a complicated and risky task. We can assert that the cave started to form before 0.78 Ma according to the palaeomagnetic data from the oldest dated cave fill in Ovlna Passage. The roof of that passage is the highest preserved bevel to have developed under the succeeding vadose conditions. We may tentatively link the formation of this highest bevel with the oldest sedimentary fill. Therefore, vadose conditions probably were established 0.78 Ma. The phreatic phase of cave development has to be older (Late Tertiary/Pleistocene), but it cannot be dated properly. It appears that the age of cave origin is close to that suggested by Kubny (1959) and Homza, Rajman and Roda (1970). The sequence of cave development summarised in Table 4 is based primarily on the U series dating of flowstones and aragonite/calcite speleothems. Acknowledgements The authors wish to express their thanks especially to Dipl.-Ing. Jozef Hlav Director of the Slovak Caves Administration in Liptovsk Mikulš and Mr. Jn Ujhzy, Head of the Ochtin Aragonite Cave for permission to conduct research and to take samples during a period of 1996 to 2001. We acknowledge the contribution of: Dr. Karel Melka and Mr. Ji # Dobrovoln (X-ray analyses and interpretations), Dipl-Ing. Anna Langrov (microprobe analyses; all from Laboratory of Physical Methods, Institute of Geology, AS CR Prague), Dr. Daniela Venhodov (production and evaluation of palaeomagnetic data; Department of Palaeomagnetism, Insitute of Geology AS CR Prague), Dr. Karel Žk (isotopic analyses; Czech Geological Institute, Branch Barrandov, Prague), and Dr. Jana Ederov (DTAGTA analyses, University of Chemical Technology, Prague). The research was carried out under the Agreement on Scientific Co-operation between the Slovak Caves Administration and the Institute of Geology AS CR. Costs were covered from sources of the Caves Administration (Task B. of the Main Activity Plan in 1998), Plan of Scientific Activity No. Z 03-013-912 of the Institute of Geology AS CR, and Grant No. A3013201 of the Grant Agency of AS CS. U series analyses by Ford at McMaster University were supported by a grant in aid of research from the National Scientific and Engineering Research Council of Canada.

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.15 References Andrejchuk V.N. and Klimchouk A.B. 2001. Geomicrobiology and Redox Geochemistry of the Karstified Miocene Gypsum Aquifer, Western Ukraine: The Study from Zolushka Cave. Geomicrobiology Journal 18 : 275-295. Andrusov D. 1958. The geology of Czechoslovak Carpathians, Part I. Publ. House Slovak Acad. Sci. 304 pp. Bratislava (in Slovak). Bajank Š. and Vozrov A. (Eds.). 1983. Explanations to the geological map of the Slovensk rudohorie Mts. – eastern part (in Slovak). Geol. Inst. D. Štr 223 pp. Bratislava (in Slovak). Bella P. 1995. Principles and theoreticaly-metodical aspects of classification of cave morphological types. Slovensk kras 33 : 3-15. Liptovsk Mikulš (in Slovak, Engl. summ). Bella P. 1997. The views to the genesis of the Ochtin Aragonite Cave (in Slovak). Aragonit 2 : 13-14. Liptovsk Mikulš (in Slovak). Bella P. 1998. Morphological and genetic features of the Ochtin Aragonite Cave. Aragonit 3 : 3-7. Liptovsk Mikulš (in Slovak, Engl. summ). Choppy J. 1994. La premire karstification. Synthse splologique et karstique. Les facteurs gographiques 3. Splo-Club de Paris, Club Alpin Franais, 1-70. Paris. Clek V. and Bedn # ov J. 1993. Silcretes in the Bohemian Karst. esk kras (Beroun) 18 : 4-13 (in Czech, Engl. summ.). Clek V., Bosk P., Melka K., Žk K., Langrov A. and Osborne A. 1998. Mineralogical investigations in the Ochtin Aragonite Cave. Aragonit 3 : 7-12. Liptovsk Mikulš (in Czech, Engl. summ.). Clek V., Langrov A. and Melka K. 1998. Pseudokarst phosphate-all ophane speleothems from ice caves of the Podyj National Park. Speleo (Praha) 26 : 20-27 (in Czech, Engl. summ.). Clek V. and Šmejkal V. 1986. The origin of aragonite in caves. The study of stable isotopes. eskoslovensk kras 37 : 7-13. Praha (in Czech, Engl. summ.). Cordingley N. J. 1991. Water stratification in active phreatic passages. Cave Science 18(3): 159. London. Curl R. L. 1966. Cave conduit enlargement by natural convention. Cave Notes 8, 1 : 4-8. Castro Valley. Droppa A. 1957. Ochtin Aragonite Cave. Geografick asopis 9, 3 : 169-184. Bratislava (in Slovak, Russ. and Germ. summ.). Ford D. C. and Williams P. W. 1989. Karst Geomorphology and Hydrology Unwin Hyman, 601 pp. London-Boston-Sydney-Wellington. Gal $ 1996. Exploration and protection of aragonite caves in surroundings of the Hrdok Hill. In Sprstupnen jaskyne – vskum, ochrana a vyuvanie jask # zbornk refertov P. Bella (Ed.): 130-133. Liptovsk Mikulš (in Slovak, Engl. summ.). Gal $ 1998. Karst phenomena in the surroundings of the Hrdok Hill and their relation to the Ochtin Aragonite Cave. In Vskum, vyuvanie a ochrana jask # zbornk refertov P. Bella (Ed.): 44-52. Liptovsk Mikulš (in Slovak, Engl. summ.). Gal $ and Ženiš P. 1986. Karst of the Revcka Highland. Slovensk kras 24 : 27-60. Liptovsk Mikulš (in Slovak, Russ. summ.). Goodman L. R. 1964. Planes of repose in Hllern, Germany. Cave Notes 6, 3 : 17-19. Castro Valley. Ivani ka J., Snopko L, Smopkov P. and Vozrov A. 1989. Gelnica Group – Lower Unit of Spiško-gemersk rudohorie Mts. (Early Paleozoic, West Carpathians). Geol. Zbor. Geol. Carpath. 40, 4 : 483-501. Bratislava. Hill C. and Forti P. 1997. Cave Minerals of the World Second edition. Natl. Speleol. Soc., 463 pp. Huntsville. Homza Š., Rajman L. and Roda Š. 1970. Origin and evolution of karst phenomenon of the Ochtin Aragonite Cave. Slovensk kras 8 : 21-68. Liptovsk Mikulš (in Slovak, Engl. abs., Germ. summ.). Ivanovich M. and Harmon R.S. 1992. Uranium Series Disequilibrium Applications to Environmental Problems. 2nd Ed. Clarendon, 910 pp. Oxford Jelnek V. 1966. A high sensitivity spinner magnetometer. Studia geophysica et geodetica 10 : 58-78. Praha. Jelnek V. 1973. Precision A.C. bridge set for measuring magnetic susceptibility and its anisotropy. Studia geophysica et geodetica 17 : 36-48. Praha. Jennings J.N. 1985. Karst Geomorphology Oxford, Basil Blackwell, 293 pp. Kempe S. 1998. Siderite weathering, a rare source of CO2 for cave genesis: th e Eisenstein Stollen

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P.Bosk et al. / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.16 System and adjacent caves in the Iberg, Harz Mountains, Germany (abs.). Journal Cave Karst Studies 60, 3 : 188. Huntsville. Kempe S., Brandt A., Seeger M. and Vladi F. 1975. „Facetten“ and „Laugdecken“, the typical morphological elements of caves developed in standing water. Annales de Spleologie 30, 4 : 705-708. Moulis. Kirschvink J. L. 1980. The least-squares line and plane and the analysis of palaeomagnetic data. Geophysical Journal of the Royal Astronomical Society 62 : 699-718. Oxford. Klimchouk A. 1997a. Artesian speleogenetic setting. Proc. 12th Int. Congr. Speleology 1: 157160. La Chaux-de-Fonds. Klimchouk A. 1997b. Speleogenetic effects of water density differences. Proc. 12th Int. Congr. Speleology 1: 161-164. La Chaux-de-Fonds. Kubny D. 1959: Aragonite cave in Slovakia. Ochrana prrody 14, 1 : 17-18. Bratislava (in Slovak). Lange A. 1962. Water level planes in caves. Cave Notes 4, 2 : 12-16. Castro Valley. Lange A. 1963. Planes of repose in caves. Cave Notes 5, 6 : 41-48. Castro Valley Li W.-X., Lundberg J., Dickin A.P., Ford D.C., Schwarcz H.P., McNutt R. and Williams D. 1989. High-precision mass-spectrometric uranium-series dating of cave deposits and implications for paleoclimatic studies. Nature 339 : 534-536. Mišk M. 1953. Geology of the area between Jelšava and Šttnik. Geologick sbornk Slov. Akad. Vied 4, 3-4 : 557-587. Bratislava (in Slovak). Nicod J. 1974. Les rgions karstiques de Slovaquie et de Hongrie septentrionale. Bull. Soc. Gogr., N.S. 82, 12, 14 : 11-25. Marseille. Pruner P., Bosk P., Kadlec J., Venhodov, D. and Bella, P. 2000. Paleomagnetic research of sedimentary fill of selected Slovak caves. In Vskum vyuvanie a ochrana jask # zbornk refertov z 2. vedeckej konferencie P. Bella (Ed.): 13-25. Liptovsk Mikulš (in Czech, Engl. Summ.). Rajman L., Roda Š. jr., Roda Š. sen. and Š uka J. 1990. Physico-chemical investigation of the karst phenomenon of the Ochtin Aragonite Cave. Final Report. MS, Slov. Muz. Ochr. Prr. a Jaskiniarstva archive : 1-46. Liptovsk Mikulš (in Slovak) Rajman L., Roda Š. jr., Roda Š. sen. and Š uka J. 1993. Untersuchungen ber die Genese der Aragonithhle von Ochtin (Slowakei). Die Hhle 44, 1 : 1-8. Wien. Slabe T. 1995. Cave Rocky Relief and its Speleogenetical Significance. Zbirka ZRC 10 128 pp. Ljubljana. Šev k R. and Kantor J. 1956. Aragonite cave in the Hrdok Hill near Jelšava. Geologick prce, Zprvy 7 : 161-170. Bratislava (in Slovak). Tsler R. (Ed.) 1991. Owen 90-New Zealand. Czech Speleol. Soc. 60 pp. Trutnov. Tsler R. and Clek V. 1999. Decoration of giant domes in the Bohemia Ca ve, New Zealand: the World biggest aragonite cave? Speleofrum 99 18 : 35-40. Praha (in Czech, Engl. abs.). Zhu R.X., Tschu K.K. (Eds.) 2001. Studies in Paleomagnetism and Reversals of Geomagnetic Field in China Beijing, Sciences Press, 168 pp.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info Prediction of condensation in caves C. R. de Freitas (1) and A. Schmekal (2) (1) School of Geography and Environmental Science, University of Auckland, New Zealand Email: c.defreitas@auckland.ac.nz (2) School of Geography and Environmental Science, University of Auckland, New Zealand Abstract Condensation is an important process in kars t environments, especially in caves where carbon dioxide enriched air can lead to h igh rates of condensation corrosion. The problem is there has been ve ry little research reported in the literature dealing with con densation as a microclimate process. This study addresses the problem and reports on a meth od for measuring and predicting condensation rates in a limestone cave. Electr onic sensors for measuring condensation and evapor ation of the condensate as part of a single continuous process of water vapour flux are tested and used to co llect 12 months of data. The stud y site is the Glowworm touris t cave in New Zealand. Condensation is a function of the vapour gradient between rock surfaces in the cave and cave air. The size of t he gradient is largely determined by air exchange with the outside. The results s how that the numerical model to predict condensat ion works well. Given that rock-surface temperature in the cave doe s not vary much, condensation is essentially a function of cave air temperature and the processes that affect it, mainly, air exchan ge with outside. The results show that condensation can be cont rolled by controlling ventilation of the cave. Keywords: Condensation, Cave microclimat e, Evaporation, Tourist cave management. Introduction Condensation is an important process in karst environments, especially in caves. For example, Hill and Forti (1986) cite seven different types of speleothems formed from condensation coupled with the evaporation of condensate. The condensation process plays a variety of roles, but two of these are particularly important. The first occurs where water condensing onto cave walls that are made of a soluble rock mineral (calcite, dolomite, gypsum, halite, carnallite etc.) is undersaturated with respect to the mineral, the potential exists for dissolution to occur. The process is called “condensation corrosion” (Ford and Williams, 1989, p. 309). It may create surface impressions on attractive speleogen features. Water from condensation can cause this because its chemistry makes it aggressive. Carbon dioxide, water and calcium carbonate (limestone or calcite) react to give soluble calcium and hydrogencarbonate ions in water. Condensation water becomes considerably more corrosive if it contains substantial amounts of dissolved carbon dioxide. In tourist or show caves, for example, visitors breathe out warm air saturated with water vapour together with over 4% by volume of carbon dioxide at a temperature usually much higher than the cave air. This air containing high concentrations of carbon dioxide will condense as it comes into contact with the colder surfaces of the cave. The second process occurs during times when condensation water evaporates and carbon dioxide is removed from saturated solutions of calcium and hydrogencarbonate ions causes precipitation of calcite. This process produces soft unattractive microcrystalline, flaky deposits of calcite. This cycle of condensation and evaporation of condensate is believed to enhance condensation corrosion (Tarhule-Lips and Ford, 1998). Cave resources are essentially non-renewable and human impacts are cumulative and often irreversible (Gillieson, 1996). Increasing cave tourism worldwide presents problems because of this irreversible degradation. Previous work on caves, especially tourist caves, has shown that an understanding of cave micr oclimate processes is crucial to understanding, managing and protecting the cave ecosystem (de Freitas, 1998; de Freitas and Banbury, 1999), but gaps in understanding certain key processes remain, in particular, those governing

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 condensation. The problem is that there has been very little research reported in the literature dealing with condensation as a microclimate process in caves. Papers by Dublyansky and Dublyansky (1998, 2000) that review the topic confirm this. Explanatory models of causal process are speculative and remain untested. Recently, however, de Freitas and Schmekal (2003) devised a reliable method for measuring condensation and evaporation as part of a single continuous process of water vapour flux. The aim here is to report further on this research, specifically on the method for predicting condensation rates on cave rock surfaces. Background The study site is the Glowworm Cave, New Zealand, widely regarded as an attraction of considerable aesthetic and ecological significance. It has one of the highest visitor usage rates of any conservation land in New Zealand. Four times the number of people visit the Glowworm Cave than the next most popular cave in either New Zealand or Australia. For this reason it is considered to be a valuable national resour ce and one that requires careful management if its attractiveness is to be protected and the resource sustained. The Glowworm Cave is located in the North Island of New Zealand at latitude 38o15Â’S, longitude 175o06Â’E. The region has a subtemperate climate with an average annual rainfall of 1530 mm. Average daily maximum and minimum air temperatures in the warmest month, January, are 24.1 and 12.6 oC, respectively. Average maximum and minimum temperatures in the coolest month, July, are 13.1 and 3.3 oC, respectively. The water vapour content of the air is relatively high throughout the year in the region, with a mean vapour pressure of 13 hPa. The cave is situated in a ridge of Oligocene limestone. The area above the cave is a scenic reserve of native vegetation administrated by the New Zealand government agency called the Department of Conservation. The Glowworm Cave is made up of 1,300 m of interconnected passageways with an estimated volume of approximately 4000 m3. It consists of three levels an upper, middle and lower level (Fig. 1). The cave has two entrances, an upper entrance and a lower entrance, 14 m vertically apart. The upper entrance is equipped with a solid door that, when closed, seals the opening preventing airflow. The upper level of the cave consists of the Blanket Chamber and the Blanket Chamber passage. The Blanket Chamber is 40 m long and ranges in diameter from 1 to 4.5 m2. The Main Passage is a 39 m long section with an elliptical cross-section varying between 3 m2 and 7 m2. This passage leads past the Tomo, which connects the lower level (Grotto) and the upper level to the Catacombs, a much larger Chamber (Fig. 1). Fig.1. Schematic isometric plan of the Glowworm Cave showing measurement sites and named cave features (from Barthow, 1988).

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 Another part of the upper level of the cave is the Organ Loft and the Organ Loft Side Passage. The Organ Loft is a cul-de-s ac passage. The Banquet Chamber and Cathedral form the intermediate level. The Cathedral is 40 m long, 11 m at its widest and up to 13 m high. It is the largest chamber in the cave and links all chambers The third level is the Grotto, which is part of the stream passage of the Waitomo River. The Grotto is a large chamber approximately 30 m long and 10 m wide. The Grotto has the main displays of the glowworm ( Arachnocampa luminsosa ) in the cave. From here the stream flows down through a passage and sump and then past the Demonstration Chamber. After this the stream flows for approximately another 180 m before leaving the cave (Fig.1). Airflow in the cave due to both convectional or gravitational forces (de Freitas et al ., 1982) and this airflow is a key component of a cave’s climate (de Freitas and Littlejohn, 1987). The speed and direction of flow is determined by the difference of mean density of the outside and inside air (de Freitas et al, 1982). Since air density is mainly a function of air temperature, the latter can be used as the main indicator of airflow (de Freitas et al ., 1982). When the outside air is cooler and thus denser than the cave air, the warmer cave air rises and flows towards and then through the Upper Entrance and replaced by cold air at the Lower Entrance. This process, dr iven by convection, is called “winter” flow. In contrast, “summer” flow occurs when cave air is co oler and denser than the air outside the cave. The flow of air is driven by gravity through the cave and out the Lower Entrance (de Freitas et al. 1982). In transitional times where the temperature gradient inside and outside the cave is small, there is little or no airflow. Method Modelling condensation Condensation per se is a dynamic process of moisture flux that involves both condensation and evaporation of the condensate, depending on the direction of the vapour gradient between the air and moist surface. When the amount of condensation over a given period exceeds the evaporation of condensate over that same period, condensation is observed to have occurred. Condensation water will accumulate if this condition re-occurs, otherwise it will dissipate. In the case of caves, the flow is to or from the surface of cave walls and features in the cave. The assumption is that at the surface there is a boundary layer of air that is saturated and has the same temperature as the surface. This boundary layer interacts with the surrounding air causing condensation or evaporation of condensate in a dynamic relationship that is driven in large part by the vapour gradient. The moisture flux across this gradient strictly speaking the resistance to the diffusion of vapour across the boundary layer is controlled by the rate of air movement and the roughness of the surface (Monteith, 1957), collectively referred to here as the combined convection moisture transfer coefficient. Condensation occurs when the dewpoint temperature of the cave air is higher than the temperature of the rock surface. However, to quantify the movement of a mass of water vapour, specific humidity rather than dewpoint temperature must be used. The rate of condensation (C) is given as: C = (qr – qa) kv (1) where C is rate of condensation (g m-2 s-1), qa is specific humidity of the air (g kg-1), qr is saturation specific humidity at surface temperature (g kg-1), kv is the combined convective water vapour transfer coefficient. Specific humidity terms qa and qr are a function of vapour pressure and can be calculated from Neiburger et al (1982): qr = 0.622 esr (2) qa = 0.622 e (3) where esr is saturation vapour (hPa) pressure at rock-surface temperature and e is vapour pressure of air (hPa). Vapour pressure and saturation vapour pressure terms can be found using any of number of formulae as, for example, from Grace (1983): e = es – 0.666(Tdb – Twb) P (4) where es is saturation vapour pressure of the air (hPa), Tdb is dry bulb temperature (oC), Twb is wet bulb temperature (oC) and P is atmospheric pressure (hPa). Saturation va pour pressure is: es = exp[a+(b T c)/( T d)] (5) where a is 1.80956664, b is 17.2693882, c is 4717.306081, d is 35.86 and T is air or surface temperature (K). However, where vapour gradients are very small, as is frequently the case in cave environments, more precise formulae are required for the calculation of vapour pressure and saturation vapour pressure, such as provided by Jensen (1983).

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 The combined convective water vapour transfer coefficient, kv, is a function of air movement and surface roughness (Pedro and Gillespie, 1982; McAdams, 1954). Compared to the boundary layer outdoors, surface roughness is relatively constant in most caves. In an open environment where wind speed varies greatly and can reach much higher levels than in caves, wind is an important variable. In caves, however, airflow is limited and rate of flow in the case of the Glowworm Cave is extremely low. As a consequence, variability is small. With this in mind, de Freitas and Schmekal (2003) show empirically that kv = 3.7 kg m-2 s-1 in equation (1) fits well with observations of C regardless of the location within the cave. Measuring condensation De Freitas and Schmekal (2003) devised a novel method for measuring the vapour flux to and from a surface using what they called “condensation sensors”. They are simple to construct and their size can be customised so it is possible to install them on uneven surfaces such as a cave wall. The condensation sensors consist of an electrical grid of two sets of parallel wires mounted on a circuit board. When condensation occurs or evaporation of the condensate takes place on the sensor’s surface, the resistance between the wires changes. To provide greater sensitivity, the wiring consisted of multiple fingers of interleaved conductive tracks made of copper (Fig. 2). Sensitivity can be altered by varying the number of conductors. Fig. 2. Condensation sensor. To obtain rates of condensation, conduction readings have to be conve rted to equivalent vapour fluxes. To do this the sen sor are weighed when dry and the conductivity reading set at zero. Using an atomiser, very fine drops of water were sprayed onto the sensor in stages and weighed at each step (de Freitas and Schmekal, 2003). The sensors showed no influence of ambient temperature over the range tested (10.0 C – 20.0 C). Data collection Data were assembled using a fully automated system of sensors and recorders and supplemented by direct measurement using hand-held instruments. Automated measurements were made of wet (Twb) and dry bulb (Tdb) air temperature, rock temperature (Tr), and airflow rate and direction. Wet and dry bulb temperatures (Campbell 107B thermistors) were measured at the Tomo, Banquet Chamber and at the Jetty. Another dry bulb thermister and humidity sensor (Vaisala Hummitter 50Y) was installed outside the cave. Readings were recorded every thirty minutes by a data logger (Campbell CR10) Rate of airflow and direction into and out of the cave are measured using a sensitive Pulse Output Anemometer (A101M) and an airflow direction sensor (Potentiometer Windvane W200P). The airflow instruments were located in the entrance area, just inside the cave door. An electronic sensor records periods when the entrance-door is open and airflow readings are taken every three seconds. The data logger then records the maximum wind speed for each one-minute interval, and these are then averaged for the length of the time the entrancedoor is open. Rock temperature was measured using a thermister (Campbell 107B). Internal rock temperatures give an indication of trends in the longer-term thermal state of the cave, as well as the direction of heat flow to and from the rock-surface (de Freitas, 1998). Rock temperatures are measured at the Tomo recorded every six hours. To sample more extensively through the cave, direct measurements were made using hand held instruments. Wet bulb temperature and dry bulb temperature were measured using a full-sized Assmann Psychrometer (Casella, Type 8900/1). The instrument can be read with accuracy to a resolution of 0.1 oC. From these data, saturation vapour pressure, humidity and dew-point temperature were determin ed using the procedure described earlier. For detailed measurements of airflow in various parts of the cave, a Dwyer hotwire anemometer (Series 470), accurate to 0.05 m s-1, was used. Rock-surface temperatures were

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 measured using a portable electronic instrument (Ultrakust, Type 4444-1B) and probe especially designed for measuring surface temperature of flat, solid objects. The flat temperature-sensing element of the probe is covered with an insulating epoxy and fibreglass resin attached to Teflon insulated leads to protect it from the thermal influences of air when it is pressed against the surface to be measured. The sensor is a small thermister pearl of high thermal-conductivity material (silver and gold) so that short response times and small heat capacity are achieved. Accuracy of the instrument is better than 0.1 oC with a full-scale response time of four seconds. Two readings were taken with the Ultrakust instrument at the condensation measurement sites described below. One reading was of the surrounding cave wall and the second reading of the “dummy” metal plate used to check that sensor surface temperatures were the same as rock-surface temperatures. To ensure that the assembled data was characteristic of the cave as a whole, four measurement sites were selected that represented different parts of the cave, namely the Organ Loft (deep cave), Cathedral (cavernous interior), the Banquet Chamber (transiti onal zone) and Blanket Chamber (near entrance zone). The locations of these sites are shown in Fig. 1. The Organ Loft is a cul-de-sac passage. Here there is little air exchange with the outside and conditions are stable. The Cathedral site is also within the deep cave zone, but in this case along the main airflow route. This area is also the biggest chamber in the cave. The Banquet Chamber is within the transitional zone and like the Cathedral site is in the main airflow route. The fourth site was in the Blanket Chamber, which represents an area where the cave air can readily interact with the outside air. Four condensation sensors were installed at each measurement site on a vertical portion of the cave wall 900 mm above the floor and attached to four dedicated Campbell Scien tific CR 10 data loggers. Readings were taken every five seconds and recorded as 10-minute averages. Measurements were taken over a 13-mont h period from December 1999 to December 2000. Results To assess the performance of the model, the difference between observed rates of condensation (Co) and calculated values (Cc) were tested using Pearson’s product moment correlation (r2). The sample size is 750. The mean difference between Cc and Co is 0.062 g m-2 h-1and the standard deviation 0.165 g m-2 h-1. The correlation coefficient is 0.97 (Fig. 3). Overall the results show that the model performs well. Fig. 3. A comparison of calculated and observed condensation rates, Cc and Co respectively (g m-2 h-1). The standard deviation is 0.165 g m-2 h-1, the correlation coefficient r2 is 0.97 and the sample size is 750. Note that over plotting of data points occurs frequently because condensatio n rates change only very gradually over time.

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 Next, the data were examined to assess the influence of outside conditions on condensation rates. Since rock-surface temperature is relatively stable, cave air temperature is the main factor that influences condensation rates. Cave air temperature is jointly determined by outside air temperature and cave ventilation rate, which is itself a function of outside air temperature. While temperature fluctuations outside the cave are much larger (0.2oC to 28.1oC) than inside the cave (12.4 oC to 18.9 oC), they both tend to follow th e same pattern. Figures 4 and 6 are typical examples of this. It follows that, as outside air temperature influences the cave climate, the different outdoor thermal conditions play a vital role in conden sation rates. In general, when it is warm outdoors qa exceeds qr during the daytime and condensation occurs. When temperatures are low outside, qa is in general lower than qr and no condensation occurs. The closing and opening of the entrance-door can be used to control airflow through the cave and consequently cave air temperatures. Fig. 4. Air temperature in the cave at the Banquet Chamber site and outside the cave during the closed-door experiment, 23-26 February 2000. Fig. 5. Results of closed-door experiment for the Banquet Chamber site showing condensation and evaporation rates. The entrance door was closed from 18:00 h on 22 February 2000 to 09:00 h on 26 February 2000.

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 With above in mind, two experiments were conducted to determine what influence the exchange of cave air with outside air has on condensation rates. In the first experiment, the solid cave entrance-door remained closed for 85 hours, thereby minimising cave ventilation. The door was opened for two-to-three minutes about twice an hour during the business day (0900 to 1730 h) to give entry to tour groups In the second experiment, the solid door at the upper entrance was left open continuously for 87 consecutive hours, thus facilitating continuous air exchange with the outside. Conditions in the cave are represented by measurements taken at the Banquet Chamber site (Fig. 1). Thermal conditions inside and outside the cave during these experiments are shown in Figures 4 and 6. The effects on condensation are shown in Figures 5 and 7. On both occasions airflow in both directions through the cave was recorded. When outside temperatures were lower than the cave air temperatures, upward flow occurred and the cave cooled. When the outside temperature was higher than cave air temperatur es, downward flow took place and a warming of cave air occurred. In Figures 5 and 7, a rising trend indicates that condensation (C(+ve)) is occurring while a downward trend indicates that evaporation of condensate (C(-ve)) is taking place. Fig. 6. Air temperature in the cave at the Banquet Chamber site and outside the cave during the open-door experiment, 2-5 March 2000. Fig. 7. Results of the open-door experiment for the Banquet Ch amber site showing condensation and evaporation rates. The entrance door was kept open from 09:00 h on 2 March 2000 to 18:00 h on 5 March 2000.

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 In the door-closed experiment (Figures 4 and 5) airflow through the cave was kept to a minimum, despite a strong cave-to-out side thermal gradient (Fig. 4). The results show a small vapour flux hovering just above and just below zero (Fig. 5). A near equilibrium moisture balance was sustained over the period: C(+ve) = 10.3 g m-2, C(-ve) = 9.9 g m-2. In the door-open experiment (Figs. 6 and 7) evaporation rates in the cave are up to five times larger than on nights when the entrance-door was shut (Figs. 4 and 5). The largest evaporation rate was recorded on the third day of the door-open experiment at 0700 h, when the temperature dropped to 14.8 C in the Banquet Camber and evaporation rose to 2.41 g m-2 h-1 (Fig. 7). Generalised statements can be made about controls on cave microclimate. In conditions where outside air are warmer than the cave air, the relatively cool cave air drains from the cave via the lower entrance and is replaced through the upper entrance by warm outside air (de Freitas et al ., 1982). As the air moves deeper into the cave it is cooled more, reducing its moisture holding capacity further which causes more condensation. Condensation occurs when the dewpoint temperature of the air is equal to or greater than the dewpoint temperature of the surface boundary layer of the cave rock. During conditions in which outside air is cooler than cave air, the process is reversed. Cool and relatively dry air enters the cave through the Lower Entrance. There is an immediate transfer of sensible heat and vapour into the colder air because of the large heat and vapour gradient. Evaporation then occurs. The further the air moves into the cave, air temperature increases and the heat and vapour gradient decreases until equilibrium is reached with the cave environment (de Freitas et al ., 1982). If the air is saturated, an increase in temperature increases its moisture holding capacity and further evaporation occurs. This is the reason why significant amounts of evaporation can occur even when relative humidity reaches 100 per cent (de Freitas and Littlejohn, 1987). Conclusions Condensation is an im portant atmospheric environmental process but it has been neglected in climate research, especially in cave microclimatology where condensation is recognised as a vital component of the cave environment. It is important because the condensation/evaporation process leads to weathering of cave surfaces. Water vapour loaded with carbon dioxide condenses on the limestone or calcite leading to corrosion, while evaporation leaves residual flaky, unsightly deposits of calcite. High carbon dioxide levels in the cave brought about by the presence of large numbers of visitors may exacerbate this. Also, air exchange with the outside and, therefore, the potential for condensation and evaporation, is affected by air movement to and from the cave through entrances. However none of this can be reliably assessed, and then if necessary controlled, until amounts and rates of condensationevaporation can be predicted and the processes that determine them understood. Here the nature and performance of an explanatory model of processes leading to condensation is described using data based on measurements of condensa tion and evaporation as part of a single continuous process of water vapour flux. The results show that the model works well. Condensation is a function of the vapour gradient between rock surfaces in the cave and cave air. The size of the gradient is largely determined by air exchange with the outside. Given that rock-surface temperature in the cave does not vary much, condensation is essentiall y a function of cave air temperature and the processes that affect it, mainly, air exchange with outside. The results show that condensation can be controlled by controlling ventilation of the cave, in this case, by opening or closing the entrance-door. By facilitating ventilation during warm conditions outside, condensation occurs and condensation rates rise as air temperatures rises. During cooler conditions outside or at night, the cave ventilation leads to evaporation and cave drying. To increase condensation rates, ventilation needs to be encouraged (cave entran ce-door opened) whenever outside temperatures are higher than cave air temperatures; downward "summer" flow will then occur, the cave air will warm up, and rates will increase. To reduce condensation rates or induce negative rates (evaporation and cave drying), the cave entrance-door needs to be kept open when outside air temperatures are lower than the cave air temperature. During the cold months, as cave air temperatures are lower than rock-surface temperatures, condensate evaporates because the vapour flux is away fro m the rock surfaces. Generally speaking, only very small to nil rates of condensation occur during the cold months. Condensation rates will only increase during mild winter days when outside temperatures exceed cave air temperatures. To incr ease condensation at these times, the solid upper entrance-door should be opened.

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C. R. de Freitas and A. Schmekal / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 The results provide insight into the environmental effects of management induced changes, but there is need for more work on caves in other climate regimes. Future research should also aim to develop an understanding of the role of condensation in the water and energy balance of caves, especially large systems. Other work might focus on spatial variation of condensation through large caves and factors that affect the geochemical composition of condensate. Acknowledgements This work was funded in part by New Zealand Department of Conservation Research Grant 3272. References Barthow D.C. 1988. The Waitomo Glowworm Cave. In Crossley P (ed). The New Zealand Cave Atlas: North Island New Zealand Speleological Society, Occasional Publication No. 6: Waitomo, 287. De Freitas C.R. 1998. Cave monitoring and management: The Glowworm Cave, New Zealand. In: Cave and Karst Management in Australasia XII. Proceedings of the Twelfth Australasian Conference on Cave and Karst Management Waitomo, Australasian Cave and Karst Management Association, Carlton South, Victoria, 55-66. De Freitas C.R. and Banbury K. 1999. Build up and diffusion of carbon dioxide in cave air in relation to visitor numbers at the Glowworm Cave, New Zealand. In: Cave Management in Australasia XIII Proceedings of the Thirteenth Australasian Conference on Cave and Karst Management, Mount Gambier, South Australia. Australasian Cave and Karst Management Association, Carlton South, Victoria, 84-89. De Freitas C.R. and Littlejohn R.N. 1987. Cave climate: assessment of heat and moisture exchange. International Journal of Climatology 7 : 553-569. De Freitas C.R. Littlejohn R.N., Clarkson, T.S. Kristament IS 1982. Cave climate: assessment of airflow and ventilation. International Journal of Climatology 2 : 383-397. De Freitas, C.R. and Schmekal, A.A. 2003. Condensation as a microclimate process: Measurement, numerical simulation and prediction in the Glowworm Tourist Cave, New Zealand. International Journal of Climatology 23 (5), 557-575. Dublyansky V.N. and Dublyansky Y.V. 1998. The problem of condensation in karst studies. Journal of Cave and Karst Studies 60 (1): 3-17. Dublyansky V.N. and Dublyansky Y.V. 2000. The role of condensation in karst hydrogeology and speleogenesis. In: A.B Klimchouk D.C., Ford A.N. Palmer & W. Dreybrodt Eds : Speleogenesis, Evolution of Karst Aquifers : Huntsville, National Speleological Society: 100112. Ford T.D. and Williams P.W. 1989. Karst Geomorphology and Hydrology Unwin Hyman: London. Gillieson D. 1996. Caves: Processes, Development, and Management Blackwell, Oxford, 324pp. Grace J. 1983. Plant-Atmosphere Relationships. Chapman Hall, London, 92 p. Jensen D. 1983. Computer simulation of an aerological diagram. Meteorology Australia 3 (2): 13-16. Monteith J.L. 1957. Dew. Quarterly Journal of the Royal Meteorological Society 83 : 322-341. Neiburger M., Edinger J.G. and Bonner W.D. 1982. Understanding Our Atmospheric Environment (second edition). W.H. Freeman and Company, San Francisco. Pedro M.J. and Gillespie T.J. 1982. Estimating dew duration. I. Utilizing micrometeorological data. Agriculture Meteorology 25 : 283-296. Tarhule-Lips R.F.A. and Ford D.C. 1998. Condensation corrosion in caves on Cayman Brac and Isla de Mona. Journal of Cave and Karst Studies 60 (2): 84-95.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info Condensation corrosion: a theoretical approach Wolfgang Dreybrodt (1,2), Franci Gabrovšek (2) and Matija Perne (3) (1) Karst Processes Research Group, University of Bremen, Germany (2) Karst Research Institute ZRC SAZU, Postojna, Slovenia (3) Student, Faculty of Mathematics and Physics, University of Ljubljana, Slovenia Re-published from: Acta Carsologica 2005, 34 (2), 317-348 Abstract Condensation of water from warm, humid air to cold rock walls in caves is regarded to play a si gnificant role in speleogenesis. The water condensing to the cave walls qu ickly attains equilibrium with the carbon dioxide in the surrounding air, and conseque ntly dissolves limestone or gypsum fo rming various types of macro,m eso-, and micromorphologies. In th is paper we present the basic physical principles of condensation and give equations, which allow a satisfactory estimation of condens ation rates. Water condensing to a cooler wall releases heat of condensation, whic h raises the temperature of the wall thus reducing the temperatu re difference T between the warm air and the cave wall. Furthermore one has to take into account the heat flux from the air to the cave wall. This defines the boundary conditions for the equation of heat conduction. For a constant te mperature of the air initial condensation rates are high but then drop down rapidly by orde rs of magnitude during the firs t few days. Finally constant condensation rates are attained, wh en the heat flux into the rock is fully transmitted to the su rface of the ka rst plateau. Fo r spherical and cylindrical conduits these can be obtained as a function of the depth Z below the surface. When diurnal or seasonal variati ons of the air temperature are active as is the case close to cave entrances, condensation rates can become quite significant, up t o about 106 m/year. The theoretical results are applied also to corrosion of speleothems and the formation of rhrenkarren as described by Simms (2003). To convert condensation rates into retreat of bedr ock the saturation state of the solution must be known. In the appendix we present experiments, which prove that in any case th e solution flowing off the rock is saturated with respect to li mestone or gypsum, respectively. Keywords: condensation in caves, condensation corrosion. Introduction Water vapor from a cave atmosphere condensing to the wa lls of a cave creates a water film in equilibrium with the partial pressure pCO2 of the cave atmosphere. This solution is therefore aggressive to limestone and the dissolution process based on it has been termed condensation corrosion. The most recent and comprehensive review on conden sation in karst and its role on hydrology and speleogenesis has been published by Dublyansky and Dublyansky (2000) and Klimchouk et al. (1996). They report that in summer condensation supplies a significant amount of water (up to 20% of the total dry reason run-off) to karst systems. They also discuss the role, condensation corrosion plays in sculpturing cupolas in limestone and in gypsum caves, when evaporation from open water surfaces at elevated temperature, above that of the cave wall, produces warm cave air saturated with water va por. Condensation of water to the cave walls supplies an aggressive solution, which runs down the cave wall and is replaced by fresh, condensing water. Cigna and Forti (1986) and Calaforra at al. ( 1993) have reported on fiel d measurements of conde nsation. Large cupolas in limestone and gypsum caves are explained by this mechanism which schematically is shown in Fig. 1 (Audra et al. 2002).

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 2air CO p 2CO p Fig. 1. Water evaporates from an open surface of elevated temperature Ta to the cave air in thermal equilibrium with the water. Vapor pressure at temperature Ta is given by pa. The partial pressure in the air is in equilibrium with in the water. The vapor condenses at the cave wall with temperature Tf < Ta and flows back as a thin film, designated by arrows. These water films become saturated with respect to the mineral composing the rock (limestone or gypsum). The free surface of water at elevated temperature Ta could be a lake of hydrothermal water. Water condenses at the cave walls if their temperature Tf is below the dew point of the air. The condensed water flows back in a thin film covering the cave walls. As long as the temperatures Ta and Tf are constant in time this is a continuous, everlasting process. To estimate the dissolution rate one has to know the wall temperature Tf. If initially the cave wall and the surrounding rock is at temperature Tfi, water from the air with temperature Ta" Tfi starts to condense to the wall. This causes a continuous flow of heat into the rock A small part of it is due to heat transfer from the warmer air into the cold rock. The major part st ems from release of heat of condensation from the condensing water. This heat transfer increases the temp erature of the cave wall, causing reduction of the rate of condensing water. Eventually a stationary state is reached and the temperature of the cave wall, Tf stat becomes independent of time. It is also possible that the temperature of the op en surface changes diurnally or seasonally, e.g. a river flowing into the cave at elevated temperature during day time (summer) and at low temperature during night (winter). In this paper, we address the question, what are the rates of condensation and what average annual retreat of bedrock follows as consequence. The first part of this paper presents a theory of cond ensation rates for the stationary state of condensation for time independent temperature of the cave air, full y saturated with water vapor. Then we will turn to situations where the air temperature in the cave cha nges diurnally or seasonally. These results will be also applied to corrosion of speleothems, which is an impo rtant topic in conservation of tourist caves (Avramidis et al. 2001). A recently observed form of small scale c upolas of a few cm diameter and a length of about 10 cm, growing upwards from bedding planes of limestone at a lake shore (Simms, 2003) is also explained by our theoretical findings. To translate rates of condensation into bedrock retreat in m/year one has to know the saturation state of the water, when it flows off the rock. In the appendi x we present laboratory experiments on limestone and gypsum, which prove that those waters are saturated w ith respect to limestone and gypsum, respectively, in all relevant situations. This work shall provide a theoretical basis to the interpretation of field data. Therefore we assume two scenarios of boundary conditions: a) stable, time independent temperature of the cave air or b) periodic variations in temperature, diurnally or seasonally. We do not ask the question, how these conditions arise in detail. This question deals with cave climate and is a complex subject (Badino 1995; Wigley and Brown 1976). Our question to be answered is: How effective under given boundary conditions is water condensation to the cave walls or to speleothems, and wh at are the rates of condensation corrosion?

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 Basic theory The amount of water condensing to an exposed rock surface per unit of time and surface area determines the retreat of rock by dissolution. Therefore, it is of utmost importance to give a reliable estimation. We assume that the exposed surface of the rock is cove red by a thin film of condensed water, about 10-2 cm deep. Such a thin film will quickly come to th ermal equilibrium with the temperature Tf of the cave wall. It keeps its constant depth since there is flow from the rock surface down to the cave floor, which compensates for condensing water. Water condenses to the film, if the partial pressure Pa of vapor in the cave air exceeds the vapor pressure Pf at the temperature Tf of the water film. In the following, we assume that the cave air is saturated, i.e. relative humidity is 100%. For condensation, water molecules must be transported to the cave wall, where they attach to the water film. We assume that the cave air is well mixed, but close to the rock a diffusion boundary of thickness #D [m] exists. Transport of water molecules is effected th rough this layer by molecular diffusion. The flux of vapor Fv to the surface of the film is given by Fick’s law as -2-1() [gms] !"m Vaf DD Fcc# 1 where Dm is the constant of molecular diffusion of water molecules in air (Dm = 2.5 10-5 m2/s), ca and cf are the concentrations in g/m3 of water vapor at temperatures Ta and Tf respectively. By use of the equation for ideal gases this can be written as -21() [gms]"" !$$af m v DaPP D FM RT# 2 Pa and Pf are the vapor pressures measured in Pa, Ta is the temperature of the cave air in K. In all following calculations we use Ta=300 K. M = 18 g/mol is the molecular weight of water and R is the universal gas constant R = 8.314 J mol-1K-1. Pa-Pf can be approximated by () [Pa] % "!" %afafP PPTT T 3 where P/ T = 100 Pa/K for the temperatures of interest. The vapor condensing to the water film releases heat of condensation qc = 2450 J/g. This is a very high amount of energy. A condensation rate of 1 mm/day = 1 kg/(m2day) causes a heat flux of 28 W/m2, equivalent to the heat produced by lightening the cave by electrical bulbs of 28 W, with one bulb on each square meter of the cave walls. The flux of heat Fq released by condensation is given by -2 -2 af4.4210 () = (T-T) [Wm] %$ !$!$$$$" %m qcvcaf DaDDMP FqFqTT RTT## 4 An additional flux from the warm cave air to the cooler water film is given by -2 -2 af2.610 () = (T-T) [Wm] $ !"a caf TTk FTT## 5 ka is the thermal conductivity of air ( 2.610-2 Wm-1K-1) and #T is the thickness of the thermal boundary layer, which is related to #D by #D = #T(Sc/Pr)1/3. Sc is the Schmidt number for diffusion and Pr the Prandtl number for heat convection(Beek and Muttzall 1975). For air, Sc=Pr=1. Therefore #D= #T= # The total heat flux to the water film is the sum of Eqs. 4 and 5

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 -1-1(), k=0.0702 WmK !"totafk FTT# 6 This heat flux causes an increase of the temperature of the water film, until the heat flux transported into the rock equals the heat flux from the cave air into the water film. If one knows the temperature Tf of the water film, equal to that of the rock surface the amount of condensing water can be obtained by inserting Eq. 3 into Eq. 2, as 5-1-1 v(), h1.810gmK"" &' % &' !$$$!"!$ () () % *+ *+af mv vaf aTT DPh FMTT TRT## 7 In a first approach, we reduce the problem to a si mple one-dimensional setting. Fig. 2 shows this. An extended large cave is located at depth Z below the su rface. We regard only the vertical heat flux towards the surface, thus reducing the problem to one dimension in z-direction. This gives the following situation. Fig. 2. A large extended hall at depth Z below ground is filled with air of 100% humidity at constant temperature Ta. Due to heating from below Ta is independent of time. Initially the temperature of the covering rock is the average annual temperature T0 of the surface. As soon as heating by the cave air starts the wall temperature Tf increases. Steady state is reached wh en the heat flow (gray arrows) to the cave ceiling equals that at the surface. At z = 0 the temperature T0 is kept constant for t > 0. The cave roof at depth Z experiences a heat flow Ftot given by Eq. 6. A solution to this problem is given in Carslaw a nd Jaeger (1959, p.125). We will not give the complete solution here. We extract what we need for further discussion: The solution has a stationary part and the surface temperature of the rock at the cave wall becomes () 1 "!" ,stat ffiafiA TTTT A 8 where A is given by !r Z k A k # 9 kr is the termal conductivity of the rock. kr=1.3 W/(m $ K) for limestone and 0.5 W/(m $ K) for gypsum. The stationary temperature Tf stat is reached after an exponential appro ach of the wall temperature with time constant %S = Z2/( &2' ), where is the thermal diffusivity of the rock, 5.610-7m2/s for limestone and 3.610 -7m2/s for gypsum. The value of & is a function of A by the relation & cot & -A = 0, and is listed by Carslaw and Jaeger (1959). Its value ranges between ( /2 and ( from A = 0 to A = Initially, for times t<0.5 years, however, the rise of surface temperature is much faster It can be approximated by the solution, where Z is infinite, or in other words, where the rock in the cave is limited by a semi -infinite region of rock. In this case the surface temperature can be written (Carslaw and Jaeger 1959):

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 2 22()1experfc &' &'&' "!""$ () ()() () *+*+ *+ffiafi rrkk TTTTtt kk.. ## 10 Tfi is the initial temperature of the rock. For our case with finite Z, this solution is valid for t << Z2, when the temperature front migrating into the rock has penetrated only a distance dt ., very small against its thickness Z. Fig. 3. Time dependence of the temperature difference Ta-Tf(t) normalized to initial temperature difference Ta-Tfi. a) Results for the first 24 hours. b) Results for long times. Until time %f normalised temperature difference rise to 0.84. After this time it rises according to Eq.11 until finally it approaches a stationary state after time 5 s / The black curve is calculated with # =10-3m and Z=10 m and the blue one for Z=30 m. The red curve is calculated for Z= a semi infinite medium. Note the coincidence of the curves in early evolu tion. The normalized temperature differen ce in the stationary state is given by A/(A+1).

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 Fig. 3 shows the exact temperature dependence for limestone with Z = 10 m, 30 m, and ) It is of utmost importance to note that the initial rise of temperatur e is independent of Z and controlled by the quantity rk k. #. For times 29 &' 0 () *+rk t k# ., Eq. 10 can be approximated by 1212 11 1 &' "!""$ () *+r ffiafik TtTTT k t# 3. 11 The rates of condensing water for the the fast initial rise of temperature can now be obtained from Eqns. 7, 10 and 11 12 2 22experfc &' &'&' % &' !$"$ () ()() () () % *+ *+*+ *+m Vafi arrDPMkk FTTtt TRTkk.. ### 12 For times 29 &' 0 () *+rk t k# ., a slow decline of the temperature difference Ta-Tf(t) towards stationary state yields rates 12-2111 gms"% &' &' 4 5 !$" () () 6 7 % *+ *+slow r Vmafi aPMk FDTT TRTk t3. 13 It is important to note th at this is independent of # The minimum rate of condensation is reached at stationary state after time 5 %s. 12 5-2-11.810 gms 1"" % &' 4 5 !$"!$ () 6 7 %, *+ ,afi stac m Vafi a rTT DPMA FTT k TRTA Z k# # 14 Note that for Z >> kr# /k the condensation rate becomes independent of # This is true for Z 10 m and # <10-3 m. The decay time towards this stationary state is given by 2 2 [s] 8sZ/ 9. 15 At a depth of 100 m this is about 127 years, a short time with respect to the time scales of cave evolution. Fig. 4a shows the rates of condensa tion for Z = 100 m and various values of # Fig. 4b depicts the amount of condensed water after time t, which is obtaine d by integration of the curves in Fig. 4a. In summary, the following picture emer ges. For times smaller than 0.5 %S, the temperature increase of the cave wall can be well approximated by Eqns.10 and 11. Later, for times t > 5 %s the stationary state is reached by an exponential approach an d the temperature becomes constant. Fig. 3 shows this for # = 0.001 m. The black line shows (Tf – Tfi)/(Ta-Tfi) for a depth of Z = 10 m, the blue line gives the temperature rise for Z = 30 m. The red curve shows the temperature for the semi-infinite plane with Z= ) Condensation in stationary boundary conditions Our one-dimensional model is idealistic because it requi res a cave chamber with horizontal dimension, one order of magnitude larger than its height.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 Fig. 4. a) Rates of condensation and retreat of bedrock as func tion of time divided by the initial temperature difference Ta-Tfi. Values of # are written at the curves. Note that for large times all curves coincide, i.e the rates become independent of # b) Amount of condensed water per m2 and K of temperature difference and total retreat of bedrock for the cases of Fig.4a. Note that although the variation of # covers two orders of magnitude, the amount of condensed water only weakly depends on # Other geometries are more suitable, e.g. a spherical room with diameter Ds or a cylindrical conduit of diameter Dc and length L, both buried at depth Z. These are shown in Fig. 5a-c. For such situations analytical results are not available. But the general be havior is similar to the idealistic one-dimensional case, and it is possible to obtain the temperature at stationary state, reached after t > 5 %s, by the theory of conduction shape factors (Incropera and DeWitt 2002). At stationary state, the wall of the rock is at

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 temperature Tf stat and the surface temperature at z = 0 is To. See Fig. 2. The total amount of heat flowing from the conduit to the surface is then given by 12 [W] !$"stat rrfoQSkTT 16 where S [m] is the shape factor. The expressions for some shape factors are given in Table 1. The total heat transferred to the wall of the cave is 12!$" s tat wafk QTT: # 17 where is the surface area of the cave. TABLE 1 Shape factors for some typical geometries as shown in Figs. 5a-c. System Restriction Shape factor Sphere in a semi infinite medium (Fig. 5a) z > D/2 2D 1D/4Z 3 Cylinder in a semi infinite medium (Fig. 5b) L >> D z > 3D/2 2L ln(4Z/D) 3 Slab (Fig. 5c) /Z Fig. 5. a-c) Geometrical configurations for the shape factors in Table 1. d) Condensation rates and retreat of bedrock in the stationary state in dependence on depth Z for spheres and cylindrical conduits with various diameters. D(0) is the initial diameter. The blue curve represents one dimensional scen ario (slab with thickness Z) of Fig. 5c. See also Fig. 2.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 Conservation of energy requires Qw = Qr. Therefore one finds 12 121 1 "!" $ $$stat afao rTTTT k kS: # 18 From this the rate of condensed wate r can be calculated by use of Eq. 7. Fig. 5d shows these rates per degree K for circular cave rooms and cylindrical conduits, buried at depth Z. Note that for # << 0.038 /S the rates are independent of # which is true for all practical applications. Spheres show the highest condensation rates, almo st independent of their depth Z below ground, but dependent on their diameter DS. As we will show in the appendix, film s of water running down rock-walls of gypsum or limestone can be regarded as saturated wi th respect to gypsum or limestone. Therefore at a pCO2 = 0.00035 atm, for limestone with saturati on concentration of 60 mg/L, 1 g/m-2s-1 of condensed water corresponds to retreat of rock by 7.5 $ 10-4 m/year using 32500 kg/m ;! as density of a slightly porous limestone. For gypsum with a saturation con centration of 2.5 g/L (independent on pCO2) and density 2300 kg/m3, the retreat of rock can be approximately obta ined by multiplying the values for limestone (pCO2 = 0.00035 atm) by a factor of 50. Now we assume the following situation. A spherical cave of initial diameter Ds(0) is invaded by geothermal waters with temperature Ta, higher than the initial rock temperature Tfi. Water evaporates from a lake, stable for long times. We assume the cave 100 m below ground. Then after about 100 years the stationary state is attained. The change of diameter Ds is then approximated quite accurately by 12541 21.8107.510 2""8<$<$<" ,$s afi s rdD TT kD dt k# [m/year] 19 For limestone one gets 1261 110 [m/year]"8$$$"s afi sdD TT dtD 20 Note that growth rates decrease with 1/DS. Integration of Eq. 20 yields 226()(0)210() ""8$$$"ssafiDtDtTT 21 where t is in years. A cave with initial diameter of 1 m evolves into a cupola of 10 m diameter independent of its depth Z for D/(4Z)>10 in 5 107 years for Ta Tfi = 1 K. A sphere with initial diameter of 0.1 m needs, howev er, only 530000 years to reach a diameter of 1 m for Ta-Tfi= 1 K. For a cylindrical conduit buried at depth Z, 12 =>61 110 m/year 4 ln "8$$"C afi C CdD TT Z dt D D 22 which can be integrated 12226()1(0)1 ()ln(0)ln210 [m/year] 4242"?@?@ "","!$<<" ABAB CDCDafiDtD DtDtTT zz 23

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.10 For Z = 100 m growth from initially 1 m to 10 m takes 2108 years for Ta Tfi = 1 K, and correspondingly 2107 years if the temperature difference is 10 K. Fig. 6 shows growth times for spheres a nd cylindrical conduits at various depths + nd with various diameters. Fig. 6. Diameters of spheres and cylindrical conduits with initial diameter of 0.1, 0.5 and 1 m at depth of 25 m as a function of time. The curves are calculated from equations 21 and 23, respectively. A final statement must be given. In this section we have assumed that temperature difference between the cave and the surface and 100% relative humidity of the cave air are independent of time. This may be the case when hydrothermal waters form open surfaces, fr om which evaporating waters condense at cave walls, delivering a constant flow of water in equilibrium with the pCO2 of the cave atmosphere. These waters attain equilibrium with respect to the so luble rock, either limestone or gypsum (see appendix). The saturation concentration for limestone is 24 3 eqcoc60mg/Lp/3.510 !$$ mg/L, where pCO2 is the partial pressure of CO2 in the cave atmosphere, measured in atm. Our findings of Eq. 21 explain the existence of sphe rical niches and cupola in caves of Hungary (Mller, 1974) in Italy, (Cigna and Forti 1986), and in hypog enetic caves in France (Audra et al. 2002). Sarbu and Lascu (1997) report on the measurements of active condensation corrosion in Movile Cave, Romania. Mostly in nature external boundary conditions are not constant in time. Annual fluctuations in the temperature of the cave air might occur when a river fl ows into a cave, with warm water in summer, which evaporates and condenses at the cave walls. In the wint er time, however, the water is colder than the temperature at the cave walls and condensation stops. In view, that to attain stationary state under time independent boundary conditions takes several years, su ch cases cannot be described by the considerations above. Condensation at periodic boundary conditions As already shown Fig. 4a presents the time dependen ce of the rate of condensing water. A first rapid decline until time t= %f (see Eq.12) 21 9 &' !$ () *+r fk k# / 24 is followed by a slow decline (see Eq. 13) exhibiting a t-1/2time dependence.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.11 After the time t = 5 %S stationary state is attained (see Eq. 15). No te that the time to reach stationary state is controlled by the depth Z of the cave below ground and the thermal diffusivity whereas the decline toward stationary state depends solely on the thermal properties k, and kr and is independent of the depth Z (See Eq.13). Variations of the cave temperature can be caused by diurnal fluctuations when warm humid air flows through the cave during day time, but during the night cold dry air enters. If the initial temperature of the rock is Tfi and the temperature of the cave air with a humidity of 100%, is Ta within the time t = 9(kr# /k)2/ the temperature Tf of the rock has increased such that Ta-Tf = 0.18(Ta-Tfi ). From then on Ta-Tf is given by Eq. 13. Note that for # = 0.001 m, t = 9(kr# /k)2/ = 5.5 $ 103 s whereas the equilibration time 5 %s= 9 $ 105 s for Z = 1 m. Therefore at depths Z > 1m daily variations in Ta cause temperature changes of the rock, which are governed by the entity (kr/k). When cold air, at temperature Tn colder than the actual temperature Tf at the cave walls enters into the cave during the night, condensation stops. Two possibilities can be envisaged. The water, whic h has condensed during day time evaporates during the night. In this case, in a first approximation, the temperature of the cave wall drops towards the temperature of the cold air in about the same time as is needed to appr oach to the temper ature of the warm air during day time. This is shown by Fig. 7a. As a result calcium carbonate dissolved during the day, will be precipitated during the night, and this process disintegra tes the texture of the rock, leaving a weathered rind of corroded material (Auler and Smart 2004). For # = 0.001 m, the amount of water condensing within 10 hours is about 100 g/(m2K) as can be obtained by integrating Eqs. 12 and 13. See Fig. 4b, which shows the amount of condensed water per square meter as a function of time for various values of # and a temperature difference Ta-Tfi of 1 K. Fig. 7. Wall temperature for diurnal temperature variations of Ta in the cave. For simplicity an abrupt change between day temperature Tday and lower night temperature Tnight is assumed. The initial temperature of the cave wall Tfi
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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.12 The time dependence of cooling or heating is given by the dimensionless variable k( t)1/2 /(kr# ). With k lower by a factor of 2.7 cooling takes a time, longer by a factor of 7 than heating. Therefore during the cooling period the wall of the rock cools down slower. This way, after several cycles a stationary amplitude is established with constant temperature differences of the rock wall, and condensation rates are reduced. Fig. 7b shows this schematic concept. These two scenar ios are extremes. Depending on the relative humidity of the air during the day and during the night, rate s will be between these tw o extreme limits. We will discuss this latter in detail. Seasonal variations can be caused when warm water fr om the surface enters into the cave in summer time and condensation takes place, whereas in winter time, when cold water flows into the cave, condensation stops. For most cases of interest, with caves deeper than 5 m, the time to r each a stationary state is longer than 2 years. The amount of water condensing during th e summer period can be obtained by integration of Eqs. 11, 12. See also Fig. 4. It can be approximated by 12 20.520 g/m 45 45 !", 67 67afiCQTTt 25 for tc 4 $ 104 s, where tc is the period of condensation in seconds. For tc = 8.4 $ 106 s 100 days this corresponds to about 1500 g/(m2K), equivalent to a retreat of cave wall by 3.510-8 m/(yK). During winter time condensation stops and the cave walls have sufficient time to cool to low temperatures. Then in summer condensation starts again. Note that so far all calculat ions of retreat of rock are based on the assumption that pCO2 in the cave is at atmospheric level with 3.5 $ 10-4 atm, and that the solution flowing off the wall has attained saturation with respect to calcite. For elevated values of pCO2 in the cave all numbers given so far must be multiplied by (pCO2/0.00035)1/3 to account for elevated pressure pCO2 in the cave. Condensation corrosion on speleothems Many researchers have observed features of surface corrosion on speleothems, which they interpret as a results of condensation corrosion (Auler and Smart 2 004; Dublyansky and Dublyansky 2000; Tarhule-Lips and Ford 1998). In this section th e physical background is discussed. Fig. 8. Heat rates transferred via corresponding surface of a stalagmite with length L and radius r. k=0.07 Wm-1K-1, kr=1.3 Wm-1K-1. Even at a temperature grad ient of 10K/m at the base, F3<0.1K. T is the surface temperature of the stalagmite. Fig. 8 depicts the thermal boundary conditions for a stal agmite. At the outer surface heat flux is given by the action of condensation. At the base heat is transf erred into the colder base rock. As can be calculated from Fig. 8 this heat flow is negligible. The decay ti me which is needed in such a case to approach thermal equilibrium (Luikov, 1968, page 217) is %Stal = D/(41 2' ), where 1 is the first root of the equation r2k tg Dk # E!$E. For D = 1 m, 1 is (2/4 and %Stal = 2 days. When thermal equilibrium is reached after 5 %s

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.13 condensation stops. Therefore to keep condensation ac tive, diurnal variations in cave temperature are necessary. When the equilibration time 5 $%Stal is longer than a day, and when during the night evaporation takes place a stationary state with high average rates is reached after some days. Lower rates are effective when evaporation is excluded during the night. We restrict to a more simple approximation. We assume stalagmites with diameters less than 0.3 m corresponding to %Stal = 0.2 day. In this case the stalagmite practically reaches temperature Ta during day time. Even if no evaporation is present during the night, the time to cool down is longer only a factor of 2. This results from the dependence of the root on the heat transfer coefficient k (Luikov 1968, page 217). The maximal amount of condensing water during one da y can be estimated from conservation of energy. The total heat transferred to the stalagmite must be equal to the increase of internal energy due to heating from the initial temperature Ti at t = 0 to the stationary temperature Ta. According to Eqs. 4 and 5 about 70% of the total heat transfer results from condensation. Therefore to a sufficiently good approximation one has 0.7 $!$$$%CCpsqMcMT 26 where Mc is the mass of condensed water during 1 day, cp is the specific heat of limestone (0.88 kJ/kg K) and T % is the change of temperature until th e end of heating. It is close to T % = (Tday-Tnight). Ms is the mass of the stalagmite with diameter D and length L. From Eq. 26 one finds the amount of water which conde nses during one period of heating. Dividing this by the surface area of the stalagmite giv es the amount of water condensed per m2 during the heating period as 2 00.7 0.17 kg/m 4 45 !$%8% 67p Cc FDTDT q; 27 this corresponds to retreat of surface by 9 63.810 [m/day] or 1.410 [m/year]" "!$% !$%c yDT DTF F 28 for D < 0.3 m. Note that for compact st alagmites we use the density of 2700 kg/m3. For larger diameter D > 1 m the equilibration times for heating and cooling are longer than 5 days. If such a stalagmite experiences condensatio nal heating with equilibration time %1= %Stal during the time t1 and cooling with equilibration time %2, during the time t2 after several cycles the maximum temperature and the minimum temperature become independent on time. The temperature difference Tmax –Tmin can be estimated in the fo llowing way. During heating the temperature Th(t) of stalagmite to an acceptable approximation is given by 1212 minmin 1()1exp &' &' "!""" () () () *+ *+hht TtTTT/ 29 where Tmin is the minimum temperature at the end of prior cooling, and Th is the temperature of the cave air during heating. For the cooling period t2 with equilibration time %2 the temperature Tc of the stalagmite is given by 1212 max 2()exp &' "!"" () *+ccct TtTTT / 30

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.14 where Tc is the air temperature in cave during cooling and Tmax the temperature of the stalagmite at the end of prior heating. From 29 and 30 we get 121212maxminmin11 minmax21()1exp/ ()exp/ "!""" "!""h ccTTTTt TTTTt / / 31 Solving this equation one finds 1212 1 2 1 21212 1122 maxmin 11221exp/1exp/ () 1exp/exp/ """" %!"!" """hctt TTTTT tt / / // 32 To estimate the amount of condensation corrosion, this value of T % must be used in Eq.28 T % in Eq. 32 becomes0.25 () "hcTT for 12 120.5!!tt// and 0.12 () hcTT for 12 120.25!!tt// For 1 / >>1texpansion of the exponents and regarding %1= %Stal yields 2 11 maxmin 2()()%!"$!"$hc Staltt TTTTT D.3 / 33 Therefore one finds from Eq. 28 2 6 1 -7 11.410() [m/year] () =3.310 [m/year], with t12 h"!$"$ $!yearhc hct TT D TT D.3 F 34 It should be noted that this number is valid for D > l m and presents an upper limit, because we have assumed tacitly that the temperature inside the st alagmite is homogenous and equal to the surface temperature. Eqs. 28 and 34 show that for stalagmites with diameters D between 0.1 m to several meters and temperature differences during night and day of 10C, corrosion rates are on the order of 10-4 to 10-5 cm/year. Auler and Smart (2004) have estimated rates of corrosion on stalagmites by measuring the depth of the weathered rind and determining the ag e of the unaltered speleothem calcite below. The highest values they observed was 410-5 cm/year. In most cases the rates were lower by one to two orders of magnitude. Field experiments to measure condensation corrosion Tarhule-Lips and Ford (1998) su spended gypsum plates of about 1 cm in thickness on nylon strings for about one year in flank-margin caves of the Caribbean From the measured weight loss they report corrosion rates of 2.4 10-2 cm/year. The thermal behavior of such isolated plates is very si milar to that of stalagmites with diurnal variation of temperature and Eq. 28 remains valid if one replaces the diameter D by the thickness of the plate. From this one obtains 51.410"F8$ cm/year, assuming T=10 K. This is three orders of magnitude lower than the experimental findings. It should be noted that these gypsum plates reach thermal equilibrium after a time of only a few minutes. After this time condensation stops and renewed cooling and subsequent heating are necessary to revive it. Sarbu and Lascu (1997) report conde nsation rates of water in Movile-Cave, Romania, where they collected water from a 10 cm by 10 cm glass plate suspended at a distance of 10 cm from the cave wall. Movile cave is heated by a hydrothermal lake with wate r temperature of 21C. In its upper level, where the glass plate was suspended air temperature is 21C, but the temperature of the cave walls is between 19.4

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.15 and 15.7C. In this case the glass pl ate achieves thermal equilibrium after a few minutes. The observed high run-off rates of about 20 g/month cannot be explaine d by condensation. From Eq. 27 one obtains only 2.5 g/month for a plate of 1 cm thickness. Summarizing, we state that condensation to small scale objects (0.1-1 cm) is subject to diurnal variations of temperature. In thermal stable cave environments, or where variations of temperature change seasonally, it can be excluded. Rhrenkarren, a small scale example of condensation corrosion Recently Simms (2003) reported on vertical, upwar d tapering tubes in limestone exposed in the epiphreatic zone at the shores of several lakes in Ireland. The dimensions of these solutional features comprising almost perfect circular tubes are between 1 to 5 cm in diameter and up to 30 cm in length. Simms suggests that these “Rhrenkarren” originat e from condensation corrosion within air pockets trapped by seasonal high stands of the la ke. During March 2000 water temperature TW stayed nearly constant, whereas the temperature of the overlying rock showed diurnal variati ons, fluctuating around the water temperature with amplitude of about 5C. In winter time, when surface temperatures become low extended periods of rock temperatures TR below water temperature are likely. Therefore conditions for condensation corrosion to operate are valid. Fig. 9. Rhrenkarren: A slab of rock with a depth of a few meters is exposed to the surface. Its bottom is flooded by the water from the lake. This way air is entrapped. An initial irregularity of radius R (dotted line) grows into a cylindrical shape of length L. The temperature of the air inside is cl ose to the water temperature. In winter time the surface temperature is lower than that of the lake water and condensation is active. The lake water is saturated with respect to calcite such that dissolution of the rock in contact with water is excluded. Since the region of the lakes was glaciated about 15 ky ago, rates of dissolution were estimated about 2.10-3 cm/year, if one assumes continuous dissolution to present. With this information it is possible to test our theoretical predictions. The distance between th e surface of the rock to th e apex of the Rhrenkarren is on the order of meters. Therefore thermal equilibrium is attained in a few days. During winter time an extended period of rock temperatures TR several degrees below that of the lake water supports condensation in a steady state. To estimate the rat es we use Eq. 19. In the initial state of the evolution of the Rhrenkarren air is entrapped on irregularities of the rock. The geom etry of this bubble could be approximated by a halfsphere with radius R as depicted in Fig. 9. The ai r entrapped is stagnant, due to the closed conditions. Therefore, diffusion of vapor from the water surface to the rock wall is through this stagnant air. # in Eq. 19 must therefore be replaced by the length L of the tube. The retreat of bedrock is focused to the apex, because h eat flow is highest there, whereas the walls exhibit lower heat flow. This way, a circular tube can propagate upwards into the bedrock. Using Eq. 19 one finds 8 2() 1.3510 [m/year] 5.410" "" !$ ,$WRdLTT dtLR 35 Initially L = R, therefore to a good approximation 8() 1.3510 [m/year]"" !$WRdLTT dtL 36

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.16 Initially with L = 0.01 m the growth rate is 1.3 10-4 cm/year K. Since condensation is active only during cold weather this must be reduced by a factor of 0.3 or so, to about 5 $ 10-5 cm/(year K). When the length increases the growth rate drops with 1/L. During the growth of the tube the entrapped air is in contact with the lake water and consequently pCO2 in this air is in equilibrium with the pCO2 in the water, which can enhance disso lution rates by a factor of two for pCO2=2 $ 10-3 atm, likely in lake water. Therefore with a temperature difference of 10 K maximal rates are 10-3 cm/year, dropping to 10-4 cm/year when a depth L of 10 cm is reached. Discussion To illustrate our theoretical results we give a numerical solution of a representative example. Fig. 10. Modeling domain of a rectangular conduit. Only the right half is shown, because of the mirror symmetry. The surface temperature is T=To. All other boundaries are impermeable for heat flow. Fig. 10 represents a rectangular cave conduit parallel to the surface of a limestone plateau. It is located at a depth 25 m below ground and its cross section is 10 m x 10 m. Inside this cave the air temperature is 10C and relative humidity is 100%. The temperature at the surface is 0C. Note that only the temperature difference is significant. The other boundaries are assume d to be adiabatic, and cannot transmit heat. This is an approximation, which is valid for t < y2/ where y is the distance of the cave to the outer limit. t = y2/ is the time when the thermal front reaches the adiaba tic boundaries. We solve the differential heat conduction equation 22 22TTT tyz &' GGG !.$, () GGG *+ 37 with the boundary conditions, as discussed above by a finite difference program. Fig. 11 illustrates the results. It shows isotherms as they evolve in time. At the beginning (1 $ 106 s, Fig. 11a) an almost circular temperature field has deve loped, symmetrical around the conduit. The temperature has changed only in the vicinity of the cave. After 107 s (Fig. 11b) the distance the temperature front has propagated is about, (107' )1/2 = 2.4 m, close to what we see in Fig. 11a. After 2 $ 108 s (Fig. 11c). After 1 $ 109 s (Fig. 11e) a thermal gradient develops, directed fro m ceiling of the cave toward the surface. Most of the heat from the ceiling flows vertically to the surface. This can be seen from the flow lines depicted in Fig. 11. In this region the temperature distribution becomes stable in time, as can be visualized in Figs. 11d, e, f. Keeping in mind that 228 s25/10s /!3.8 (confer Eq. 15) this is a reasonable result.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.17 Fig. 11. Evolution of isotherms and heat flow, represented by gray arrows in the domain of Fig. 10 at various times. The dotted lines in Fig.11d separate the regions of heat flow from the ceiling, the side walls, and the bottom surface of the cave. The dotted lines in Fig. 11f border the regions of heat flow from the ceiling, the side wall and the floor towards the surface. The averag e distance of the heat flow from the side wall to the surface is larger by about a factor of 2 compared to that of the ceiling. Th erefore the time until the temperature field becomes stationary there, is longer by a factor of 4. Finally the heat flow from the floor exhibits still longer effective flow paths towards the surface. C onsequently the time until a stationary temperature distribution is established there, becomes also longer by a factor of about 4, compared to its neighboring region.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.18 This behavior is shown in Fig. 12. Its left hand ordinate shows the rates of condensing water in g/m2s as a function of time and the right hand or dinate the retreat of the cave walls for a temperature difference of 1 K. For times t < s 10 19$ all curves are identical because as show n by Figs. 11a-d the boundaries do not yet influence the temperature distribution. Initially, as predicte d, a sharp decline is seen which is followed by an intermediate and a slow region of decreasing rates. Compare to Figs. 3 and 4. Fig. 12. Time dependence of condensation rates and retreat of bedrock for the ceiling, the side wall and bottom wall of the conduit for times smaller than 100 hours (a) and for long times (b) Note that initial rates (up to 10 y) are equal. First, after 20 years the curve for the ceiling become s constant, whereas the curve for the side wall shows still a slow decline and becomes constant after 40 years. A similar behavior is also seen for the rates from the floor. From the rate of condensation to the ceiling one finds a retreat of bedrock by 2.310-8 m/(year $ K). From the approximation expressed by Eq. 22 one finds 2.410-8 m/(year $ K). In view of the approximations used, this proves our general concept. Summarizing, the general properties of heat flow a nd condensation rates estimated from the idealized onedimensional model is confirmed by numerical calc ulations on 2D-models, a pproximating reality more accurately. This shows that the simple analytical expressions of Eqs.7, 10, 14 and 18 are sufficiently accurate to estimate rates of condensati ons in relevant geological situations.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.19 We now turn to periodic variations of the temperat ure of the cave air. To this end we assume Ta=15C during the first 12 hours of the day and Ta=5C during the 12h of the “night”. Fig. 13a shows the results if one assumes that the water, which accumulates during th e day, evaporates during the night. If no evaporation takes place during the night, cooling becomes slower than heating during the day. Fig.13b shows the variations of wall temperature in this case. It tak es about 20 days until the temperature amplitudes become constant. Fig. 13. a) Variation of temperature for diurnal temperature changes in the domain of Fig.10. The dotted line shows the temperature Ta of the air in the cave, varying between 5C and 15C Surface temperature is 0C. Condensation during day, evaporation during night. b) Accumulated total amount of condensed water and total retreat of bedrock as a function of time. Fig. 14 shows the total amount of water which has c ondensed to one square meter of the cave wall as a function of time for both cases. The stationary rat es depend only on the difference between day and night temperatures, and are independent on the surface temperature at z=0. Independently on that temperature, in the heating period the temperature of the rock approaches Th closely. Then during cooling is cannot drop below Tc. Finally a stationary state is reached when this minimum temperature at the cave wall determines the initial conditions at each heating period, and maxminTTT %!" is independent on the surface temperature at z = 0. Fig. 14. Boundary conditions as in previous Fig. 13. Accumulated total amount of condensed water and total retreat of bedrock as a function of time.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.20 One has to consider that such diurnal variations are operative only during the summer season, about one third of a year. Therefore the retreat of bedrock must be reduced correspondingly. This result is independent on the depth Z, as can be visualized from Fig. 3. During first 24 hours and up to 180 days the rise of temperature and condensation rates ar e independent of depth Z. When stationary state is reached one obtains a retreat of bedrock by 10.4 m/year E for the case with nocturnal evaporation and T10C %!H, and 2.7 m/year E if evaporation is absent. See Fig. 14. This is higher by one order of magnitude than retreat of bedrock of 0.2 m/year E for stationary temperature with T10C % !H (see Fig. 5) 2.7 m/year E This is an important result, which shows that signifi cant retreat of rock by cond ensation corrosion is possible at locations, where wet warm air with a dew point temp erature above that of the cave wall enters during the day. But during the night cool air with dew point above the actual wall temperature must enter. Such locations are favorable only close to the entrance. Finally we consider seasonal variations with the sa me thermal boundary conditions as in the previous case but with changes of temperature every six m onths. The results are shown in Fig. 15. Fig. 15. a) Variation of temperatures for seasonal changes in temperature between 5C and 15C. Condensation during the warm season, no evaporation during the cold season. b) Accumulated total amount of condensed water and total retreat of bedrock as a function of time. Steady state amplitudes of temperature are obtained after about 10 years. They are about 50% of the initial amplitudes. The corresponding amount of condensed water is given in Fig. 15b For the linear region a retreat of bedrock 0.33 $ 10-6 m/(year K) is obtained. This is about one tenth of the maximal value calculated by the initial condensation by using Eq. 25. These results are independent on the depth Z for the same reasons as discussed above. A final comment must be given. In all our considerations we have assumed that the relative humidity of the cave air is 100 %. Therefore our results give maximal values of bedrock retreat. Furthermore all values gi ven relate to limestone. These values can be approximately converted to gypsum by multiplication by a factor 50, which takes into consideration the higher solubility and lower density of gypsum. Conclusion Condensation of vapor from warm humid air to the cold er walls of a cave is controlled by transport of heat from the warmer air and the heat of conden sation into the surrounding rock. At the onset of condensation to the initially cold cave wall, the wall temperature rapidly increases and then slowly approaches a stationary state. Therefore initial condensa tion rates are high, but at stationary state they are lower by several orders of magnitude.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.21 We have shown that a one dimensional heat transfer model is suitable to estimate the condensation rates within the order of magnitude. From these rate s the retreat of bedrock can be calculated. For constant air temperature Ta in the cave, rates are proportiona l to the temperature difference Ta-To, where To is the average annual temperatur e at the surface of rock massi ve hosting the cave. Retreat of bedrock in this case is in the order of tenths of a micrometer per year at Ta-To=10C. When the cave temperature varies diurnally or seasonally, the driving force for condensation is the amplitude of temperature variations. For diur nal variations, retreat of bedrock is about 3 m/year Efor amplitude of 10C, whereas seasonal variations with the same amplitudes cause a retreat of bedrock by 0.3 m/year E. APPENDIX In our theoretical considerations we have assumed that in any case the water flowing down from the cave walls has become saturated with respect to limestone or gypsum, respectively. This crucial assumption needs experimental proof. We therefore have set up an expe riment, with condensation rates similar to those in nature. Fig. A1. Experimental set up. See text. Fig. A1 shows the experiment schematically. A closed vessel contains water at the bottom, which can be heated by a heating wire, controlled by a contact thermome ter (4). It is stirred by a magnetic stirrer (5). A conductometer (6) is used to measure the conductivity. At the top we have a limestone or gypsum cupola (2) which is cooled by water with temperature of 20 0.2C entering at inlet (3). Inlet (7) provides the possibility for exchanging the interior atmosphere. Thus it is possible to establish a partial pressure of CO2 up to 1 atm. At the upper part of the vessel there is a ring, which collects the condensed water flowing down from the cupola and guides it to the collection bottle (8 ). The temperature difference between the water and the cooler rock cupola can be maintained between 3 C up to 8C. The experiment s were performed with a temperature difference of 5C. It should be noted that this set up does not simulate th e situation in nature, since the water film will attain a steady temperature in a very short time due to c ooling of the cupola. Nevertheless it will answer the question about saturation of the condensed water fl owing down from the rock. Two kinds of experiments have been performed. In the first experiment we collected the water flowi ng down from the rock and analyzed it by standard titration methods for Ca. We first perf ormed the experiment with a cupola ma de of plaster of Paris. In this case the amount of condensed water flowing into bottles was about 15 g/day, equivalent to 1.7 kg/(m2day). The result of the analysis for Ca show ed saturation with respect to gypsum.

PAGE 22

W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.22 Similar experiments were performed using a limestone cupola. Here because of a different geometry the amount of condensing water was reduced to 3 g/day equivalent to 0.69 kg/(m2 day). The experiments were performed with an atmosphere of different CO2-partial pressures (1 $ 10-3, 5 $ 10-3, 1 $ 10-2, 2 $ 10-1, 1 atm). In all cases the collected water was saturated with respect to calcite. In a second kind of experiment the ring was remo ved and the condensed water flows down to the reservoir, thus increasing its conductivity. In all expe riments we observed a linear increase of conductivity with time for several days. For limestone at pCO2 = 1 $ 10-3 atm the rise was 2 S/day which increased up to 20 s/day for pCO2 =1 atm. If the down flowing water is satur ated, then in these experiments the increase !* in conductivity must be proportional to the saturation concentration Cs. From the equilibrium chemistry of the system H2O-CO2-CaCO3 it is known that Cs is given by 1/3 S1CO2 1/3 2CO2CC(P) C(P) !$ %:!$ A1 C1 and C2 are constants. We have plotted the experimental data logarithmically versus pCO2 as depicted in Fig. A2. Clearly we obtain a straight line with slope of 0.33 0.02. This proves the underlying assumption of saturation with respect to calcite. Fig. A2. Increase of conductivity for condensation corrosion on the limestone cupola as a function of CO2pressure. We have furthermore estimated the thickness of the f ilm by removing the condensed water from the rock surface with a tissue and weighting its weight increase. We found a thickness =5 $ 10-3 cm. Using this value it is possible to confirm the experimental results by the following arguments. The dissolution kinetics of calcite covered by thin films of water is given by the rate equation (Buhmann and Dreybrodt, 1985) S 2mmol R(CC) [] cms !I" A2 C is the actual calcite concentration in the film covering the calcite surface. Values of are tabulated and range from 0.3 $ 10-5 up to 3 $ 10-5 cm/s for natural environments increasing with increasing pCO2. The time necessary until such a film achieves 95% saturation is given by T3 [s] F !$ I A3 This gives a maximum value of about 1 hour. From the total amount of water constituting the film (0.22 cm3), and the amount condensing per day (3 cm3) the average time water spends at the surface before flowing back is about 1.5 hours, sufficient to come close to saturation. In the case of gypsum dissolution proceeds by molecular diffusion. Therefore is given by D/ where D 1 $ 10-5 cm2s-1 is the diffusivity of the Ca-ion. Thus 3 $ 10-3 cm/s and T ~ 2s. Thus the achievement of saturation can be taken for granted in both gypsum and limestone caves.

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W. Dreybrodt, F.Gabrovšek and M Perne / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.23 References Audra, P., Bigot, J.Y. and Mocochain, L. 2002. Hypoge nic caves in Provence (France): Specific features and sediments. Acta Carsologica 31/3 : 33-50. Auler, A.S. and Smart, P.L. 2004. Rates of conde nsation corrosion in speleothems. Speleogenesis and Evolution of Karst Aquifers, www.speleogenesis.info 2/2 Avramidis, P., Hong, J., Barnes, C. and James, J.J. 2001. A new method of measuring condensation corrosion. 13th International Congress of Speleology, Brasilia, DF. Badino, G. 1995. Fisica del Clima Sotterraneo. Instituto Italiano di Speleologia, Bologna, 136 pp. Beek, W.J. and Muttzall, K.M.K. 1975. Transport phenomena. Wiley, London, New York, 298 pp. Buhmann, D. and Dreybrodt, W. 1985. The kinetics of calcite dissolution and preci pitation in geologically relevant situations of karst areas.1. Open system. Chemical Geology 48(1-4) : 189-211. Calafora, J.M., Dell'Aglio, A. and Forti, P. 1993. Pr eliminary data on the chemical corrosion in gypsum karst: 1 The Sorbas region (Spain). XI Internationa l Congress of Speleology, Beijing, China, 77-99. Carslaw, H.S. and Jaeger, J.C. 1959. Conduction of heat in solids. Oxford University Press, Oxford, 510 pp. Cigna, A. and Forti, P. 1986. The speleogenetic role of air flow caused by convection. International Journal of Speleology 15 : 41-52. Dublyansky, V.N. and Dublyansky, Y.V. 2000. The role of condensation in karst hydrogeology and speleogenesis. In Speleogenesis: Evolution of karst aquifers, A. Klimchouk, D.C. Fo rd, A. Palmer and W. Dreybrodt (Editors), National Speleological Society, 100-111. Incropera, F.P. and DeWitt, D.P. 2002. Fundamentals of heat and mass transfer. J. Wiley, New York, 981 pp. Klimchouk, A., Cucchi, F., Calafora, J.M., Aksem, S., Finocchiaro, F. and Forti, P. 1996. Dissolution of gypsum from field observations.International Journal of Speleology 25(3-4) : 37-48. Luikov, A.V. 1968. Analytical Heat Diffusion Theory. Academic Press. Sarbu, S.M. and Lascu, C. 1997. Condensa tion corrosion in Movila cave, Romania. Journal of Cave and Karst Studies 59(3) : 99-102. Simms, M.J. 2003. The origin of enigmatic, tubular, lake -shore karren: A mechanism for rapid dissolution of limestone in carbonate-saturated waters. Physical Geography 23 : 1-20. Tarhule-Lips, R.F.A. and Ford, D.C. 1998. Condensa tion corrosion in caves on Cayman Brac and Isla de Mona. Journal of Cave and Karst Studies 60 : 84-95. Wigley, T.M.L. and Brown, M.C. 1976. The physics of caves. In The Science of Speleology, J.D. Ford and C.H.D. Cullingford (Editors), Academic Press, London, 329-58.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info Underground drainage syst ems and geothermal flux Giovanni Badino Dip. Fisica Generale – Universit di Torino, Associazione La Venta E-mail: badino@to.infn.it Republished from : Acta Carsologica 2005, 34 (2), 277-316 Abstract The paper presents an analysis of the in teraction between the geothermal flux and the water or airdeep drainage networks. The problem of geothermal pow er intercepted by deep structures and, in general, the temperature field calculations, is converted to classical th ermo-engineering problems in terms of shape factors. It is shown that the fluid flow in a conduit perturbs the whole deep rock temperature field until the geothermal flu x of a large area is focalised onto the conduit. It is shown that either small water masses flowing into a mountain are able to perturb the rock temperature up to the surface, on sizes that do not depend on water mass dimensi on, but on its depth, and then on enormous volumes. The introduction of the “geothermal cross section” of an undergr ound drainage structure allows us to improve the classical formula of minimum provenance depth of geothermal water. Enlarging factors ar e applied to the classical estimation in dependence of the ratio between the actual av erage discharge and the critical discharge Qc, which depends on the conduit geothermal cross section. The geothermal “umbra cones” crea ted in the overlying rock by d eep underground structures are described. It is shown that the geothermal flux can play a significant role in the underground drainage phenomenology. Keywords: geothermal flux, heat transfer in karst massifs 1. The geothermal flux and the energy contents of rocks An introduction to the geothermal energy flux It is widely known that the rocks below us have temperatures that increase with depth. The reason is that the internal part of Earth is hot and the surface cold; there are then two “heat sources” (but in this work I have adopted the suggestion of Bohren (Bohren, 1998), avoiding use of the word “heat”), and the thermal energy flows between them with the rules given by the usual conduction equations. Table 1, adapted from (Lee, 1966) gives typical values, widely variable, of geothermal flux, estimated by measures of deep underground temperature gradients. The world average (Verhoogen, 1956) is 2m W 06 0!"gtF The flux is some 60 kW per square kilometre, which corresponds to a total release of 3 # 1013 W on the whole planet. For comparison the energy flux received from the Sun is 1.7 # 1017 W, therefore the geothermal flux is around 5000 times smaller than the main energy source for Earth. It cannot play a role in the free atmosphere phenomenology, but we are going to see that in the case of underground atmospheres it can and it does. At first, this appears not to be true in the case of cave atmospheres, that are really quite cold (essentially at the yearly average external temperature, -from hereafter Tave ; Badino, 2004) while the atmosphere of mines can be very hot (Badino and Forti, 2005). Actually the first aim of this work was to show that the geothermal flux could not play any role in the deep karst microclimates and genesis, because it is shielded by deep drainage conduits: This is exactly the contrary of what we are going to show...

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 TABLE 1 Typical values of geothermal flux Regions Mean geothermal flux Fgt [mW m-2] or [kW km-2] Number of measures Continents Africa 36-61 15 America 25-150 85 Australia 35-160 65 Asia 22-150 60 Europe 26-140 60 Oceans Atlantic 3.4-250 250 Indian 5.9-220 250 Pacific 2.9-340 600 Arctic 33-62 20 Geothermal regions Larderello (I) 450 Oora Kei Korake (NZ) 4200 Matsukawa (J) 630 Mines, caves and tunnels The first point to discuss is the difference between the temperature of artificial and natural cavities in order to clarify the common confusion between “thermal flux” and “temperature”. It is useful to compare two Underground Neutrino Observatories, in Mont Blanc (between France and Italy) and in Gran Sasso (in central Italy), which are assembled in halls in motorway tunnels. The depth of the first, dismissed in 2001, was some 1800 m below the surface, at an altitude of 1300 m asl, whilst the depth of the second is around 1050 m at an altitude of 1000 m asl. Their temperatures are nevertheless completely different, in the first it is around 32 C, in the second at 6 C, the two unsuitable for working. Therefore it is necessary to act in the opposite sense, in the first to cool, in the second to warm the experimental halls. The reason for the two different temperatures is the different rock permeability. The Mont Blanc rock is mainly granite, the waters met by the tunnel were essentially fossil waters, the meteoric water circulation being quite epidermal (up to depth of 100-200 m below the surface), with some exceptions localised along large major rock discontinuities. The Gran Sasso rock is essentially limestone, and drillings have shown a cave layer 550 m above the tunnel altitude and a general water table extended up to that level. The infiltration waters at the surface are essentially at Tave, and cross the whole mountain in nearly adiabatic conditions, which means that they are only very slowly heated along the fall. Therefore in the Mt Blanc’s depth there are essentially “mine” waters in thermal equilibrium with hot, deep rocks, whilst the deep Gran Sasso waters are essentially meteoric waters, in equilibrium with the atmosphere. It is useful to discuss a little more the internal water heating in karst. The reason for water temperature increase during underground fall is the gravitational energy which is converted in thermal (a process that gives a water adiabatic lapse rate –2.34 C km-1) and, in the case of flowing in vadose conditions, also to thermal exchanges with moist air, always characterised by a different adiabatic lapse rate, around -5 C km-1. The actual caves’ lapse rates are between these two extremes (Badino, 2000; Luetscher and Jeannin, 2004), a fact that has huge consequences on the caves energetic balances, which nevertheless are outside this work aim; we return briefly to it in the next chapter, but a discussion can be found in (Badino, 2005). We concentrate here on the fact that the infiltrations

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 create a thermal connection between the atmosphere and the deep karst. In short the thermal contact between deep waters and atmosphere in the Mont Blanc case is due to the recent tunnel, in the Gran Sasso to the natural ancient conduits that have permitted a deep water flow that h as been able to shield the geothermal flux from below and to cool the whole mountain above the water table. More in general we can compare the temperature in the bulk of large mountains looking at the temperatures encountered during the tunnels construction. The worldÂ’s deepest tunnels are essentially in the Alps and it is possible to discuss their internal temperatures. The figures, adapted from (Szechy, 1973) show the rock profile above the tunnels and the corresponding local temperature. The first figure (Fig. 1) shows the situation of the St. Gotthard tunnel, in Switzerland, mainly in granite and gneiss. It is possible to see that the temperature dependence with depth is quite regular. Next figure (Fig. 2) shows the situation in the Simplon tunnel (between Italy and Switzerland), a geological structure in gneiss and, roughly in the Italian part, limestones. It shows low temperature anomalies in the sedimentary part around PK 15 (Luetscher and Jeannin, 2004). The Mt Blanc situation is quite different (Fig. 3). Its rock is mainly protogine, but there is an important fault that allows very deep glacial water circulation, which has lowered the rock temperature near PK 8, where extreme excavation problems where encountered (Guichonnet, 1967). The ge neral behaviour is regular, but the whole mountain has been cooled a little by the fault. The next figure (Fig. 4) shows a completely different situation, the Gran Sasso (Catalano, 1993). Not only the temperature does not increase with depth, but also it shows a tendency to decrease with it, because as deep as it is, as colder are inflowing waters. We have seen above that they meet a very small warming crossing the mountain (Badino, 1995). Other more complex phenomena can occur in determining the rock temperature. For example Szechy cites the case of the Great Appennine tunnel (Italy), mainly in limestone and clay, where a local temperature of 60 C, 500 metres Fig. 1. Depth profile and rock temperatures in the St. Gotthard tunnel. Fig. 2. Depth profile and rock temperatures in the Simplon tunnel. Fig. 3. Depth profile and rock temperatures in the Mont Blanc tunnel. Fig. 4. Depth profile and rock temperatures in the Gran Sasso tunnel.

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 below the surface, has been encountered. The Author says that “discrepancies are due to intensive methane infiltration through the Eocene clay shale”, which looks quite strange. Nevertheless it is quite common to meet high temperatures connected with local thermal water infiltrations, often connected with hyperkarstic phenomena (Badino and Forti, 2005). Here we deal only with a “standard” situation to analyse the geothermal impact on our usual karst. Convective mountains The graphics (Fig. 5) show in short the above discussed data (temperature vs. depth) where each type of data point indicator describes a different tunnel. It is possible to see two completely different behaviours, the “hot” mountains without internal water fluxes, with positive temperature gradient, and the “karstic” mountains with slowly negative or zero gradients. Fig. 5. Rock temperature versus depth in the large alpine tunnels. In fact these mountains are in thermal contact, and in equilibrium, with the local climate, that they follow with a delay that depends on the mountain depth (Badino, 2004). This thermal contact surface-underground due to water vein, is the base for the traditional and fundamental “remote sensing” during tunnel excavations. Continuous rock temperature measurements are performed during work (Guichonnet, 1967); a regular temperature increase with the surface distance is a signal of compact rock. Occasional water veins are in thermal equilibrium with geothermal flux, which is possible only if these water reservoirs are relatively small and without a hydrogeological connection on long ranges, and then cannot be extremely dangerous. Otherwise, if during excavation appears a tendency to a temperature reduction, it is a sure signal of an approaching water stream in direct contact with the surface, that at these depths has obviously enormous pressure, which is able to create extremely dangerous situations, also because it is surely associated to dramatic rock discontinuities. This is the reason why a lot of work is made in the field of rock temperature estimations in deep tunnels (Koenigsberg, 1906; Goy et al., 1996; Badino, unpubl. 2005). Underground high temperatures are connected with good thermal insulations, which means that a tunnel, or a mine, can be excavated across rocks that are very hot (high temperature) because, i) they have almost no contact with the surface and ii) they have then acquired equilibrium with the geothermal flux. Then the relative rock insulation has allowed to a so small thermal flux to heat up to high temperatures enormous quantities of matter. As we have seen above, it is possible to consider mountains with caves as good thermal conductors and then in general (let us forget for a while the geothermalism) they are in thermal equilibrium with the external atmosphere As larger are the caves, as smaller it is the impedance for water crossing the whole mountain, then deep water circulation is possible and the whole structure is crossed, which is a big difference from a mountain without caves where only water circulation is quite epidermal. Another consequence to be pointed out is that if the mountain is highly permeable to water fluxes the thermal energy transfers inside it are absolutely dominated by fluid motion, then the pure conductive terms into the rock (“heat” diffusion) become negligible. It is possible to add th at low-impedance water (or air) transfers inside a mountain can be considered a special case of thermal convective movements, and then the karstic mountains are examples of convective thermal contact with the atmosphere, whilst the impermeable rocks have only a conductive contact with it, with a very poor efficiency. The hot mountains energy contents Let us estimate the energy contents of a hot rock prism, with a surface A and an altitude of

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 H. Its temperature at the beginning is T0=Tave, local average temperature. The geothermal flux will heat the rock until the temperature gradient in it becomes able to evacuate the geothermal power on the surface. This equilibrium condition (“stationarity”, i.e. no time dependence) is obtained when the gradient is z T K FR gt$ $ Where we have used partial derivative because in general the temperature is function also of time. This is exactly the condition that allows the measure of geothermal flux from temperature gradient data, assuming a steadystate condition. In this case we can assume a rock thermal conductivity KR=2.5 Wm-1 K-1, typical for granite, whilst the limestone conductivity is some 10% less. Then the temperature gradient at the equilibrium is 1m K 024 0 5 2 06 0!" " $ $R gtK F z T Which is 24 C per kilometre. It is easy to calculate the energy necessary to heat at this constant temperature gradient a rock prism of surface A from depth Z up to the surface; its total available energy is Q ACR%R& T0 Z'dz ACR%RFgtKR Z22 Where CR is the rock specific thermal capacity (800 J kg-1K-1) and %R its density (roughly 2600 kg m-3). For example, assuming Z=1 km, with typical values the total energy contents per kilometre square of surface is J 10 5 2 2 1000 024 0 2600 800 10 216 2 6 2# # # # # " Z K F C A QR gt R R% For comparison, a 20 kTon atomic bomb yields 1014 J. It is important to note that this energy, that a power plant of 1 GW produces in two years, has been released to our prism rock by the small geothermal heater, only 60 kW. It is a small power, but it has worked from a very long time, the rock is a very effici ent thermal insulator and the final result is an efficient energy storage. Rock heating time To estimate correctly the time needed for heating, it is necessary to take into account that the rock prism is in contact with others all around; the calculation would have to include these in the estimation, reconstructing the whole temperature field and its dependence on time. The problem is complex but we can do a crude estimation of the prism heating time scale considering it as thermally isolated from its surroundings (which is equivalent to assume a flat surface and uniform, uni-dimensional heating), then 2 1 22 2Z K C AF Z K F C A AF Q tR R R gt R gt R R gt heat% %" ( ( ) + + " & The last term can be rewritten in terms of thermal diffusivity coefficient defined as R R R RC K a%" In the case of rock 1 2 6s m 10 2 1! !# .Ra And then R heata Z t22" & Typical width of limestone mountains are around 1000 m, then the heating time scale is around 104 years, not so much for geological time scale, but longer than the typical global climate fluctuation (Badino, 2004). Penetration lengths of temperature fluctuations Let us recall the classical thermal fields solution of a homogeneous thermal conductor to a sinusoidal and to a sudden (step) temperature change. In the first case a thermal wave propagates inside, fading exponentially (Badino, 2004;

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 Lismonde, 2002). If tsin is the period of sinusoidal wave we have that the “penetration length” is sin 2t a lR p/" If the temperature fluctuation has a step shape, the propagation length is more difficult to define. The penetration of thermal shock of amplitude T is described in terms of rock temperature increase at depth x at time t, by equation (Isachenko, 1969) Tx t01"& T 1-erf x 2 aRt -2 ,2 +2 +2 *2 )2 (2 (2 3 2 42 52 52 62 7 2 8 2 8 2 Where erf(u) is the Error Function. The figure (Fig. 6) shows the results at different times. A discussion about the properties of this solution can be found in (Lismonde, 2002). Fig. 6. Diffusion in rock of one step temperature increase on its surface. The last equation suggests that a step thermal wave is able to reach a depth x in a time t or, vice versa, after a time t the thermal disturbance has reached the depth x, and the relation between the two quantities is given by the argument of Error Function. More precisely it is possible to show (for example, classically (Laidler and Meiser, 1995) that x2. 2 aRt Then the penetration in rock up to depth Z of a cool wave requires a time-scale & tcool & tcool" Z22 aR The estimation is heavily approximated, because we are not looking for the complete cooling of the mountain, but for the equilibrium temperature field formation inside it, what is attained not when the whol e rock is at the same temperature of the surface, but when its temperature has attained the “stationary temperature field” geothermal gradient seen in the previous chapter. Nevertheless it makes no sense to try to perform exact calculations, which in any case work with unrealistic forms of mountain. We can conclude that the heating and the cooling time-scale up to the equilibrium configuration are almost the same, and they do not depend on the temperature drop, then at depth Z in metres we have & teq" Z22 aR .& tcool.& theat [1.1] In the case of compact rock we can assume that the equilibration time scale in years is given by & teq 0.01 Z2 [y]

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 This estimation is very important for our discussion. In the next chapters we are going to consider water and geothermal fluxes that have attained stationary conditions. It is then obvious that if the water flow has begun from a time that is much smaller than that of equilibration time, it is not possible to assume that the system has attained a general equilibrium configuration, what in fact happens in the artificial excavations. This is the reason why the mines are hot. 2. The basic approach to the problem The problem of infiltration temperature The classical Desio’s formula (Celico, 1986) that gives the minimum depth attained by geothermal water is based on the assumption that underground there exists a first layer ( heterotermic ) in thermal contact with the atmosphere with essentially the temperature of local inflowing fluids, which have a temperature T0 quite precisely (but in general, a little lower, (Badino, 2004)) equal to the local yearly temperature Tave. Below this first layer the rock temperature increases in conductive regime ( homotermic region ). The heterotermic layer is also called “active layer” (US Bureau of Mines, 1996), mainly because the seasonal variations can create icing. It is necessary to note, nevertheless, that some problems do exist in the definition of homotermic layer (Schoeller, 1962; Celico, 1986), because it is considered the layer where temperature does not depend on atmospheric temperature variations, not the layer where temperature is equal to the local average of the atmosphere. It is then defined on the basis of its temperature stability, not on the basis of its thermal contact, if with atmosphere or with deep rocks. Really in rock with deep aquifers we meet a heterotermic layer with seasonal fluctuations (some score of metres), a layer above the aquifer (included) at T0 (Luetscher, 2004; Badino, 2005), a thin layer of thermal contact deep rocks-aquifer, where a relatively sudden temperature increase, probably dependent of aquifer permeability, is possible (Goy, 1996; Badino, unpubl. 2005). We would call it the “geothermal exchange layer”. Below this region, a regular temperature increase in deep rocks is found. The scenario is then more complex, but it is better to postpone a detailed discussion to a future work. We can spend some words about the exact value of T0 and its relation with Tave. Really Tave depends on the altitude, and on average decreases of 6-6.5C per kilometre, as described with the International Standard Atmosphere. Also at a first approach it looks better to think that the rock assumes not the average yearly temperature of the atmosphere Tave, but the average temperature of wa ters at the infiltration point, that is quite lower because the rain waters in alpine karst are generally associated with colder periods (but in tropics with warmer periods); really, many other corrections are necessary to estimate the local rock temperatures (Badino, 2004). Another corrective term, already cited above, appears during the underground flow because the temperature increase of underground waters along their travel is different from outside, where the waters follow the ISA mean lapse rate (6 C per kilometre of fall). Underground, in adiabatic conditions the theoretical value of water temperature increase is 2.34C per kilometre of fall, but th e experimental values into the caves, where en ergy exchanges with the air are possible, are around -3 and -4C km-1. Really, we can assume that the water temperature that has infiltrated at altitude H [km], when it arrives at level 0 is some (3-4) # H [C] hotter, then sensi bly colder than the corresponding infiltrating waters at that altitude. In detailed calculations these effects, that create a difference between the actual cave temperature T0 and the local yearly temperature at its same altitude Tave, have to be taken into account, but in our discussion they are completely negligible. The energy release to groundwater There is an obvious approach to consider the role of geothermal flux. In the upper Earth surface layers the geothermal energy is essentially intercepted by water that releases it to the atmosphere when it goes out from springs. The energy that comes onto a large surface A is obviously FgtA, and it is very regular in time. Let us suppose that it is absorbed by a mass M: Its temperature increase rate is then given by the condition

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 dT M C Adt Fw gt" Where Cw is the water thermal capacity, that is M C A F dt dTw gt" Then if we know the shielding mass M we can calculate its temperature variation with time, but how large is M? Reasonably it is the mass of groundwater, but it changes seasonally, depends on rock characteristics and so on. Still worse, the previous equation says that the temperature continues to increase because the mass is stable and the energy flux continues. The equation gives us a temperature-changing rate, not a temperature variation. It is easy to avoid these problems. Really the mass M can change, but its water is almost continuously flowing, which means that some water enters cold in the mass M, and on average the same quantity flows away from some other side, warmer. We will impose a condition of steady state and calculate the temperature variation of the flowing water, independently on M and, above all, on time. The geothermal energy flux is very regular but the groundwater flow in general it is not. We then may expect that the temperature change fluctuates, also if we consider the oversimplified system of a single water mass M and we neglect other problems like the drainage network structure and mixtures between different branches that depend on the water flow rate. So, the temperature changes fluctuate, but we are going to estimate the average value of temperature increase due to geothermal flux. The groundwater temperature increase It is easy to estimate the flowing water temperature increase, assuming the powerful and reasonable hypothesis that on average the system is stationary. This means that the thermal energy inflowing in M from Earth is, on average, going out as water flux enthalpy increase. In this way we have not to consider the mass M, that we cannot know, but only the outgoing flux from it, because our steady state assumption states also that the temperature of M does not change with time (on average...). Let us estimate the water flux out-flowing from a region of surface A, which thermally interact underground. If the precipitation is P (in kg m-2s-1), the infiltration is P minus the part Pout “lost” outside due to evaporation. This part depends on climate, surface type and so on; in temperate regions ranges between 30 and 40% of total, but in deserts can rise to 90% (Celico, 1986). With this assumption the enthalpy extracted from the system is (P-Pout) & TgtA, where Tgt is the water temperature increase during deep flow (Fig. 7). The condition of stationarity implies that on average the temperature cannot change in time and then FgtA P Pout 0 1 Cw & TgtA That solves our problem. We can change units, calling P* the infiltration in [mm a-1], to obtain [2.1] Fig. 7. Interception of geothermal energy flux by a flat aquifer. In stationary approximation the water is heated and the upper rocks are completely shielded. The upper parts of drainage systems (for instance the caves, highly permeable) are almost exactly at the exte rnal average local temperature, therefore the water is in general warmed of & Tgt between the lowest cave parts and the springs, that is along the flow in the phreatic systems. 01 9:C 500 10 2 4 06 03; # & P P P Tout gt

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 In alpine karsts P* is some 1000 mm a-1 and then the water average temperature increase due to geothermal energy is some 0.5C. It is a small term, very difficult to measure, and really it has never been measured. But it is not always so small, a temperature increase of 5C between the caves and the springs has been measured in Cuatro Cinegas, a figure that corresponds to an infiltration (P*-Pout)=100 mm a-1 in this desert region of Coahuila, Mexico (AA. VV., 2004). Therefore it would theoretically be possible to estimate the average infiltrating flow measuring this temperature increase, but it is a measure quite difficult to do with some accuracy, because it is the difference of two uncertain parameters, and other processes surely occur along the water rock crossing. The plane watertable We have used these calculations in a previous work (Badino, 1995) to explain why the karstic mountains are so cold, therefore excluding a significant role of geothermal flux in underground climate definition, because the energy flow from Earth depths is easily shielded. This very natural approach to refuse, in general, any role to the geothermal flux, it was not original, it is a quite traditional point of view. Bgli (Bogli, 1980) estimates reasonably 0.2 C “to prevent the karstified zone above from being geo-thermally heated up”, that is a very interesting idea that we shall meet again. In a very interesting and co mplete paper Mathey (Mathey, 1974) estimates a maximum of 0.55 C. More recently Jeannin et al. (1997) estimate the specific discharge of a karstic spring (the equivalent of infiltra tion) between 30 and 3000 mm a-1. The first figure seems too low (almost three times dryer than the North Mexico deserts), but in this paper it is used an energy flux too low by a factor thousand, to finish to say that the temperature increase, that has physical dimension [C], is “less than 0.1 C/a”, where really the 30 mm a-1 case would be warmed by some 15 C. Let us return to this estimation. We have obtained the average water temperature increase and we could now begin to calculate the fluctuations dependence on discharge and so on. But we are dealing with conduits, not with plane watertables. Are these calculations and assumptions true for similar “discrete” systems? The general answer is that no, they are not generally correct. The water flows along definite branches, that do not cover a large surface and with a general shape that is far from regular. The complete shielding assumption is not reasonable for karstic drainage. 3. The underground temperature field with a drainage network The problem The scenario described above (a regular, flat, diffuse water table) can sometimes be correct but in general it gives completely wrong results in the internal rock temperature field estimations. To study a more real model it is necessary to estimate the energy interception made by a system (a thermodynamical sink) that is buried in a semi-infinite medium where a thermal energy flux is flowing from infinite. Let us consider the problem details. We have a semi-infinite rock volume in which a thermal flux Fgt is coming from below. It creates a temperature gradient given by: $ T$z F gtKr Where we have assumed the depth z as positive downwards. The temperature field at depth H below the surface (or, better, below the heterotermic layer) it is therefore given by TH01" T0<$T$z -2 ,2 +2 *2 )2 (2 H T0< FgtKr 2 ,2 +2 2 )2 (2 H We have previously seen that T0 is essentially the local Tave. It is very natural to suppose that the geothermal energy intercepted by a deep structure is that given in previous chapters, that is (geothermal flux) # (structure area), therefore [3.1] This means that the energy intercepted by a “cave” is proportional to its surface in the direction of energy flux. It is a very natural assumption, but it is false. gtF A W

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.10 The fact that, up to now, has not been considered is this: If we bury a system able to intercept and to evacuate elsewhere the energy, the whole temperature field in the rock is altered and then the flux itself changes The geothermal field with a cave The problem of temperature field calculation in this configuration has to be solved with these boundary conditions: 1) The geothermal flux from infinite is constant, Fgt; 2) The temperature on the surface is T0 constant everywhere; 3) The cave temperature is T0, the same as on the surface (we have sufficiently discussed the limits of this assumption); 4) At the infinite the temperature field is not disturbed by the cave existence. These conditions imply a flat external topography and assume that the infiltrated water heating from surface to the cave is negligible. We said that if the cave has an area A, it is natural to assume that th e absorbed geothermal energy is Fgt# A [Eq 3.1]. Is this correct? Let us firstly discuss it qualitatively, drawing the isothermal surfaces in th e rock. The figure (Fig. 8) shows a reasonable situation that respects the boundary conditions. It is possible to see two things: 1) The isothermal surfaces have a tendency to converge, then to be focussed, onto the cave; 2) They are “compressed” around the cave. If we remember that the thermal flux flows along the maximum T variation (i.e. along the grad (T), which means perpendicularly to the isothermal surfaces) and that its value is proportional to the gradient of T, we have that the two features are equivalent to say that: i) the cave focus on itself the geothermal flux and, ii) in the rock surrounding the cave the geothermal flux (and the geothermal gradient) is much more intense than the natural one... So, the assumption that gave us the [Eq. 3.1] is surely wrong. But is it possible to calculate the correct value? Quite surprisi ngly it seems that nobody has studied this important problem. Before we look for the solution, we have to make some note about the enormous weight of the stationarity assumption. Whatever initial temperature field condition will converge to asymptotic values which are solution of Laplace equation, but this convergence requires time. During this time, which is of the order of equilibration time scale introduced above [Eq. 1.1], the difference between the real field (in transient phase) and the asympotic one (stationary phase) can be important. If the equilibration time is comparable with the typical changes of boundary conditions (global climatic changes, infiltration of hot waters or so on), the system can never be considered in a true stationary phase, and the equation given by this assumption has to be considered heavily approximated. This limit of stationary assumption gives strong uncertainty in the temperature fields estimations for the new deep alpine tunnels, which have very long equilibration times (Badino, unpubl. 2005), and affects also our next considerations. Fig. 8. Qualitative view of the stationary thermal field due to the interaction between the geothermal flux and a conduit with strong drainage (system S). The isothermal surfaces are affected in a wide region, and the cave temperature is T0. The general solution Therefore the problem of calculation of stationary temperature fields is very difficult to solve just with the easy boundary conditions given in the previous chapter. What to expect, then, when we will have to assume finite energy transfer rate inside the cave or situations in which the temperature of the cave itself is

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.11 determined by the geothermal flux? And what monstrous form may assume the solution in a non-stationary case, if we want to consider, for instance, the cooling of a mountain during the karst creation inside it? Nevertheless there is a simple way to lead this problem to typical situations of engineering thermal exchange. Let us show how, considering three different systems, S, S’ and S”: 1) The main system just now described, the real case, which we call from hereafter S (Fig. 8); 2) The system composed only of semi-infinite undisturbed rock, without caves and with external temperature equal to 0, which we call from hereafter S’ (Fig. 9); 3) The more complex system S”, (Fig. 10), composed by a cave at a particular temperature T” buried in a semi-infinite rock, that releases energy to the surface at temperature T0. Also the rock at the infinite is assumed to have temperature T0. In this last ideal system S” there is no geothermal flux at all. We are then ready to do the final step, assuming that T” in S” has a value given by: [3.2] where H is the cave depth from the surface. We then assume that the cave in S” has exactly the temperature of the rock at depth H in the system S’ plus T0. Let us consider now the three temperature fields. They are the solution of general Fourier equation (Isachenko, 1969) 01 t T a z y x T $ $ = > =1 , That in our case, not time dependent, it reduces to the Laplace equation 02" = T It means that the T fields behave like a huge class of phenomena for which the sum of the three spatial second derivatives is zero. The functions that satisfy these conditions are called “Harmonic Functions”, and are among the most important and studied functions in Physics (Carslaw and Jaeger, 1959; Bejan, 1993; Balcerzak and Raynor, 1961; Nashchokin, 1979). By the way, it would be possible to solve our fields using solutions given for different and well-studied problems like, for instance, the electric field due to particular charge distributions, but we can do better for our purpose. Fig. 9. Qualitative view of the stationary thermal field due to the interaction between the geothermal flux and a conduit with no drainage (system S’). The isothermal surfaces are unaffected, and the cave temperature is higher than T0. Fig. 10. Qualitative view of the stationary thermal field between a “hot condu it” at temperature T” and the surface at T0 (system S”). The isothermal surfaces are finite, closed and contain the conduit. Heat diffusion problems like this one are quite usual in thermal exchanges engineering. H K F T Tr gt< "0"

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.12 The field structure around S, that we have qualitatively shown (Fig. 8), is our unknown term. The field structure of the second S’ it is obvious, it is composed by many horizontal parallel lines (or planes, in 3-d) for T-field and vertical vectors for thermal fluxes. The third system is the most exciting. There are almost no lines all around (neither for T nor for thermal fluxes) unless in the region between the cave and the surface, because the temperature difference between the “hot cave” and the surface and also the “infinite”, drives an energy flux. From the other side, this situation is a very usual situation for engineering, because the “hot cave” can be a tube transporting hot fluid buried in some engine... Now we can do the last step: We state that the unknown T field of the system S is given by 010101z y x T z y x T z y x T , , , [3.3] That is, we can subtract the (very complex, but very common) T” field from the trivial T’ to obtain our solution T. It is possible to prove this theorem in three steps: 1) The Laplace equation is linear, then if T1 and T2 are solutions, also (T1-T2) is a solution: here in particular T’ a nd T” are solutions, then also T is; 2) The boundary conditions written above are satisfied by a T field given by (Eq 3.3); 3) Then T is a solution of our equation with these particular boundary conditions, but the solution is unique for the Uniqueness Theorem for Harmonic Functions then, T is the solution... The base of this proof is the linearity of the grad operator, which allows the first property. But also the temporal derivative is linear, and then we have another independent fundamental result: The T field may be calculated by this way also for transient conditions, if we use the equivalent transient solutions for the T”. We are not ready to use here this corollary, probably very important. The geothermal energy focusing on caves It is interesting to look for other consequences. Let us return to our [Eq. 3.3] to apply the grad operator (that in equation will be noted with = ) and multiplying for the rock conductivity KR K R = T K R = T K R = T These terms are now the energy fluxes that flow through the systems S, S’ and S”, and then 0 1 0 101z y x F z y x F z y x F, , , ! ! But the energy flux in the system S’ is simply -Fgtk where k is the unit vector in the zdirection, and then 0 1 0 1z y x F k F z y x Fgt, , ! ! This means that we are able to calculate the energy fluxes in the system S with vectorial subtractions between the S” system, complex but already studied, and the trivial S’. If we multiply this equation by the surface element dS and integrate on a wide surface A that contains all the surroundings of the cave we have A F A F FAgt" ! The term F # A describes the flux outgoing from the surface in presence of the cave, Fgt# A the total flux if it would not be the cave, than the energy flux captured by the cave is the difference between the two [3.4] In this way the problem of energy interception of a cold cave buried in an energy flux is reduced to the energy transfer between a hot cave and the surface. Now it is possible to study the T” solutions, going to the heat transfer engineering to use its results. The shape factor The thermal transfer engineering uses a very effective approach to the problem of complexshape systems. Let us return back to the fundamental equation of conductivity, now written in three dimensions. The thermal flux through a surface element dA is given by dW K R = T dA Where KR is the body (rock) conductivity. We may consider two sources at definite temperature T1 and T2. The thermal energy is drained between the two by an intermediate A F A F A F A F FA A F A F Wgt gt gt capt capt" " < " "

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.13 temperature field confi guration, which depends on the system shape in an extremely complex manner. Let us call Aiso the isothermal surfaces that we can draw between the two sources; these surfaces can be infinite, also if neither of the two bodies is infinite. The grad (T) must be perpendicular to these surfaces because the thermal flux vector cannot have any component along an isothermal surface, and then the thermal energy flows normally to these surfaces. Let us call n the coordinate along the thermal flow path. The previous equation gives the thermal flux as dW "! KR $ T $ n dA The total flux between the two sources is then given by the integral of thermal flux on any of these surfaces. It is not important to choose one or the other, because we have assumed that the only two “heat producer or destroyer” are these two sources, and the energy must be conserved. The thermal transfer through one of these surfaces is therefore W "! KR $ T $ n dAAiso' We define now a new dimensionless temperature T* (that is in fact a relative temperature variation in the path between the two sources) as 1 2 1* T T T T T ! Where T1 and T2 are the sources temperatures. Then it is possible to write W "! KR T2! T101 $ T *$ n dAAiso' In this way the sources temperatures are analytically separated from the system geometry, which now is completely included in the last integral, which is nevertheless extremely complex also for trivial configuration. This equation has to be compared with the usual equation which describes the thermal energy transfer between two sources separated by a uniform distance & z through an area A W "! KR A T2 T1& z 2 ,2 +2 *2 )2 (2 We see that the integral behaves like the ratio between A and the sources distance, and then the system geometry is included in this term A & z 2 ,2 +2 2 )2 (2 $ T $ n dAAiso' The term in brackets is the ratio between the surface crossed by the thermal energy and the distance between the two sources. It is a “length” that characterises each system shape that exchanges energy among two sources. This scale-length is called “shape factor” in literature (Carslaw, 1959; Hahne, 1975; Holman, 1996; Ozisik and Necati, 1993; Kays, 1966). We adopt unwillingly the usual notation, that uses the “S” for a length, but we shall write SF, hoping to reduce (perhaps…) confusion with the subscript “F”. Then [3.5] The geothermal power absorbed by the cave can be then written as [3.6] Now we have to study the way to use this result. The “shape factor” calculation We do not have to study in details the way to calculate the shape factors. Still in simple configuration the isothermal surface calculation is very complex and the in tegration is in general extremely difficult, but the heat-exchanges literature contains many shape factors worked out for the most common geometrical configurations. Most of these results have been obtained based on advanced analytical methods (conformal mapping, superposition, special transforms, analogies with the electrical potential studies and so on); it is not useful for us here to study these approaches. Table 2 shows these shape factors in interesting situation (Holman, 1996) We are going to use these shape factors, but first it is necessary to answer an important question: Are the shape factors linear? Let us discuss the question with an example. SF" $ T *$ n dAAiso'T S K WF R &

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.14 TABLE 2 Semi-infinite medium with isothermal surface and isothermal spherical cavity of radius R at depth H H R R SF2 / 1 4 / H R T2 T1 Semi-infinite medium with isothermal surface and isothermal disc of radius R parallel to the surface at depth H 01 H R R R SF2 / tan 2 / 41!! "/ / H R T2 T1 Semi-infinite medium with isothermal surface and isothermal cylindrical cavity of length L of radius R, parallel to the surface at depth H 01 R H L SF/ cosh 21!" / H T2 T1 R L Semi-infinite medium with isothermal surface and isothermal cylindrical hole of radius R drilled to a depth H normal to the surface. 01 R H H SF/ 2 ln 2 / H T2 T1 R Semi-infinite medium with isothermal surface and isothermal plate (width W, long L, H>>W) parallel to the surface at depth H 01 W L W SF/ 4 ln 2 / H T2 T1 L W If we know the shape factors of two independent systems S1 and S2, say, one composed by two cylinders and the other of a sphere and a cylinder, can we consider a third system S3 (in this case, two cylinders and a sphere) as composed by some “sum” of the two firsts, and consider that its shape factor is given by some “sum” of the two? Unfortunately the answer is: No, we cannot. The temperature fields are linear and then the S3 temperature field can really be calculated from the S1 and S2, but it changes completely the equipotential surfaces on which the integration is performed to “average” the flux in the integral [Eq. 3.5]. It is therefore necessary to recalculate these surfaces and to repeat the integration that will give a result that has no direct connection with the integrations of S1 and S2 fields. This means that, for instance, the knowledge of the shape factor of a conduit buried in a semiinfinite medium tells us almost nothing about the shape factor of two parallel conduits in the same medium, unless their distance would be so large that each temperature field is not affected by the other. Only in this latter case the shape factor of the two conduits is the sum of the shape factor of two single conduits, but in general it is not so. 4. The interception of geothermal flux by caves The geothermal cross sections of caves It is now possible to apply the previous results to the problem of interaction between caves and geothermal flux. It is stated above that, to satisfy the boundary conditions, the T” temperature has to be H K F T Tr gt< "0" As a consequence, the equation that gives the power intercepted T S K WF r &

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.15 Fig. 11. A deep draining conduit can create a very large geothermal shadow on the surface. Its scale size is not the conduit size, but the conduit depth It is reduced to the very simple result W KR SFFgtKr H FgtSFH [4.1] That is 010101depth cave factor shape cave flux geothermal flux d Intercepte # # " Table 3 shows that the scale-size of shape factor is some 1-10 times larger that the scalesize of the underground structure. The last equation has to be compared with [Eq. 3.1]. Therefore, the effective area (we call it “thermal cross section”) for thermal flux absorption of an underground conduit (Fig. 11) is not its geometrical area but instead SFH, roughly 10 times the cave size multiplied by its depth then it is always enormously greater than the cave’s actual area! This amplification is due to the “converging lens” effect created by the presence of cold fluids in the cave that affects the whole structure of the rock temperature field. For example, let us estimate the geothermal power intercepted by a conduit at a depth H=500 m with radius r=0.5 m and long L=300 m. We may use the shape factor given by m 250 6 7 1900 cosh 21" ( ) + "!r H L SF / With Fgt=2 # 106 J m2a-1 we have W FgtSFH 2 # 106250 # 500 0 1" 2.5 # 1011Ja 1 Which is really a big power. It is possible to study how much this figure changes with conduit radius. Table 3 shows the energy collected, by conduits of different sizes. The result in the third column at first appears surprising, because it shows that the variation of the conduit size does not affect so much the intercepted power, but it is reasonable because the power is not absorbed by the conduit surface, but by the focusing effect of conduit on the temperature field. The fourth column shows the surprisingly high average thermal flux (note that are Watt per square metre!) that enters through the conduit transverse surface. For comparison, the Sun deposits on average 1.4 kWm-2 on the Earth surface: The geothermal energy deposition on small conduits is then of the same order! This appears to be absolutely unbelievable, but is it true? Roughly, the answer is that: Yes, it is true. But there are other important details to be taken into account. The heating of water in deep conduits Let us discuss the effective water heating in the focusing conduit, calling T its temperature at the springs. In the previous discussion, it was made the fundamental assumption that the water temperature T0 in the conduit does not change and really it is its low temperature that changes the whole temperature field of surrounding rock. This is equivalent to assume that the water flux (or air flux in case of dry caves) is so large that the enthalpy intercepted by the conduit flows away in the form of a small temperature increase of a very large amount of fluid, and does not really affect the conduit temperature.

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.16 TABLE 3 Conduit radius r [m] Shape factor SF [m] Intercepted power [J a-1] Average Flux on conduit [W m-2] 0.01 165 1.7 # 1011 900 0.1 210 2.1 # 1011 110 0.5 250 2.5 # 1011 25 2 310 3.1 # 1011 8 5 360 3.6 # 1011 3.8 We have then that the maximum energy extraction efficiency is obtained if the exit temperature T is almost equal to T0. We have the opposite limit, if the warming is so large that the water temperature T becomes T”, the temperature of undisturbed rock. In this case the conduit becomes completely “transparent” to the geothermal flux, the temperature field assumes a regular geothermal gradient not affected by the cave presence and, as a consequence, just a little energy is intercepted. The water into the conduit is then a hot “mine water” and the classical, “wrong” solution [Eq. 3.1] becomes correct. The real cases are intermediate between these two extremes, because in first approximation the water really warms but, as a consequence, its capability to intercept geothermal energy is reduced, because the temperature difference between the rock and the water becomes smaller. Really the scenario is still more complex, because the water temperature increases along its path underground: It enters cool, very efficient in geothermal energy focussing, but as long as it warms downstream its capability to intercept the geothermal flux decrease. The non-linearity of SF forbids correct analytical solutions, but we can make some other step. The critical shielding discharge It is possible to estimate the water heating along a deep conduit and its final temperature as a function of conduit parameters. We have just seen that two extreme scenarios are possible. If the fluid flux is very large the rock temperature field is completely changed, the geothermal energy flow interception is maximum and the water flows out at T0. At the other extreme, if the water flow is very small, the rock temperature field is completely undisturbed, the geothermal energy flow to the water is minimum and the water flows out quite hot, at T”. Let us define the “critical fluid flux” Qc that divides these two scenarios in a usually idealised way. We look for a water flux Qc that enters at temperature T0 and flows out at T” in stationary conditions. The enthalpy subtracted to the system is dE CwQcdtT T0 0 1 If we admit that the system parameters do not depend on time, this enthalpy deficit must be given by the incoming geothermal flux Wdt. Then using [Eq. 4.1] we have FgtSFHdt CwQcdtT"! T0 0 1 But T” is given by H K F T TR gt< "0" And we obtain [4.2] Then the critical flux is simply the conduit shape factor “scaled” by a dimensional term (rock conductivity divided by the thermal capacity of flowing fluid). In terms of volume flux dV dt Qc%w K R S FCw%w m3s!19: From another point of view, Qc can be considered the critical flux below which it is possible to consider that the rock temperature field is undisturbed. Or, from still another point of view, we can be sure that a water flux much larger than Qc perturbs the rock. Qc" dM dt K R S FCw kgs!19:

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.17 The solution [Eq. 4.2], which is surprisingly simple in comparison with the bloody analytical difficulties of the problem, is able to give also the velocity that water must have to satisfy the equation. If the conduit radius is r, then w w F R w cr C S K r Q v% / % /2 21 1 " In the case of a water draining conduit in limestone we have Qc water" 2.3 4.2 # 103 SF" 5.5 # 10! 4SFkgs! 19: [4.3] If the conduit drains air, the critical flux becomes Qc air" 2.3 103 SF" 2.3 # 10! 4SFkgs! 19: The shape factor is in general comparable with the conduit length, usually some 103-104 m, therefore the water flux able to create changes in the rock temperature fields is in general quite small. Neither the air flux requested to cool the rock it is too large, in absolute, but its small thermal capacity and its very small density cause a large volume flow request. It is nevertheless necessary to remember that temperature field changes are only possible if the fluid fluxes have had sufficient time (more than & teq defined in [Eq. 1.1]) to converge at the stationary (equilibrium) state. The critical flux Qc for air or water are extremely important for another reason: They are the air or water fluxes able to shadow the upper rock from the geothermal flux forcing its temperature near to the average yearly temperature of external atmosphere, Tave. Finally, it is important to note that this work was originally a chapter of an underground climate physics book. Its purpose was to estimate when we could consider a cave as “shielded” from the upward geothermal energy flux. If the flux is much larger than Qc we are sure that the mountain parts (and caves) above the conduit are shielded; if the flux is well below the critical value the conduit is “transparent” to the geothermal energy and we have to include also its contribution to analyse the underground climate in the rock above the conduit. We are going to improve this point of view and discover that Qc has another, still more important, meaning. Geothermal power intercepted We can calculate the effective cave temperature T at the equilibrium and solve the inverse problem, the estimation of flowing depth of hot spring waters. The final system temperature must lie between T0 (near it for high water discharges) and T” (near it for low discharges). Let us call WM the maximum power that it can be subtracted by our system WM FgtSFH 0.06 SFH W9: [4.4] Which is valuable for very large water flows and outflowing temperature around T0. If we call T the real (unknown) conduit temperature at the equilibrium, the residual outgoing upward flux is not zero, because it “filters” an energy given by 0 10T T S K WF R up The difference between WM and Wup is the net flux entering inside the cave from below. If the conduit is at temperature T, the energy conservation states 010T T S K H S F W W WF R F gt up M in! " But from [Eq. 3.2] we have 0" T T H F Kgt R! Then Win" FgtSFH 1 T T0T T0 2 ,2 +2 2 )2 (2 WMT T T T0 2 ,2 +2 *2 )2 (2 [4.5] If T=T0 the equation reduces to Win=WM (which describes the “system S” situation), and if T=T” the term Win vanishes, as expected. Therefore, the geothermal flux intercepted by the cave is reduced as long as its final temperature T increases: The cave is becoming “transparent”, and this equation describes its “fading” inside the temperature field.

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.18 Temperature and deposited power versus discharge Usually we have very small possibilities to establish a natural conduit discharge. We deal with actual discharge Q, and we can only try to understand how this water (or air) flow is affected by the geothermal flux. Let us then return to the application of the First Principle [Eq. 4.5]; if we call Q the total actual water flux in the system, it gives WMT T T T0 2 ,2 +2 2 )2 (2 CwQT T001 [4.6] Where we have assumed that the whole entering energy flux goes to warm the water flux. This statement is true if we admit that the thermal energy is released on the whole system i.e. it is true if the temperature T, which is actually the output water temperature, can be used to describe the whole system, also in its further parts. It is a strong, and in general false, hypothesis, but it is better to assume it as true and only afterwards have a look on what happens in more real situations. With the previous results and trivial calculations, we have 01 01 ( ( ) + + ! ( ( ) + + ! " ( ( ) + + ! "0 0 0 0 0" " " T T T T Q T T T T T T C W T T T T T T C W Qc w M w M [4.7] This important equation relates the actual flux Q to the output temperature T in terms of the critical shielding flux Qc and the surrounding temperatures. It is very easy to solve it to obtain [4.8] Where we have called [4.9] The excess temperature above the “external average” T0 is then T-T0, but its natural scale is the ratio between this difference and the theoretical, maximum difference T”-T0. Then T T0" T < qT01 < q T0" T T0011 1 < q 2 ,2 +2 *2 ) 2 ( 2 And calling the “excess temperature ratio” of groundwater, that is the amount of actual heating in comparison with the maximum attainable, we have [4.10] With the assumption [Eq. 4.9] we can rewrite the [Eq. 4.5] [4.11] We have then two funda mental equations, [Eq. 4.10] and [Eq. 4.11], which connect the internal drainage Q to the outgoing temperature and to the intercepted geothermal flux. The two graphics (Fig. 12) describe the behaviour of the out-flowing water temperature T and of water absorbed energy as a function of discharge, obviously measured in function of our nice scale-discharge Qc (it can be adapted to air flow with trivial changes). We have previously discussed the Qc as the “shielding flux” and WM as the “maximal intercepted flux”. Now we see that they are mainly the natural scales of fluid flow and of geothermal power flux intercepted, exactly as happens with the Similarity Numbers, always a ratio between a parameter and a scale-value for it. Fig. 12. Water temperature increase and intercepted power by a deep conduit versus water discharge. rT" T T 0T T0 1 1 < q M M inW q q T T T T W W < ( ( ) + + ! 1 "0 q qT T Q Q T Q Q T Tc c< < < < 1 1 "0 0 cQ Q q "

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.19 Really, the apparition of a natural scale-value of the main variables, the fluid flow and the geothermal power, may suggest to use this typical thermoor fluid-physicists approach, calling the Q/Qc and W/WM terms as Someone Number, to create two underground companions to Nusselt, Reynolds, Peclet, Froude and so on Numbers. Nevertheless we do not like this way to describe physical processes, preferring (aesthetically, because mathematically it is absolutely equivalent…) to work with the scale dimensions. Therefore we do not propose this Similarity approach although, if a name must be given to the geothermal “heat” term W/WM, we strongly suggest “Alighieri Number”... The inverse problem It is obviously of main interest the problem of deducing the provenance depth of a hot spring. The traditional Desio formula (Celico, 1986) assumes essentially that a water flux at temperature Tw comes from a depth H0 at which the temperature of undisturbed rock is Tw, that is [4.12] This is true for “mine waters”, deposits of resident waters, which have essentially the T” temperature, but we have seen that this assumption is in general false, because a natural water flux had surely succeeded to disturb the whole rock temperature field. We have then to use [Eq. 4.8] and [Eq. 4.2] to write 01 1 4 1 0 010 5 5 1 1 "! !( ( ) + + # < " ( ( ) + + < !F R gt c wS Q H K F Q Q T T T T That correlates the temperature increment above the local temperature to the unknown depth H, to the discharge Q and the conduit shape factor SF. Then H KRFgt Tw! T0011 < Q 5.5 # 10! 4SF 2 ,2 +2 *2 )2 (2 Comparing with [Eq. 4.12] we obtain the solution for thermally disturbed rock (Q average yearly discharge) ( ( ) + + # < "!FS Q H H4 010 5 5 1 This solution shows that the evaluations made with the Desio formula [Eq. 4.12] are deeply underestimated, unless for very small discharges. In fact we can write H H01 < q 0 1 [4.13] It is then possible to see that the q-number is essentially the “amplification” term of estimated depth H0. The main difficulty in these formulas it is the estimation of critical discharge Qc, which requires the knowledge of the conduit shape factor, in generally unknown. In case studies it is necessary to take into account the geological context to estimate the probable conduit shape in order to calculate the critical discharge Qc. Let us do an example. In many case, for instance, we can assume a “U” shape for the whole drainage system. A similar conduit can simply be approximated with a conduit of length L at depth H, because its two vertical branches do not matter for the shape factor, being merged in a rock shielded by the deep drainage. Then [4.14] The Table 4 gives the value of inverse hyperbolic cosine term for typical conduit radius and depths. TABLE 4 H # R $ 0.1 m 1 m 10 m 100 m 200 8.3 6.0 3.7 1.3 500 9.2 6.9 4.6 2.3 2000 10.6 8.3 6.0 3.7 It is possible to see that, unless the nearby “pathological” cases of the last column, the denominator in [Eq. 4.14] is not far from 2 % ; therefore for this conduit it is possible to assume a shape factor roughly equal to its length L SF The formula that estimates the water provenance depth can then be approximated as [4.15] 010 0T T F K Hw gt R! "01 R H L SF/ cosh 21 !" / ( ) + # < .!L Q H H4 010 5 5 1

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.20 In general a typical deep circuit has a very large size, many kilometres. If we call L* its length expressed in kilometres we have, for the circuit described above 5 0 L Qc. And finally [4.16] It is a simple formula that estimates the drainage depth as a function of discharge. As an example, let us consider a spring with an average discharge Q=20 kg s-1, and a temperature 20C above the local average. The depth estimation [4.12] gives H0=830 m. If the hydro-geological context suggests a circuit length L*=20 km, the critical discharge is Qc=11 kg s-1, and then the q term is almost equal to 2. Our formula estimates then a depth H of 2.5 km for the circuit, much more than supposed… With this value we can return to the geological context and, if we have other information (like the circulation time), to more accurate estimations of the ratio between the depth and the radius of a deep drainage system. A better estimation of provenance depth with [Eq. 4.13] and [Eq. 4.14] is then possible. It is nevertheless better to remember that we are working with the assumption that the system has attained stationary condition; the above formula is then correct for water fluxes that persist from very long times, much longer than & teq of [Eq. 1.1]. Temperature changes into the system It is possible to perform a last “calibration”. We have already noted that we are assuming that it is possible to define univocally a system temperature T, but this is not always true. It is possible to speak univocally of “system temperature” if each part of the system is uniformly heated by the energy flux, for instance if the head and the tail of water flux are mixed (for instance, when the water enters in a spherical deposit). Nevertheless this is not the usual situation because in a real conduit the water enters with a temperature T0 and flows warming up to the final temperature T. It is possible to perform a last step, considering a long conduit L, along which the water is heated. The term T is now the temperature at the length x, in a section dx with shape factor sF. The thermal flow and the temperature increase in that section is then given by [Eq. 4.5] 01dT Q C T T s T T H F H s F dW dW dWw F gt F gt up M" ! " 0 0" And then dT S s T T T T Q C H S FF F w F gt" ( ( ) + + !0" This is an equation that could solve the problem, if we would able to integrate the left part, but unfortunately this is not possible. We have written sF and not dSF (as would be natural) because it is not possible to pass from the equation that gives SF as a function of L, to the contribution of a part dx of L to SF. At the end of the third chapter, we have noted that SF does not linearly depend on each dx part, because it comes from an average on the whole space and system, and it is not possible to consider it as the result of an integration on some dx. For instance, the contribution of the dx at one conduit edge covers a cone above it, and the local sF is like that given by a small sphere, whereas the dx in the middle of conduit gives a very small contribution. We can nevertheless integrate the last equation assuming the false approximations sF=dSF. It makes no analytical sense but it probably introduces an error smaller than the assumption of uniform system temperature. Then for a conduit buried in a semi-infinite medium at depth H dT L dx T T T T Q C H S Fw F gt" ( ( ) + + !0" Using Qc becomes 1 q FgtH Kr T T T T0 2 ,2 +2 2 )2 (2 dx L QcQ T T01dx L dT And integrating on x from 0 to L and on T from T0 to T we have ( ) + < 2 10L Q H H

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.21 T T T T001exp QcQ 2 ,2 +2 *2 ) 2 ( 2 Adding and subtracting T0 from the equation left side, we have T T0" T T001! T T001exp 1 q 2 ,2 +2 *2 )2 2 [4.17] This has to be compared with [Eq. 4.10]. It is a simple and nice equation that describes the water heating during a flow. The Qc term has returned, and continues to be the scale discharge of deep conduits. If the effective discharge Q is large (in comparison with Qc) the water temperature at the outflow is near the T0, and if q is near 0 the T=T”. The shape factor has disappeared, because this equation is valuable everywhere the (strong and false) sF=dSF approximation is valuable. Nevertheless we can suppose that it is reasonable model, and we guess that the last equation gives a fair approximation of natural heating processes along a conduit. It is useful to invert again the problem to obtain the estimated depth crossed by water of a spring at temperature Tw. We have 8 7 6 5 4 3 ( ( ) + + ! q H K F T Tr gt w1 exp 10 And with simple passages [4.18] That corresponds to [Eq. 4.13], and reduces to it for q~0 and q>>1. With the same assumption of [Eq. 4.16] on shape factor, and L* in [km], it gives [4.19] That can be considered a reasonable formula to estimate the water provenance depth, for water flows that have attained a steady state situation with rock. Returning to the previous example, of a spring with Q=20 kgs-1, and a temperature 20 C above the local average, H0=830 m, Qc=11 kg s1, and q=2. The corrective term to be applied to H0 with [Eq. 4.16] it is a factor 3, but now [Eq. 4.19] gives a factor 2.54. Really the temperature variability along the conduit gives [Eq. 4.19] a final heating at a depth smaller than in the case of a “global” heating [Eq. 4.13], but the difference does not appear as too significant if compared with the intrinsic uncertainties of such problems. Steady State Geothermal Power Plant In the previous chapters, it has been shown implicitly a way to extract power from underground, using a deep conduit that focuses on itself large amounts of geothermal energy. This is deeply different from the usual Geothermal Power Plants, which extracts energy (or, better, are believed to extract, because cool water always focus on itself the temperature field) from hot rock, directly cooling it. In principle, when the rock is cooled the power plant stops its work. Here we have shown that the deep cooling effect acts as an energy attractor on the cooled rock, and then that a power plant working in such way, it will never end its fuel. We want here to make the next step, looking for its “constructive” efficiency. Consider a fluid that transfers energy QH from a “hot” source at TH to a colder source at TL. Does this process produce work? If the energy transfer is made with “special” systems it does: They are called “thermal engines” and use the energy flow from TH to TL to produce work. A power plant is said to “produce” energy, but this is trivially false because the energy cannot be neither produced nor destroyed: It stores energy at a very low entropy (“work”) in an entropy flow from a high temperature (low entropy) to a low (high entropy). The Second Principle of Thermodynamics states that the efficiency -that is the ratio between the work rate given L and the heat rate absorbed Wof a reversible thermal engine working between the two sources is 01 ( ( ) + + ! " 8 7 6 5 4 3 ( ( ) + + ! "!q H q F T T K Hgt w r1 exp 1 1 exp 10 1 0 ( ( ) + + ! Q L H H 2 exp 10

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.22 H L H L H H MaxT T T T T Q L " 1 But we have that ( ( ) + + &H L L HT T T Q S 1 And then S T LL Max& That is, the maximum work available is given by the product of the temperature of cold source and entropy change during cooling, which is the Free Energy variation in the transformation. In the case of interaction cave-geothermal field, the cave acts as a thermal sink, in two phases. In the first step it intercepts a flux of geothermal energy W (low entropy, temperature T”) from downward, which results in a water temperature increase from T0 to T. In the second step the energy it is released as “disordered energy” to the atmosphere (high entropy, T0) at the spring. For instance, considering the deep conduit as a geothermal power plant, we have that its entropy production per time unit & t -we are dealing with discharge Qis & S "! W T < W T0 2 ,2 +2 2 )2 (2 & t SFFgtH q 1 < q 2 ,2 +2 2 )2 (2 1 T < 1 T0 2 ,2 +2 2 )2 (2 & t Calling 0 0 0" T K H F T T Tr gt" "? With some work and using [Eq. 6.1] & S & t WMT0 ?q 1 < q 2 ,2 +2 2 )2 (2 1?< 1 < q 2 ,2 +2 *2 ) 2 ( 2 This shows that the entropy production goes to zero for q=0 and q= @ because if the water flow is very small the fluid final temperature is quite high but the total energy removed is very small; from the other side, if a lot of water flows into the conduit, its final temperature is essentially T0, then the entropy is able to flow between the rock and the water, but it is not finally transferred to the atmosphere and to an external “final user”. It is easy to calculate the value qM for which the entropy flow attains its maximum qM"?< 1 T T0T0 < 1 T T0 That we can substitute in the previous equation to obtain the maximum of entropy flow. In natural cases the term T0 is some 280 K, the T” some 350K, then the ratio is slightly more than 1, and then qM" 1 < T T0T0 1 < 1 2 FgtKrT0 2 ,2 +2 2 )2 (2 H Which gives, erasing the second-order terms, the maximum power production of this geothermal power plant & L & t -2 ,2 +2 *2 )2 (2Max" T0& SMax" 1 4 WMT T0T0 2 ,2 +2 *2 )2 (2 It is necessary to emphasize the difference between the subtracted power WM and the maximum available work (or power) LMax. The first is interesting to make some use that requires enthalpy at constant temperature, as it is the case of ice melting or water evaporation, for which WM, not LMax, is used. But to create structures we need “work” also in the physical sense: Order, available work. Therefore, the LMax terms in each water heating and rock cooling processes are directly connected with the entropy rate at disposal for constructive processes, that is, they may appear as the building rate of ordered structures, like conduit networks. Geothermics and phreatic conduit genesis We have observed above that the initial purpose of this work was to show that the geothermal energy flux could not participate in the characterisation of cave climate and then, for instance, to speleogenesis (Badino, 2005). As the reader has seen, we are showing exactly the contrary. Here we are going to make the last step giving some ideas about the geothermal role in the genesis of phreatic conduits and in general of underwater drainage networks.

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.23 During deep flow the underground waters are warmed from their starting temperature T0 up to the final T, that has a theoretical maximum at T”, as shown above. What is the typical temperature increase? We have seen that a conduit is able to shield the geothermal flux like a plane watertable does, then the scale temperature increase is around the value given by [Eq. 2.1] & Tgt" 0.06 4.2 # 103P Pout01 500 P ; C9: Therefore, in real cases it is in the range between 0.2 and 3 C, a temperature drop that water gains during flow between the cave bottom and the springs. Now we know that this temperature change happens also along conduits, not only in the “watertable”, and that the power release is concentrated on the conduit surface walls. How does this warming affect the water chemistry? We can outline the chemical behaviour of water saturated of calcium carbonate entering in flooded conduits. It is well known that the carbonate dissolution in water is very complex (Snoeyink, 1980), (Ford and Williams, 1989), (Dreybrodt, 2000) because the equilibrium state results from the combined equilibriums of different, interconnected reactions, which depends on temperature, local pH and the presence of other dissolved salts with common ions. In the simplest case, the first equilibrium reaction gives the amount of dissolved carbon dioxide, for which in usual conditions the Henry Law holds, stating that the dissolved gas decreases with temperature and it is proportional to its partial pressure above the water surface. Therefore, its quantity depends also on the kinetics of gas transport until the surface, if it does exist. The other reactions, which involve only water and carbon dioxide, are the dissociation of carbonic acid in calcium bicarbonate and H+, the dissociation of bicarbonate and the equilibrium H+ and OHin water. These dissociations tend to increase with temperature thanks to the increase of available energy. The last main reaction describes the equilibrium between the calcium carbonate and water enriched with carbon dioxide. The carbonate dissolution releases ions that are in part the same already present in water. This complex system forces to find the solution of many different equations describing equilibrium kinetics, charge and mass conservation. General solution charts are given in (Ford and Williams, 1989); they show the saturation values at the equilibrium for various initial partial carbon dioxide pressures. It results that in open systems (with release of carbon dioxide excesses) the warming of a calcite saturated water gives, without exceptions, a super-saturation and then provokes a calcite, or aragonite, deposition. In a closed system this behaviour changes in a complex way. Generally a super-saturation is produced, but if the initial CO2 partial pressure is below 200 Pa (0.002 atm) and the temperature is below 30 C, a calcite under-saturation appears as result of water heating, as larger as colder is the water. The typical carbon dioxide partial pressure in free atmosphere is 3.5 # 10-4 atm, then at 10 C the calcite equivalent content at the saturation is around 12 mg l-1. A water temperature increase of 1 C result in a saturation value of 0.02-0.04 mg larger, i.e. with a flux of 1 m3s-1 it gives around 103 kg of dissolved rock per year. It is a small figure but it suggests that further studies are necessary to a more complete understanding of saturation conditions as a function of temperature, of chemically complex waters in a closed system. In any case indirect evidences of effectiveness of speleogenetic processes induced by geothermal heating in phreatic conduits can be found, because if these processes are possible, they have to affect the network morphologies: 1) the geothermal energy is released only in the lowest conduit walls, then the dissolving characteristics have to depend on the rock surface orientation; 2) a deep conduit shadows completely the upper rock, then the formation of a conduit that cross the rock above another is hampered, and this affects the whole drainage conduit structure. Similar processes can probably play a part also in the deep drainage network formation in glaciers (Badino, 2002), but either ice or limestone, a lot of work has still to be done for a better understanding of geothe rmal role in karst.

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.24 Conclusions The estimations of temperature fields inside mountains are important for speleogenesis and for underground climate studies, but also for many cases which require an energy balance on a sub-geological time-scale, like glacier stability, geothermal spring studies, deep hydrogeological analysis, tunnel drilling and so on. We have shown that these potentially cumbersome modelling can be reduced to simple calculations that allow quite accurate estimations of energy absorbed by deep structures and of provenance depths of geothermal waters. These results could also be applied for remote sensing of deep drainage structures and for construction of inexhaus tible geothermal power plants, but at present these applications appear to meet insurmountable practical difficulties. References AA. VV., 2004. Under the Desert: the Mysterious Water of Cuatro Cinegas. La Venta. Badino G. 1995. Fisica del Clima Sotterraneo. Memorie IIS 7, II Badino G. 2000. I Gradienti di Temperatura nei Monti, un Indicatore Esplorativo. Talp-FST 21 72-80 Badino G. 2002. The Glacial Karst. Proc. of V Int. Symp. on Glacier Caves and Cryokarst 2000, Nimbus 23, VII, 1/2002 Badino G. 2004. Cave Temperatures and Global Climatic Change. Int. J. Speleol. 33 Badino G. and Forti, P. 2005. L’eccezionale Ambiente della Cueva de los Cristales. Proc. “Le Grotte di Miniera”, Memorie IIS XVII, II Badino G. 2005. Clouds in Caves. Speleogenesis and Evolution of Karst Aquifers 2 (2) www.speleogenesis.info. Badino G. 2005. Nuovo Collegamento Ferroviario Torino-Lione, Temperature del Tunnel di Bussoleno, Geodata, Torino, unpubl. Balcerzak M. and Raynor S. 1961. Steady State Temperature Distribution and Heat Flow in Prismatic Bars”, Int. J. Heat Mass Transfer 3 : 113-125, Bejan A. 1993. Heat Transfer John Wiley and Sons. Benderitter Y., Roy B. and Tabbagh A. 1993 Flow Characterization through Heat Transfer: Evidence in a Carbonate Fractured Medium. Wat. Res. Res. 29, 11 : 3741-3747. Bohren C. and Albrecht B. 1998. Atmospheric Thermodynamics Oxford Un. Press, 402 pp. Bogli A. 1980. Karst Hydrology and Physical Speleology Springer-Verlag. Carslaw H. and Jaeger J. 1959. Conduction of Heat in Solids, Oxford-Clarendon Press. Catalano P. 1993. Laboratori Sotterranei: Relazione Geologica. Unpublished, INFN. Celico P. 1986. Prospezioni Idrogeologiche Liguori. Dreybrodt W. 2000. Equilibrium Chemistry of Karst Water in Limestone Terranes. In Speleogenesis: Evolution of karst aquifers Klimchouk, A., Ford, D.C., Palmer, A.N., and Dreybrodt, W. (Eds.), Nat. Speleol. Soc., USA: 130-135. Ford D. and Williams P. 1989. Karst Geomorphology and Hydrology, Unwin Hyman. Goy L., Fabre D. and Menard G. 1996. Modelling of Rock Temperatures for Deep Alpine Tunnel Projects. Rock Mech. Rock Engng 29 Guichonnet P. 1967. Il Traforo del Monte Bianco, Mondadori. Hahne E. and Grigull U. 1975. Formfaktor und Formwiderstand der Stationaren Mehrdimensionalen Warmeleitung, Int. J. Heat Mass Transfer 18 : 751-767 Holman J. 1996. Heat Transfer, MacGraw-Hill. Isachenko V., Osipova V. and Sukomel A. 1969. Heat Transfer, MIR. Jeannin P., Liedl R. and Sauter M. 1997. Some Concepts about Heat Transfer in Karstic Systems. 195-198. Kays W. 1966. Convective Heat and Mass Transfer. McGraw-Hill. Lee et al. 1966. Heat Flow and Volcanic Temperatures. Handbook of Physical Constants The Geological Society of America. Koenigsberger J. and Thoma E. 1906. Uber die Beeinflussung der geothermischen

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Giovanni Badino / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.25 Tiefenstufe durch Berge und Taler. Eclog. Geol. Helv. IX (1) Laidler K. and Meiser J. 1995 Physical Chemistry. Houghton-Mifflin Co. Lismonde B. 2002. Arologie des Systmes Karstiques CDS Isre. Luetscher M. and Jeannin P.-Y. 2004. Temperature distribution in karst systems: the role of air and water fluxes. Speleogenesis and Evolution of Karst Aquifers 2 (2) (from Terra Nova 16):, 344– 350 Nashchokin V. 1979. Engineering Thermodynamics and Heat Transfer. Mir, 573 pp. Ozisik M. and Necati. 1993. Heat Conduction ,.Wiley InterScience. Schoeller H. 1962. Les eaux souterraines. Masson. Snoeyink V.and Jenkins D. 1980. Water Chemistry John Wiley & Sons. Szechy K. 1973. The Art of Tunnelling, Akademiai Kiado, Budapest. U.S. Bureau of Mines. 1996. Dictionary of Mining, Mineral, and Related Terms, CDROM. Verhoogen J. 1956. Temperatures within the Earth. In Physics and Chemistry of the Earth Pergamon Press, I.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info Ground water flux distribution between matrix, fractures, and conduits: constraints on modeling William B. White (1) and Elizabeth L. White (2) (1) Materials Research Institute and Department of Geosciences, The Pennsylvania State University, University Park, PA 16802 USA. Email: wbw2@psu.edu (2) Hydrologic Investigations, 4538 Miller Road, Petersburg, PA 16669 USA Abstract Calculations are presented to show the re lative contribution of the matrix, fracture, and conduit permeability to the overall f low of ground water through a karst aquifer. The conceptual model is a cross-section spanning the full width and thickness of the aqu ifer. A constant, but adjustable head is assumed. The rock matrix is characterized by an adjustable hydraulic conductivity. Varying proportions of fractures and conduits of adjustable fracture ap ertures and conduit diameters were the calculational parameters. Calculations used DarcyÂ’s law for matrix flow, the cube law for fracture flow, and the Darcy-Weisbach equation for conduit flow The results show a surprising dominance of fracture flow in the early stages of aquifer development. A focusing mechanism is needed to localize the flow into a relatively small number of conduits. Keywords: ground water, aquifer, triple permeability, speleogenesis. Introduction Attempts to describe, analyze, or model ground water flow in karstic ca rbonate aquifers usually begin with some aspect of the triple permeability model (White, 1999; Worthington et al., 2000; White, 2002). The three components are matrix permeability, fracture permeability, and conduit permeability. Each contributes to the flow field but frequently only one or at most two of these components are included in the calculations. The choice of components and decisions concerning what can be neglected are often based on little more than guesswork. The object of the present paper is to show the relative contribution of each of these permeability components to the overall flux of moving ground water. The calculations are based on a crosssection spanning the full width and thickness of the aquifer. By varying the co ntributions of each of the permeability components a measure of their relative importance to the overall flow system is obtained. The intent is not to provide an aquifer model, but rather, by using actual numerical values, provide some insight into when and under what circumstances, one or more of the components of aquifer permeability can be neglected. Such calculations provide some constraints on the various equivalent porous media models such as the one developed for the Edwards aquifer in Texas (Scanlon et al., 2003). The triple permeability concept Table 1 gives the essential characteristics of the three components of the triple permeability model for karst aquifers. The matrix, fracture, and conduit permeability are, essentially, independent components. Any specific real aquifer will have a mix of these contributions. There also exist real aquifers in which one of the components is completely dominant.

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W.B.White and E.L.White / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 TABLE 1 Characteristics of the three components of the triple permeability model. PERMEABILITY APERTURE TRAVEL TIME FLOW MECHANISM GUIDING EQUATION DISTRIBUTION Matrix m to mm Long Darcian flow field. Laminar !" 2Nd g L v hf# $ % Continuous medium Fracture 10 m to 10 mm Intermediate Cube law. Mostly laminar; may be non-linear components 3b f C h Q % Localized but statistically distributed. Conduit 10 mm to 10 m Short Darcy-Weisbach. Open channel and pipe flow. Turbulent r g v L f hf42% Localized Calculations The framework used for the calculations is shown in Fig. 1. A fixed cross-section of aquifer is assumed. The aquifer is taken as rectangular, 100 meters thick and one k ilometer wide, giving a cross-sectional area of 105 m2. The aquifer crosssection is an adjustable boundary and can be set to any value, providing that the area is large compared to solution features that are embedded in it. The head is assumed to be constant and uniform across the aquifer area. This is an extreme assumption but one that eliminates concern for the water table and also the variable heads that would be characteristic of most aquifers containing conduits. The rock matrix is characterized by an adjustable hydraulic conductivity. Varying proportions of fractures and conduits of adjustable fracture apertures and conduit diameters are the calculational parameters. A further assumption is that there is sufficient recharge behind the aquifer cross-section to provide whatever flow is called for by the calculations. The guiding equations for the permeability (Table 1) show that the flow rate, Q, varies linearly with the head in laminar flow but with the square root of the head in turbulent flow. The head becomes a scaling variable. The head, or hydraulic gradient, dh/dL, is here set equal to 0.01, a nominal value for small karstic drainage basins. The guiding equations contain the density and viscosity of water, both of which are functions of temperature. A temperature of 10 C was selected as typical of karst ground waters but the variation in the parameters over the range of temperatures expected in karst aquifers is relatively small. Fig. 1. Sketch showing the aquifer cross-section used for calculations. The matrix component Ground water flow through the limestone or dolomite bedrock is not intrinsically different from ground water flow in any other aquifer. The guiding equation is DarcyÂ’s law. However, calculations must use hydraulic conductivities for the rock. Such data are sparse. Most hydraulic conductivities are based on pump tests on wells and those data are dominated by the fracture flow component. Intrinsic hydraulic conductivities of the bulk rock must be measured on core samples in the laboratory. Some representative data are shown in Table 2.

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W.B.White and E.L.White / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 The flux through the matrix is a fixed quantity for chosen values of hydraulic conductivity. Lines of constant flow rate we re calculated for Kentucky Mississippian limestone (K = 2 x 10-11 m/sec) which should be typical of many Paleozoic limestones and dolomites. Other lines of constant flow were calculated for the Edwards Limestone (K = 1 x 10-8 m/sec) and for the mean value of the Floridan aquifer (K = 3.65 x 10-6 m/sec). The latter is a carbonate aquifer in which matrix flow is a dominant component. TABLE 2 Hydraulic conductivities for some carbonate rock aquifers Rock Unit K (m/sec) Reference Floridan Aquifer Budd and Vacher (2002) Wackestone 1.97 x 10-7 Packstone 9.61 x 10-7 Grainstone 3.82 x 10-6 Sucrosic dolostone 9.61 x 10-6 Mammoth Cave, Mississippian limestone 2 x 10-11 Worthington (1999) Silurian dolomite, Ontario 1 x 10–10 Worthington (1999) English chalk 1 x 10-8 Worthington (1999) Pliocene limestone, Yucatan, Mexico 7 x 10-5 Worthington (1999) Edwards Aquifer, Texas, Cretaceous 1 x 10-8 Worthington et al. (2002) Swabian Alb, Germany, Jurassic 8 x 10-9 Worthington et al. (2002) The fracture component The idealized model for fracture flow assumes a fracture with plane parallel walls and a uniform aperture. For the ideal case, the cubic law can be derived theoretically from the Navier-Stokes equations. L h b g w Q & & %$ #123 [1] Here, Q = flow rate in m3/sec, w = fracture width in m, = density of water = 999.7 kg/m3, g = gravitational acceleration = 9.8 m/sec2, b = full aperture of the fracture in m, and = viscosity of water = 1.307 x 10-3 Pa sec. It has been long recognized that real fractures do not have uniform apertures and that the walls are not parallel. Witherspoon et al. (1980) resolved this problem by compacting the constants of equation [1] into a single constant and then adding an empirical friction factor, f, to give the form of the equation shown in Table 1. More recent work (e.g. Brush and Thomson, 2003; Konzuk and Kueper, 2004) has proposed more quantitative descriptions of rough-walled fractures but in general the results change by no more than a factor of 2. For the rough calculations in this paper, the plane-walled fracture (equation [1]) should suffice. The fracture width, w is the total extent of fractures measured perpendicular to the flow direction. There may be multiple fracture sets at different angles with respect to each other. In the model assumed in Fig. 1, there are both vertical fractures and horizontal bedding plane partings in an assumed horizontal bedding. The total width of vertical fractures is the aquifer width/mean fracture spacing. The total width of horizontal fractures is the aquifer thickness/mean spacing of bedding plane partings. With the aquifer dimensions given in Fig. 1 and a typical 10 meter spacing for both vertical fractures and bedding plane partings, the model would contain 18,900 meters of fracture assuming that the boundary planes are not fractures. A second curve was calculated using only vertical fractures with a 50 m spacing. Fracture flow calculations were cut off when the aperture reached 0.01 m as this is the dimension at which turbulent flow is expected to develop. The cubic equation applies only to laminar flow. If all fractures were to develop to the largest aperture, fracture flow would completely dominate the flow system. Something like this occurs in aquifers with maze cave development. In most aquifers, rearrangements of the flow path would have occurred before all of the fractures reached this very large aperture.

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W.B.White and E.L.White / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 The conduit component A single conduit is assumed as sketched in Fig. 1. The conduit radius, r, is taken as the dependent variable. Conduit flow in the laminar regime is described by the Hagen-Poiseuille equation [2]. dL dh r g Q$ # (84% [2] Turbulent flow in a cond uit is described by the Darcy-Weisbach equation. Written as volume flow and taking the cross-section as a circular conduit gives 2 1 2 5 2 12 ) + ) ) + , % dL dh r f g Q( [3] Application of equation [3] requires numerical values for the Darcy-Weisbach friction factor, f, which must be determined empirically. The friction factor relates to the wall roughness and, in the case of a conduit that behaves as a uniform pipe, estimates of f have been made from irregularities, usually scallops, on the conduit wall. Most conduits, however, are not uniform pipes. Breakdown, sediment chokes, and varying passage shapes also contribute to the effective roughness. An alternative approach is to measure all other parameters in equation [3] and then back calculate f. The two approaches give dramatically different results as shown in Table 3. Because the friction factor enters the Darcy-Weisbach equation as a square root, the effect of the different numerical values is somewhat muted. Calculations were made with the smallest value (0.039) and one of the largest values (130). TABLE 3 Estimates of the Darcy-Weisbach friction factor Location From Discharge From Roughness Reference Mendips (UK) 24 – 340 --Atkinson (1977) Castleguard (Canada) 0.87 – 2.31 0.33 – 0.90 Atkinson et al. (1983) Morecombe Bay (UK) --0.077 Gale (1984) Glomdalsvatn (Norway) 0.116 0.039 Lauritzen et al. (1985) Turnhole (KY) 27 --Worthington (1991) Friars Hole (WV) 46 – 74 --Worthington (1991) Holloch (Switzerland) --0.322 Jeannin (2001) Maligne Basin (Canada) 130 --Smart (1988) Discussion and conclusions The results of all calculations are plotted in Fig. 2. The x-axis gives the aperture, either fracture aperture or conduit radius. The y-axis shows the flow volume that would be expected under the specified conditions of aquifer cross-section and hydraulic head. Changing the aquifer cross-section and the hydraulic head would shift the numerical positions of the curves but not their relative pattern. For matrix flow the “aperture” is the total crosssection of the aquifer so the matrix components plot as horizontal straight lines whose vertical position depends only on the assumed hydraulic conductivity. As expected, in the dense, low permeability Paleozoic limestones, the matrix flow is negligible. It becomes more important in more permeable limestones and yields a significant contribution to the flow field when K exceeds values of 10-6 m/sec. A surprise in the calculations is the dominance of fracture flow. Most observed fractures in carbonate aquifers have apertures in the range of hundreds of micrometers. If these fractures were enlarged by dissolution to the millimeter to centimeter range, fracture flow would completely dominate the system. What happens in most aquifers is the focus of the flow into a few localized pathways early in the development of the system. Because of the acceleration of dissolution kinetics at the critical aperture in the one-centimeter size range, a single conduit or small set of condu its grow at the expense of near by fractures. Lowering hydraulic heads in the conduits cause the conduits to act as drains and the simplified model used for the present calculations is not applicable. In those aquifers where geologic factors prevent the focus of flow into single conduits, fracture enlargement does continue with the product of maze caves. These results are consistent with the conclusions of Worthington et. al. (2000) that although the main

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W.B.White and E.L.White / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 Fig. 2. Discharge through the components of the triple permeability system as a function of aperture. The horizontal lines are the matr ix contribution. The lines labeled 10 m and 50 m spacing are the fracture contributions at the specified fracture spacing. The lines for the conduit contribution are labeled with the chosen values for the Darcy-Weisbach friction factors, f = 0.039 and f = 130. portion of the flow in karstic aquifers is through the conduits, the main portion of the storage is in the fractures. The results are also consistent with the high well yields obtained from fractured dolomite aquifers. Conduit systems do indeed dominate the flow system in many carbonate aquifers. However, in order for them to develop there must be focusing mechanisms to drain off water from the fractures before they develop wide apertures. Lowe’s (2000) inception horizon concept is of importance as one of the focusing mechanisms. Further, the hydraulic gradient must be sufficient to drive the competitive process that leads to single conduits. Low gradient aquifers are again associated with maze caves and a dominance of fracture flow. References Atkinson, T.C. 1977. Diffuse flow and conduit flow in limestone terrain in the Mendip Hills, Somerset (Great Britain). Journal of Hydrology 35 : 93-110. Atkinson, T.C., P.L. Smart and T.M.L. Wigley.` 1983. Climate and natural radon levels in Castleguard Cave, Columbia Icefields, Alberta, Canada. Arctic and Alpine Research 15 : 487502. Brush, D.J. and N.R. Thomson. 2003. Fluid flow in synthetic rough-walled fractures: Navier-Stokes, Stokes, and local cubic law simulations. Water Resources Research 39 : 1085, doi:10.1029/2002WR001346. Budd, D.A. and H.L. Vacher. 2002. Facies control on matrix permeability in the upper Floridan Aquifer, west-central Florida: Implications for diffuse flow. Karst Waters Institute Special Publication 7 14-24. Gale, S.J. 1984. The hydraulics of conduit flow in carbonate aquifers. Journal of Hydrology 70: 309-327. Jeannin, P.-Y. 2001. Modeling flow in phreatic and epiphreatic karst conduits in the Hlloch Cave (Muotatal, Switzerland). Water Resources Research 37 : 191-200. Konzuk, J.S. and B.H. Kueper. 2004. Evaluation of cubic law based models describing single-phase flow through a rough-walled fracture. Water Resources Research 40 : W02402, doi:10.1029/2003WR002356. Lauritzen, S.-E., J. Abbott, R. Arnesen, G. Crossley, D. Grepperud, A. Ive and S. Johnson. 1985. Morphology and hydraulics of an active phreatic conduit. Cave Science 12 : 139-146. Lowe, D.J. 2000. Role of stratigraphic elements in speleogenesis: The speleo inception concept. In Speleogenesis: Evolution of Karst Aquifers A. Klimchouk, D.C. Ford, A.N. Palmer and W. Dreybrodt, Eds., National Speleological Society, Huntsville, AL, USA, pp. 65-76. Scanlon, B.R., R.E. Mace, M.E. Barrett and B. Smith. 2003. Can we simulate groundwater flow in a karst system using equivalent porous media models? Case study, Barton Springs, Edwards Aquifer, USA. Journal of Hydrology 276 : 137-158. White, W.B. 1999. Groundwater flow in karstic aquifers. Chap. 18 in The Handbook of Groundwater Engineering J.W. Delleur, Ed., CRC Press, Boca Raton, FL, pp. 18-1 – 18-36. White, W.B. 2002. Karst hydrology: recent developments and open questions. Engineering Geology 65: 85-105.

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W.B.White and E.L.White / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 Witherspoon, P.A., J.S.Y. Wang, K. Iwai, and J.E. Gale. 1980. Validity of cubic law for fluid flow in a deformable rock fracture. Water Resources Research 16: 1016-1024. Worthington, S.R.H. 1991. Karst hydrogeology in the Canadian Rocky Mountains. Ph.D. thesis, McMaster University, Hamilton, Ontario, 380 pp. Worthington, S.R.H. 1999. A comprehensive strategy for understanding flow in carbonate aquifers. Karst Waters Institute Special Publication 5 30-37. Worthington, S.R.H., D.C. Ford and G.J. Davies. 2000. Matrix, fracture and channel components of storage and flow in a Paleozoic limestone aquifer. In Groundwater Flow and Contaminant Transport in Carbonate Aquifers I.D. Sasowsky and C.M. Wicks, Eds., A.A. Balkema, Rotterdam, pp. 113-128. Worthington, S.R.H., G.M. Schindel and E.C. Alexander, Jr. 2002. Techniques for investigating the extent of karstification in the Edwards Aquifer, Texas. Karst Waters Institute Special Publication 7 173-175.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info Karst and Caves of Ha Long Bay Tony Waltham Department of Civil Engineering, Nottingham Trent University, UK Email: tony@geophotos.co.uk Updated and re-published from: International Caver 2000, pp 24-31. Introduction Ha Long Bay is distinguished by the hundreds of small limestone islands that rise steeply or vertically from its shallow waters. Its dramatic and beautiful landscape is deservedly famous as one of the worldÂ’s outstanding natural sights, but it is also a UNESCO World Heritage Site of international geomorphological significance (Fig. 1). The bay lies on the northeastern coast of Vietnam, immediately east of the Red River delta. It is bounded on the north by the mainland hills either side of Ha Long City (also known as Hong Gai), to the south by the open waters of the Gulf of Tonkin, to the west by Cat Ba Island, and to the east by islands of sandstone (Fig. 2). Ha Long Bay has an area of about 1500 km2, and contains nearly 2000 limestone islands. The caves described here were all visited during an assessment of the bayÂ’s geomorphology with respect to its position as a World Heritage Site. Records of other caves in Ha Long Bay are sparse. A British team led by Howard Limbert mapped the Hang Hanh stream cave in the mainland limestone along the north shore of the bay; and a French team led by Marc Faverjon explored caves in the islands east of the bay, and also a few in Ha Long Bay itself. Locality names are here tr anslated into English, except for the cave names which are left in Vietnamese. The key terms are: dao = large island; hon = small island or rocky tower; hang = tunnel or passage cave; dong = chamber cave. Fig. 1. The view out from Hong Gai harbour, with the limestone islands extending out into Ha Long Bay.

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 Fig. 2. Location of Ha Long Bay in northern Vietnam. Fig. 3. A simplified map of Ha Long Bay, showing the larger islands and the outline geology, and naming some of the larger caves. The islands of Ha Long Bay are all cut in a folded sequence of Carboniferous and Permian limestones that reaches to more than 1000 m thick. These pale grey limestones are strong, fine-grained materials, ideal for the development of karst. Beds vary from 500 mm to 5 m thick, and very thin shale partings occur on many of the bedding planes; chert is a minor feature, and there is patchy dolomitisation at a few localities. Across most of the bay, the limestones dip to the west, but the structure is complicated by north-south faults. Cat Ba Island is distinguished by limestone hills with slope angles lower than those of the bay islands; it appears to be formed largely in a more thinly bedded series of Carboniferous limestones than those that underlie the bay. Small areas of karst on islands east of the bay are formed on Devonian limestones that occur within the sandstone sequence. The mainland north of Ha Long Bay is formed largely of Triassic coal measures. These are complexly folded and contain anthracite seams up

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 to 50 m thick that are mined in large open pits. The boundary of the coal measures with the limestone is almost along the coast, and this is largely faulted. The geological structure of the Ha Long Bay limestones appears to be very complex, and some of the drowned valleys between islands are probably fault-guided. Bedding is close to horizontal in the eastern part of the bay, but elsewhere dips at any angle, and small overfolds occur in the western part of the bay. The structure has little influence on the karst geomorphology, as slope profiles on the islands are largely independent of the limestone dips. The karst of Ha Long Bay The strong pure limestones of Ha Long Bay have been eroded into a mature landscape of fengcong and fenglin karst. This e volved by normal subaerial erosion of the limestone, but was then invaded and slightly modified by the sea at a late stage. The hundreds of rocky islands which form the most beautiful and famous landscapes in the bay are individual towers in a classic fenglin landscape where the intervening plains have been submerged by the sea (Fig. 3). Most towers reach heights of 50 to 100 m, with height/width ratios up to about 6. Many towers have vertical walls on all or most sides; these continue to evolve by rockfalls and large slab failures. In 1997, a large slab peeled off a small island north of Cat Ba, to create a new vertical face; the main failure surface was on vertical fractures, part of which had been opened to form a cave subsequently partly refilled with large stalactites, and the fallen block of around 500 m3 now lies in the sea where it is being eroded by dissolution and wave action. Many of the towers have very old cave remnants preserved within them, and many have foot caves that are relicts of their undercutting at various levels (Fig. 4). Clusters of limestone hills form the larger islands within Ha Long Bay, and represent fine examples of fengcong karst. Summits are generally at around 100 m above sea level, and the highest peaks reach heights of over 200 m. Their profiles are mostly very steep cones; except around their marine margins, vertical cliffs are only minor components, as they have not been subject to lateral undercutting where their internal dolines do not reach down to base level. These conical hills also contain remnants of old cave passages. Some of the fengcong hills have individual slopes or sides which are formed on steep bedding planes within the limestone; beyond these sites the geological structure has very little influence on the limestone hill profiles. Fig. 4. Diagrammatic profile showing the main types of cave in the Ha Long Bay karst. Fig. 5. The Luon hong, with the sea-level cave entrance through the limestone tower towards the right. Fig. 6. The doline lake in the interior of Cong Do island. A distinctive feature of Ha Long Bay is the abundance of lakes which lie inside the limestone islands. Dau Be Island, at the mouth of the bay, has six enclosed lakes, including those of Ho Ba Ham. All these island lakes occupy drowned dolines within the fengcong karst, and are known as hongs in other parts of southeast Asia (Fig. 11). Their depth profiles have not yet been surveyed; most appear to be deep, but some have shallow areas

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 over planation surfaces just below their water level, and most have marine solution notches around their walls. They are tidal, as sea water moves freely through the limestone, at some sites through sealevel caves which are traversable by boat, but at other sites through inaccessible fissure networks. One freshwater lake is known on the eastern part of Cong Do Island (Fig. 6); it probably survives on a doline floor of clastic sediment, but details of its geomorphology are unknown. Many of the areas of shallow sea between the limestone islands appear to be karst plains which have been submerged by the sea; most of the bay is less than 10 m deep. Clastic sediments cover much of the sea floor; most are probably of marine origin, though remnants of subaerial sediments from original karst plains may survive in the buried sequences. The drowned plains across the inner parts of the bay are sited where drainage from the Red River and the mainland flowed onto the limestone. The patches of fengcong karst which now form the larger islands probably originated as areas of slightly higher ground or fewer dolines in the initial surface from which the modern karst evolved. All the limestone surfaces on the Ha Long Bay islands are fretted by dissolution. The pinnacles, ridges and blades of remnant rock all have surfaces of razor-sharp rillenkarren, creating an inhospitable terrain which is very difficult to cross. Jagged open fissures continue to depth and there is no continuous soil cover, and no sub-soil rundkarren. Largely organic soil accumulates in some limestone fissures, where it provides a rooting medium for the ubiquitous scrub vegetation. All limestone surfaces are black due to the blue-gr een algae that live in the surface crust of the rock and aid its pitting by biogenic dissolution. Marine erosion of the limestone Marine invasion of the Ha Long Bay karst has added an extra element of lateral undercutting of the limestone islands. The most conspicuous feature is the main marine notch cut into most of the rocky coastlines. Its deepest zone is generally nearly 3 m high, occupying the levels between normal high and low tides. It is a complex feature, commonly with a shoulder at its mid-height and a lesser undercut extending another metre higher; the latter may represent erosion by high wave action and at spring high tides. Further notches at higher levels were cut at times of higher sea level and are no longer active. Features below low tide level have not been observed, but there is no evidence of any wide wave-cut platforms. Across the bay, there is no variation in the size of the notches which relates to exposure to the larger waves that derive from the open sea. This indicates that the notches have been cut largely by dissolution of the limestone. Sea water is normally saturated with respect to calcium carbonate, and limestone dissolution is therefore dependant on aggressive micro-environments created by algae in the surface layers of the limestone. The dark crusts with blue-green algae, that are ubiquitous on the subaerial limestone outcrops, do extend down the cliffs into the tidal range; marine algal forms are probably equally widespread, but have not yet been documented. Limestone dissolution is also enhanced at sea level by the mixing of sea water and fresh water within the fissure systems of the islands. Notches are a feature of limestone coastal cliffs worldwide, but those of Ha Long Bay are exceptionally well developed, and at many sites extend back into arches and caves (see below). Undercutting in the marine notches is presently the major process in the erosion and retreat of the limestone cliff faces. Marine erosion has not only added the notches to the profiles of the limestone islands, but it also maintains the steep faces of the fenglin karst towers, and thereby perpetuates the spectacular nature of the karst landscape (Fig. 7). Fig. 7. A small rocky island in Ha Long Bay; the remains of a fenglin tower, it is undercut by a marine notch, and a chunk of cliff has recently fallen off so that it now rests on the submerged rock platform. Many of the bay islands have narrow peninsulas formed by high limestone ridges that are bounded by cliffs; these separate and overlook bays which are much wider than the intervening land fragments, as at the norther n ends of Dau Be Island. Ridges that link chains of peaks are created in fengcong karst where there are favourable patterns

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 of large dolines. Under normal weathering conditions, these degrade into lower cols, eventually to leave isolat ed karst peaks. The very narrow limestone aretes of Ha Long Bay are not typical of subaerial fengcong karst. They are a consequence of more rapid lateral expansion of the intervening depressions, and appear to be a feature of marine erosion. Their presence indicates a significant component of marine action in the evolution of the Ha Long Bay geomorphology. The caves in the islands Within Ha Long Bay, the main accessible caves are the older passages that survive from the times when the karst was evolving through its various stages of fengcong and fenglin. There are three types of caves in the limestone islands remnants of old phreatic caves, old karstic foot caves, and marine notch caves. Remnants of old phreatic caves Many of the bay islands contain fragments of large or small cave passages that are partly choked with debris, mud or stalagmite. These are remnants of very old cave systems that were mostly phreatic. They occur at all levels in the limestone islands, and are distinguished from the other cave types by their sloping passages and considerable vertical range. Active phreatic caves must exist within the deeper limestone, but none has yet been found. Hang Sung Sot is a large fragment of very old cave passage (Fig. 8), the largest of three caves in Bo Hon Island. From the entrance chambers that are truncated at a balcony high in the cliffs, a passage more than 10 m high and wide (Fig. 9) descends gently to the south and then rises to a massive choke of boulders and stalagmites; a small exit above this emerges in a chaos of fretted limestone under a canopy of vegetation. The main cave has various levels of flowstone and false floors, and ways may exist through the boulder floors into lower passages. There is some active stalagmite growth, and one dry basin is floored with cave pearls. Daylight reaches far into the cave, but the northern aspect prevents the intrusion of direct sunlight, and there are no phytokarren. Dong Tam Cung is a large phreatic fissure cave developed in the bedding of the dipping limestone. Massive flowstone and stalactite development has divided the single fissure cave into three chambers, obliquely above each other, over a height range of about 20 m. There appears to have been only a single phase of cave enlargement, followed by draining and calcite deposition that continues today. Fig. 8. Plan and section of Sung Sot Cave in Bo Hon Island. Fig. 9. Looking into Hang Sung Sot from its balcony entrance, with the main passage extending beyond the person in the white shirt beside stalagmite on right. Fig. 10. Plan and section of Lau Dai Cave in Bo Hon Island. Dong Lau Dai is a cave with a complex of over 300 m of passages (Fig. 10). The main chamber has an entrance passage open to the cliff, while its southwestern end appears to be a choke of boulders, mud and stalagmite which is also truncated by the cliff. The passages to the north are on three levels,

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 developed on separate bedding planes and each aligned along the strike where the dip is about 10 southwest. There are thick deposits of mud and some large stalagmites. Dong Thien Cung and Hang Dau Go are two remnants of the same old cave system that both survive in the northern part of Dau Go Island at between 20 and 50 m above sea level. Thien Cung has one large chamber more than 100 m long, blocked at its ends and almost subdivided into smaller chambers by massive walls of stalactites and stalagmites. Flowstone, false floors and gour deposits cover part of its floor; the whole cave is very beautiful, and has been developed as a tourist site by building a looping path between its two fissure entrances. The roof is a spectacular series of phreatic domes, and there are younger muddy passages beneath the chamber floor. Hang Dau Go is a single large tunnel descending to a massive choke. A high-level side ch amber contains stratified sediments that have been excavated in the distant past, probably for nitrates. Dong Huang Long is a short cave entered at beach level. A chamber over 12 m high has remnants of calcite false floors high on its walls, and a soaring shaft breaks through its roof. Two passages end in chokes, and a small passage at the top of a steep gour ramp extends to the lip of a shaft. Truncated fragments of many other very old caves are exposed in the island cliffs. They occur at various altitudes up to 50 m above sea level. Some are partly filled with calcite and clastic sediments, and flowstone false floors are common; others remain as open tunnels. The sea cliff on the south side of Cong Thau Trong Island exposes a ver tical section through a breccia pipe. This consists of a column of limestone blocks which have fallen progressively from the roof of a cave, so that the void has migrated upwards, from the original cave below sea level. Old foot caves Foot caves are a ubiquitous feature of karst landscapes that have reached a stage of widespread lateral undercutting at base level. They include small notches at the foot of cliffs and also more extensive horizontal maze caves. None is active in Ha Long Bay, because their sites have been invaded by the sea to become the marine caves. There are fewer foot caves than would be expected in such a mature karst, and they are nearly all in the larger islands th at have a cover of soil and vegetation; the smaller is lands that have little soil on their rocky tops have neither foot caves nor karstic notches. This suggests that mixing corrosion is critical to the development of foot caves, and only the larger islands gather enough rainfall to provide corrosive fresh water that mixes with the salt water. On the smaller tower islands, marine notches and cliff retreat progress more rapidly than do solutional foot caves. Hang Trinh Nu is one of the larger foot caves in the bay; it extends beneath a conical hill and right through a peninsula on Bo Hon Island, between the main entrance and the bay-overlook exit that are about 80 m apart (Fig. 11). The main passages are all at the one level, with their ceilings at about 12 m above sea level. Floor heights vary, due to accumulation of stalagmite, and two large shafts rise into darkness. Notches in the walls at various levels are partly filled by stratified calcite and clay, as evidence of multiple stages in the caveÂ’s development. Fig. 11. Plan and section of Trinh Nu Cave in Bo Hon Island. An unnamed foot cave extends for 40 m right through a small island east of Cong Do (Fig. 12). The main tunnel is at its widest about 6 m above sea level; at the same level it extends into smaller side chambers and a passage which rises gently along a bedding plane. Both the latter and a section of large tube may represent earlier phreatic relics. A foot cave on an island south of Cong Do has a single, horizontal passage with an elliptical profile 5 m high and over 15 m wide. It extends 100 m to a

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 choke of flowstone and gour dams, and may continue further. Roof domes rising up to 4 m are relics of its phreatic enlargement. Hang Bo Nau is a horizontal cave about 70 m long, containing old stalagmite deposits; it is notable for the way that its passage clearly cuts across the 25 dip of the limestone bedding, and also for its exit view of the island karst (Fig. 13). Fig. 12. The old foot cave that reaches through the small island of Xac Kha, east of Cong Do. Fig. 13. The fengcong karst of Bo Hon Island silhouetted by the entrance of Bo Nau cave. Vung Ba Cua Island has a remnant of large old phreatic cave that completely breaches the narrow limestone ridge linking to its northeastern peninsula. Almost below it, a foot cave extends in 50 m at just above sea level; its roof is a calcite false floor, and there appears to be another foot cave passage at a level about 8 m higher. Thick clay in the lower passage has be en worked for pottery. Remnants of marine oyster beds survive cemented to the walls of some of the old foot caves. They date from times when the caves were temporarily invaded by high sea levels. A concentration of the oyster beds at around 6 m above sea level may correlate with terraces at that altitude around the Red River delta, which are ascribed to mid-Holocene times. Marine notch caves Since their invasion by th e sea, the karstic hills and towers of Ha Long Bay have been modified by marine erosion; the process continues today. Dissolution of the limestone creates cliff notches that may extend into caves; many of these extend right through the limestone hills, into drowned dolines which are now tidal lakes or adjacent bays. A distinguishing feature of these marine caves is a smooth, horizontal ceiling that cuts through the limestone. Many also have phreatic roof domes recessed into their otherwise flat ceilings, and it is likely that part of their development was as subaerial foot caves, with components inherited from even older phreatic caves. One of the most distinctive features of Ha Long Bay is the Ho Ba Ham group of hidden lakes and their connecting caves in Dau Be Island (Fig. 14). From the islandÂ’s perimeter cliff a cave extends about 150 m to Lake #1; the passage is 10 m wide at water level, and curves so that the centre is almost dark and requires some care when passing through in a boat. At low tide there is a minimum of 1.5 m of airspace, over water that is at least 2 m deep. The second cave, through to Lake #2, is of similar cross section but only about 60 m long. A cave through to Lake #3 is smaller, and may only be traversed by a canoe; any links to Lakes #4 and #5 have not been checked. In the two larger caves, most of the ceiling profile is a perfectly flat corrosion surface. Bulbous stalactites up to 1 m in diameter hang from the cave roofs, and largely postdate the ceiling planation. There are similar marine caves at many other sites in Ha Long Bay. Hang Luon (Fig. 15) extends 50 m through to an enclosed tidal lake. An unnamed cave extends about the same length into Lake #6 on the east side of Dau Be Island, and notch caves link adjacent bays on the west side of Vung Ha Island and on the south side of Bo Hung Island. A cave into the eastern tip of Cong Do Island carries sea water in and out of a drowned doline lake; its passage is about 4 m high and wide, and it is a very fine stream cave in clean rock, unusual only in that its flow reverses with the ebb and flow of the tide. All these caves range in width from 8 m to 20 m. They have phreatic domes in their ceilings, and parts of them are breaking upwards in fractured limestone to create low stable arch profiles. The flat ceilings in these marine caves are above present

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 high water levels, and many of them are draped with stalactites. They are clearly old erosion surfaces, created largely by dissolution in times past when sea levels were higher in the Pleistocene. Hang Luon has a massive stalactite hanging 2 m down from a flat roof, and the stalactite is truncated at the modern high tide level. Fig. 14. Outline map of Dau Be Island showing the dolines and hong lakes that lie between the conical hills of the drowned fengcong karst. Fig. 15. Hang Luon, a marine notch cave that extends through to a large hong lagoon. Evolution of the landscape Karstic evolution of the Ha Long Bay limestone area must have continued through much of Neogene time (20M to 2M years ago), as mature fengcong and fenglin landforms can only evolve through millions of years of erosion and surface lowering. In panoramic views across the bay islands, there is generally little suggestion of any concordance of summit heights that could relate to intermediate Neogene eros ion levels. Remnants of old phreatic caves survive at various levels within the bay islands, but their altitudes are not easily related to their ages. The topography of the bayÂ’s sea floor is only known from the limited available map data, but it appears that most of it has very little relief, and the great proportion lies at depths of less than 10 m. This morphology is compatable with an origin as a subaerial karst plain, which has been subsequently drowned by the sea. The major impact on Ha Long Bay of the worldwide climatic oscillations during the Pleistocene was the periodic lowering of sea level. During cooler stages, water was locked into the icecaps of higher latitudes, and sea level temporarily declined by about 100 m. This happened many times, and between these events sea levels were close to those of the present day. When the sea level was low, the whole of Ha Long Bay was dry land. Subaerial karst processes continued on the limestone basins and plains between the cone clusters and towers. Allogenic drainage from the north fed rivers across the bay area. These excavated valleys are now drowned, but none reached contemporary sea level until it was far to the south of the modern bay site. Limited mapping of the bayÂ’s sea floor indicates that, though these drowned valleys reach depths of 20-30 m below present sea level, they are discontinuous and do not continue to deepen to the south (downstream). They may represent segments of valleys that lay between caves carrying their rivers through intervening limestone ridges. Alternatively, the sea floor bedrock topography may be masked by extensive accumulations of clastic sediment. The main open part of Ha Long Bay has a floor of thick sed iment; large ships to and from Cua Luc have to follow a channel that was excavated through this, but is now kept clear by natural scour. When sea level was low during the cold stages of the Pleistocene, karstic evolution was rejuvenated, so that dolines were deepened, while the cones and towers were degraded. Many of the caves now at or just above sea level were used as sheters from 18 000 to 12 000 years ago, when sea levels were low during the Devensian glacial stage. The cave dwellers left behind vast banks of freshwater and terrestrial gastropod shells which testify to the remoteness of the contemporary sea; some of these shell banks have subsequently been cemented to the cave walls by calcite deposition.

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Tony Waltham / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 Marine undercutting of th e limestone islands at sea level is a conspicuous feature of Ha Long Bay. Its extent is more than could have been created solely in the Holocene; the many limestone ridges that have been reduced to narrow aretes by coastal retreat are indicative of a long period of marine erosion. Much of the marine morphology of the limestone islands dates from the warmer phases of the Pleistocene when sea levels were close to that of today. The complex marine cliff notches and the multiple levels of dissolution features in the associated notch caves further indicate that sea level erosion has been active over a long period of time, when sea levels have not been absolutely constant. Acknowledgements The author thanks Hans Friederich and Tran Duc Thanh for their kind hospitality and assistance during his second visit to Ha Long Bay, also the staff of the Management Department of Ha Long Bay for excellent logistical support, and the World Conservation Union (IUCN) and UNESCO for financial support. Elery Hamilton Smith is acknowledged for suggesting the role of mixing corrosion in a co-authored entry on Ha Long Bay that appeared in 2004 as pp.413-414 in the Encyclopedia of Caves and Karst Science edited by John Gunn (published by Fitzroy Dearborn, New York).



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal ISSN 1814-294X www.speleogenesis.info The role of karst in the ge nesis of sulphur deposits, Pre-Carpathian region, Ukraine Alexander B. Klimchouk Institute of Geological Sciences, Ukrainian Academy of Sciences Email: klim@speleogenesis.info Re-published from: Environmental Geology 1997, 31 (1/2), 1-20 Abstract Most of exogenous epigenetic sulphur depos its are clearly associated w ith intensely karstified carbona te and sulphate rocks. Th is paper demonstrates, using the Pre-Carpathian region as an example, that karstification is one of the most important processes g uiding the formation of sulphur deposits. This is determined by a coincidence of some ma jor prerequisites of these two processes. In the Podol'sky and Bukovinsky regions the Miocene aquifer system is well drained by erosion valleys; the giant network caves known here in gypsum formed under past artesian conditions. In the region of sulphur deposits, as sociated with the same karstif ied gypsum strata, true artesian conditions still prevail. Hydrogeologi c data show that abundant caviti es detected in the vicinity of sulphur deposits can be interpreted as having the same origin as the relict caves of the Podol 'sky and Bukovinsky regions. The widespread belief that the gyps um/anhydrite bed in the region is an aquifuge se parating the Miocene aquifers is inadequate. Thi s belief caused much controversy with regard to the genetic inte rpretations of sulphur deposits in the region. Cave systems form ed by the upward water flow through the gypsum/anh ydrite bed govern the water exchange between the aquifers within the aquifer system A new karst model for the formation of sul phur deposits is suggested. It agrees well with the hydr ogeological features of the Miocene sequence and with biogeochemical mechanisms of sulphur origin in lo w-temperature diagenetic environments. Keywords: sulphur deposits, bioepigenetic sulphur, sulpha te karst, speleogenesis, Pre-Carpatians, Western Ukraine. Introduction The Pre-Carpathian sulphur-bearing basin is one of the largest in Eurasia. In this basin all economic sulphur deposits are spatially and genetically related to the Miocene gypsum and anhydrite bed. This is one of the classic regions of exogenous epigenetic sulphur deposits in sulphate/carbonate rocks (Sokolov, 1972; Kityk, 1979). Despite more than 40 years of geologic and hydrogeologic investigations the formation of the sulphur deposits was not clearly understood. Previous models are contradictory in their interpretations of host-rock geology, ore localization, biological and geochemical sulphur-related processes, particularly in relation to the regional settings. The widespread karst in the Pre-Carpathian region is related to the sa me gypsum-anhydrite bed as the sulphur deposits. The Podol'sky and Bukovinsky areas contain vast maze caves, including the five longest gypsum caves in the world, which are now decoupled from their formational, truly confined, environment. The gypsum-bearing Miocene sequence is now largely drained through these areas. There are a lot of data on karst at sulphur deposits, in the artesian flow area, but the structure of karst systems and their role in water exchange remained little understood. Because of the studied relict caves and sulphur deposits are in different contemporary geomorphic and hydrogeologic settings, and because karst scientists have little exchange of ideas with local ore geologists and hydrogeologists, the sulphate karst in the whole region has never before been considered in a unified theory. Many investigators of exogenous epigenetic sulphur deposits, not only in the Pre-Carpathians but throughout the world, have noticed their relation to intense karstification of sulphates and carbonates, and some have mentioned a possible relationship between karst processes and the origin of sulphur. However, the causes for this relationship remained unclear. This paper considers the role of karst in the formation of sulphur deposits, with special

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.2 reference to the Pre-Carpathian region. It integrates the karst hydrogeology in the vicinity of sulphur deposits, the structure of karst systems and groundwater flow pattern, mechanisms of karst development, and biogeochemical sulphur processes. To focus on this goal, only the karst older than, and contemporaneous with, the sulphur ore is considered here. Karst and the genesis of sulphur deposits: the state of the art Having noted the increase of the thickness of sulphur ore and of the number of sulphur-bearing horizons within highly karstified areas of the Vodinsky deposit in the Volga region, Russia, Borodaev (1936) and Markov (1954) suggested that karst processes not only destroy but also form ore bodies (as cited by Otreshko, 1966b). Peresun'ko (1960) noted the spatial coincidence of ore bodies and karst in sulphur depos its in Central Asia and the Pre-Carpathians. In his opinion (p.135) karst phenomena are inextricably related to the sulphurforming processes. In other work he noted that interpretation of karst in sulphur deposits was limited by a poor understanding of their hydrogeological settings, but that a study of the genesis, character and intensity of karst around sulphur deposits is of a great practical value (Peresun'ko, 1961). Babine ts and Tsapenko (1960) indicated that the belt of deposits of mineral water bearing H2S in the Pre-Carpathians is associated with karst processes in Miocene gypsum rocks, but they did not discuss the nature of this relationship. The association between the sulphur formation and karst in sulphate/carbonate strata of the VolgaKama region had been noticed by Stankevich (1968). Tkachouk and Koltun (1963) stated that a detailed investigation of subsurface karst in the gypsum/anhydrite of the Pre-Carpathians was needed for prospecting and exploitation of the sulphur. Pertzovich (1969) stressed that the determination of the age of karst in gypsum/anhydrite and carbonate rocks in the PreCarpathians was important to solving the problem of the genesis of native sulphur. Polkunov et al. (1979) studied the details of karst in the vicinity of sulphur deposits of the Pre-Carpathians (see below), and pointed out that karst features are important aspects of ore fields and deposits. The pronounced role of karst in the formation and localization of bioepigenetic calcite and sulphur mineralization has been shown for deposits in the Delaware basin of west Texas (Wallace and Crawford, 1992; Miller, 1992). These minerals are formed within the thick late Permian evaporites of the Castile and Salado Formations and are associated with large caves filled with late Cretaceous conglomerates and karst breccias. Paleosystems of cave channels provided permeability within the gypsum/anhydrite and served as migration paths for hydrocarbon-bearing water into the zone of sulphate reduction from the underlying aquifer, the Bell Canyon Formation. A clear statement of the role of karst in the formation of sulphur deposits was given by Otreshko (1966b): Karst, most likely, must be included among the major factors determining the genesis, transformation, and destruction of deposits ". In another work he suggested that the presence of karst is one of the basic guidelines for prospecting for sulphur deposits (Otreshko, 1966a). Otreshko (1966a, 1966b, 1974) discussed a paleogeographic regularity which had been outlined for the first time by Teodorovich (1943) and Sokolov (1956): sulphur deposits are associated with areas where the clay caprock overlaying sulphate/carbonate sequences is considerably scoured by erosional action Numerous examples of sulphur deposits in the Pre-Carpathians, Volga region and Central Asia, show that ores are localized beneath, or adjacent to, late Neogenic buried valleys. In the Pre-Carpathians in particular, there is a distinct inverse relationship between the thickness of sulphur-bearing and barren limestones overlying the gypsum and the thickness of the overlying clay caprock (Otreshko, 1974). The reasons why epigenetic limestones and native sulphur are associated with erosionally scoured structures remained unclear. Otreshko (1966a) also suggested that the association between karst and sulphur ores is determined by that ores occur beneath erosionally scoured areas exposed to descending percolation and, hence, they are more susceptible to destruction by karst processes. In this case the post-ore karst is meant that apparently contrasts with the above cited statement by Otreshko on the ore-forming role of karst. Pisarchik and Rusetskaja (1972) considered the sulphur deposits of the Pre-Carpathians to have formed during the hypergene transformation of sulphate rocks. They conc luded that the history of drainage and karst devel opment, as well as the neotectonics that controls them, are of great significance to hypergenic processes, and particularly to the origin of sulphur. Osmolski (1976) examined the role of karst in the gypsum/anhydrite stratum in the formation of sulphur deposits in Poland. The drawn an important conclusion th at karstification of

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.3 gypsum/anhydrite had facilita ted water circulation, leading to the formation of an aquifer and, in the process, opening migration paths for hydrocarbons. He believed that the processes, leading to replacement of sulphates by calcite and sulphur could have been limited to the gypsum strata exposed to two phases of karstification in preSarmatian time. He invoked the traditional concept of karst development by descending water, an idea that is not supported by data from the Ukrainian Pre-Carpathians. One of the major phases in the formation of sulphur deposits is considered by Gajdin (1983) to be the development of sul phate karst which result in the opening of migration paths for hydrocarbons from the sub-gypsum aquifer into the sulphates. He also pointed out that sulphur ores form zonally, extending from the underlying aquiferous rocks and karst cavities in the sulphates. Most studies of karst in the vicinity of sulphur deposits have been merely descriptive and have not clarified the genesis of karst, its hydrogeological and ore-forming role. Only in the above-cited works of Gajdin, Osmolsky, Wallace and Crawford and Miller, has the role of karst been outlined in the context of development of water-exchanging paths and bringing sulphates in contact with hydrocarbons. A general interpretation of the conjunction of karst development and nati ve sulphur formation can be derived from the fact that both processes share many basic prerequisites. Formation of bioepigenetic sulphur requires the presence of: (1) sulphate rocks, (2) hydrocarbons, (3) oxidizing agents, and (4) conditions favourable to the interaction of reactants, namely: (4a) transporting solutions, (4b) permeability of the host rocks, and (4c) hydrodynamic conditions necessary to cause the movement of solutions. The above conditions are sufficient only if certain favourable physical and chemical environments are brought in interaction over sufficient time. It is important to note that conditions 1, 4a, 4b and 4c are the basic requirements for karst development as outlined by Sokolov (1962). Thus, as a rule, karstification must occur for sulphur deposits to form Moreover, formation of an extensiv e sulphur deposit requires large-scale sulphate reduction, which can be maintained only in th e presence of sufficient dissolution surface of sulphate. Secondary permeability and subsurface dissolution surfaces are created by speleogenesis. Geological settings of the Pre-Carpathian sulphur-bearing basin The Pre-Carpathian sulphur-bearing basin lies within Miocene gypsum/anhydrite, and all economic sulphur deposits are related to these rock both spatially and genetically. The sulphur is located in the transitional zone between the Eastern-European and Western-European platforms and the Carpathian foredeep (Fig. 1; Kityk, 1979). Upper Badenian sulphate rocks stretch sub-parallel to the Carpathian folded region through southern Poland, western Ukraine, and north of Romania. Further discussion in this paper is limited to the Ukrainian part of the basin (Fig. 2). Economically viable sulphur deposits adjoin the foredeep edge of the platform. The gypsum/anhydrite "belt" is narrowest in the northwest section and extends southeast for more than 300 km through Ukraine. In it's southeast part, this belt widens up to 100 km encompassing some interiors areas of the platform (Podol'sky area). Miocene sedimentary rock s within the platform fringe overlie the eroded strata of Cretaceous or, less commonly, earlier age. The Cretaceous succession is represented by terrigenous and carbonate sediments, mostly by marls, sandstones, detrital, and argillaceous limestones. The Miocene succession consists of Badenian and Sarmatian deposits. The Lower Badenian, beneath the gypsum, includes mainly carbonaceous, argillaceous, and sandy deposits near the foredeep (70-90 m thick), that grade into calcareous bioherm and sandy facies toward the interior of the platform (10-30 m thick). In the vicinity of most sulphur deposits, the Lower Badenian is composed mainly of lithothamnion limestones. The gypsum/anhydrite stratum (GAS) and overlying pelitic and crystalline limestones comprise the Tyrassky Formation of Upper Badenian age. The sulphate stratum, 35-50 m thick, consist of gypsum in the interior of the platform (Podol'sky area); but toward the foredeep the anhydrite content increases. The present extent of sulphates on the platform is as much as 20,000 km2. The GAS is rather variable in structure and texture. In the Podol'sky area, it consists of threeunits, which, in ascending order, include cryptoand microcrystalline massive gypsum, bedded microcrystalline and megacrystalline gypsum. In the upper unit, the gypsum displays large spherules (Klimchouk et al., 1995). Toward the foredeep, the GAS becomes more homogeneous and aphanitic. Interbeds of carbonate and clay are rare and thin.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.4 Fig. 1. Distribution of sulphur, gas and oil deposits in relation to major tectonic structures of the Pre-Carpathians an d Carpathians (from Kityk, 1979). 1 = platforms. Carpathian foredeep : 2 = outer zone, 3 = inner zone. Carpathian folded region : 4-10 = tectonic zones, 11 = periclinal depression of eastern Carpathians, 12 = Vienna Trough; 13 = effusive formations, 14 = sulphur ore deposits, 16 = gas reservoirs, 17 = oil reservoirs. Fig. 2. Location of gypsum stratum, sulphur and hydrocarbon deposits, and large caves of the western Ukraine (modified from Polkunov et.al., 1990). 1 = Eastern-European platform fringe. Carpathian foredeep : 2 = outer zone, 3 = inner zone. 4 = Carpathian folded region. 5 = sulphate rocks in the platform. Tectonic boundaries : 6 = platform/foredeep, 7 = outer/inner zone of the foredeep, 8 = foredeep/folded region. 9 = other major faults. 10 = flexures. 11 = zones of sulphur mineralization. 12 = sulphur ore deposits. 13 = gas reservoirs. 14 = oil reservoirs. 15 = large maze caves in the gypsum.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.5 The layer of pelitic and crypto-crystalline limestones, ranging from several dozens cm to more than 25 m in thickness, normally overlies the GAS. Its origin, which is closely related with that of sulphur, is still controversial (Grinenko et al., 1966; Mamchur, 1972; Sakseev, 1972; Lein et al., 1977). The upper limestones contain two genetic varieties. The pelitic limestones (locally called "Ratynsky") are evaporitic in origin. They are 0.210 m thick and occur over almost the entire GAS area. In places they contain minor veins or pockets of sulphur within fissures and caverns. The other variety of limestone, which is cryptoand microcrystalline, formed epigenetically by metasomatic or hydrogenic replacement of the GAS during sulphate reduction. This limestone varies in thickness up to 25 m and is prevalent in the areas of sulphur deposits. In places it has completely replaced the GAS and is either sulphur-bearing or barren. There is a prominent inverse relationship between the thickness of the crystalline limestones and overlying clay cover (Fig. 3; Otreshko, 1974): the epigenetic limestone is thickest where the clay cover has been partially eroded. In places, epigenetic limestones are found also beneath the GAS (Sakseev, 1966; Composition, 1979). The Ratynsky and epigenetic limestones differ greatly in carbon isotope composition. Values of C13 vary in the Ratynsky limestone from -8 to -3 ‰, whereas the epigenetic limestone gives values ranging from -32 to -65 ‰ (Fig. 4; Mamchur, 1972; Lein et al., 1977). The isotopically light carbon comes from CO2 generated by oxidation of methane during sulphate reduction, which then contributes to epigenetic calcite deposition. The sedimentary Ratynsky limestones were recrystallised in places and enriched with epigenetic calcite. In these situations they range in C13 from -11 to -57 ‰. Considering that it is difficult to distinguish between the two varieties of limestone in the field, they will be collectively referred to as "supragypsum limestones" where their genesis is unknown or not relevant. The contact of these limestones with the underlying GAS varies from gradual to sharp, depending on the genesis of the limestone and the extent of hypergene alterations. The Tyrassky Formation is overlain by argillaceous strata of the Kosovsky Formation, which is also of late Badenian age. Near the foredeep boundary these strata are mainly clays, with sandstone and carbonates in the lower part. In the interior of the platform, red algal argillaceous limestones with minor beds of sandstones prevail. The Upper Badenian strata grade upward into the Sarmatian marls and clays, whose thickness increases to 30-50 m toward the foredeep. Adjacent to the foredeep, Sarmatian strata are lithologically similar to the Kosovsky Formation. Together they constitute a clay overburden over sulphur deposits with a total thickness up to 80-100 m. Fig 3. Relationship between the thickness of the supragypsum limestones and the Kosovsky Formation (caprock) at the Rozdol'sky sulphur deposit (from Otreshko, 1974). Fig. 4. Variation in isotopic composition of carbon from the Ratynsky pelitic limestone (at top) to the crystalline epigenetic limestone (at bottom) within the supragypsum aquifer of the Tyrassky Formation in the PreCarpathian (from Mamchur, 1972).

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.6 The Miocene formations are overlain by late Pliocene and Quaternary glacio-fluvial sands and loams (the north-west section of the gypsum/ anhydrite belt was exposed to the Mindel glaciation; Gerenchuk et al., 1972), and by high alluvial terraces of sand a nd gravel deposited in the Podol'sky area by Dniester River during the late Pliocene and early Pleistocene. There are many buried valleys entrenched 30-50 m into the Kosovsky and Sarmatian clays and, in places, into the upper part of the Tyrassky Formation (Figs 57). Some of these valleys are considered to be of Mindel-Riss age, others to be Middle Pleistocene (Kutepov and Tsjurupa, 1969; Kutepov and Kozhevnikova, 1989; Sokolov et al., 1969). Fig. 5. Location of buried valleys at the Jazovsky sulphur deposit (modified from Kutepov and Kozhevnikova, 1989). 1 = elevation contours at top of Miocene strata; 2 = buried valleys; 3 = surface karst features; 4 = zone of high hydraulic conductivity in the Miocene aquifers, from due tracing tests; 5 = regional fault zone; 6 = other faults. Fig. 6. Cross section of a buried valley at the Rozdol'sky sulphur deposit (from Sokolov et al., 1969). Quaternary sediments : 1 = soil; 2 = loam; 3 = sand; 4 = sand and gravel. Miocene sediments : 5 = clay; 6 = limestone; 7 = sandstone; 8 = brecciated limestone; 9 = gypsum and anhydrite. 10 = fissures. Fig. 7. Location of the ancient terraces of the Dniester River and other relict valleys and large caves in the Podol'sky area. Northern boundary of alluvium of the 7th terrace: 1 = according to Svyn ko (1979); 2 = according to Gofshtain (1962); 3 = relict valleys (Svynko, 1979); 4 = large caves in gypsum.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.7 The present distribution of Miocene formations and levels of their denu dation vary regularly from the interior of the platform toward the foredeep (Andrejchouk, 1988; Klimchouk, Andrejchouk, 1988). The Tyrassky Formation dips 1-3o toward the foredeep and in the transitional zone is disrupted by block faults. Concurrently, the thickness of the Kosovsky and Sarmatian clay overburden increases, and the depth of erosional entrenchment decreases (Figs 8 and 9). These differences, the result of differential neotectonic movement, played the most important role in the hydrogeologic evolution of the Miocene aquifer system, as they determined the recharge-discharge and flow conditions, partic ularly the development of karst in the GAS and the formation of sulphur deposits. Hydrogeology of the Miocene sequence The Pre-Carpathian sulphur-bearing basin occupies the southwestern part of the VolynoPodol'sky artesian basin (Babinets, 1961; Kityk, 1989; Shestopalov, 1981), and the second order Podol'sky and Bukovinsky drainage basins. The GAS spreads on an artesian monocline slope in which the regional flow is toward the southwest and south, from the interior of the platform toward the large Dniester valley and the Carpathian foredeep. On the south-west, along the fault boundary with the Carpathian foredeep, the Miocene and Cretaceous aquifers are brought into lateral contact with thick low-permeable clays of the Kosovsky formation. This causes an upward discharge of the aquifer system through the capping clays. The same situation takes place also along every major block fault of the platform fringe. Fig. 8. Hydrogeologic profile across the gypsum/anhydrite belt near the Jazovsky sulphur deposit (line I-I' on Fig. 2) depicting natural (pre-quarrying) conditions. Geolog y based on the profile from Aleksenko (1967). Fig. 9. Geologic/hydrogeologic settings of gypsum karst development in the Podol'sky and Bukovinsky areas in the southeastern section of the gypsum belt (line II-II' on the Fig. 2).

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.8 Present groundwater flow in the Miocene sequence is controlled by the geologic and geomorphic position of the individual strata. As noted above, these conditions alter regularly in the direction from interior of the platform toward the foredeep (Figs 8 and 9). North-northeast of the gypsum belt the sub-gypsum aquifer is exposed to the surface or lies beneath permeable Quaternary sediments. This is the main area of infiltration recharge (Rostochje, Lvov Plateau, Opolje, the northern part of the Podol'sky area). To the southsouthwest (particularly in the Podol'sky area) the gypsum and capping clays remain not eroded, subdivided into isolated uplands by deeply entrenched left tributary valleys of the Dniester. The water table lies within the sub-gypsum formation, or, in places, within the lower part of the gypsum. The aquifer receives point recharge through karst sinkholes, and discharges in valleys via springs. Further to the south-south-west, along the foredeep boundary, the gypsum stratum is almost intact. The aquifer system in the Miocene sequence is confined by the thick clays. Groundwater discharges upward along buried valleys and faults that breach the clay. The major regional fault that brings Miocene aquifers into lateral contact with the thick clay prevents further flow toward the Carpathian foredeep. Paleohydrogeologic evolution and the formation of the above variations in the hydrogeologic conditions have been caused by differential neotectonic movements during late Neogene Pleistocene, resulted in plunge of the platform fringe and compensating accumulation of the clay series, uplift of the interior part of the platform with accompanying denudation of clays and the gypsum, deep entrenchment of the Dniester valley and it's left tributaries. Present hydrogeologic settings of the Miocene sequence also differ considerably between the northwest (narrow) and southeast (wide) sections of the gypsum belt. The northwest section of the GAS belt In this region, the monoclinal slope is most clearly developed (Fig. 8), although it is complicated by block faulting. Most of the sulphur deposits are located there. Within the uplands of Rostochje and Opolje, beyond the GAS limit, Lower Badenian strata lie just beneath the Quaternary cover on Cretaceous rocks of low permeability. In places, the Ratynsky limestones are also present. This is the main area of infiltration into the regional aquifer. In this area the aquifer is unconfined and contains HCO3-Ca waters with TDS as much as 0.5 g/L. Beneath the Upper Miocene clays, the heads are high enough that some water flow upwar d into sandstone beds of the lower Kosovsky Formation. The Kosovsky and Sarmatian formations consist mainly of clays that constitute the cover of low permeability separating the Miocene and Quaternary aquifers. The hydraulic conductivity of clays does not exceed 10-6 cm/s, although it increases to as much as 10-3 cm/s in minor sandstone and siltstone beds in the lower part of the sequence. The permeability of the Kosovsky formation is greatest along faults, and buried valleys. In the confined flow area, the GAS is absent only in places. Local upward discharge is concentrated along the Dniester valley (which is rather shallow in this section), as well as along buried valleys, where the thickness of the clay overburden diminishes. Along the foredeep, where regional faulting has brought the Miocene sequence into lateral contact with the thick Kosovsky clays, further flow in this direction is prevented. In the confined flow area two aquifers are commonly distinguished in the Miocene sequence: (1) the "sub-gypsum" aquifer in the Lower Badenian lithothamnion limestones, sands and sandstones, and (2) the "supra-gypsum" aquifer in the Ratynsky and epigenetic limestones and lower part of the Kosovsky Formation. The sub-gypsum aquifer is the major aquifer in the system. In the few areas where the gypsum is absent, the aquifer lies directly beneath the Upper Badenian strata. The aquifer contains rather homogeneous granular and fracture porosity and has hydraulic conductivity of 1.0x10-4 to 1.6x10-3 cm/s, but in the lithothamnion limestones there are karstified zones with hydraulic conductivity as high as 6x10-3 to 7x10-1 cm/s (Fedorova, 1985). Waters in the aquifer have calcium bicarbonate composition with TDS as high as 1.0 g/L. In places there are calcium sulphate waters with TDS up to 1.1 2.9 g/L and SO4 = content up to 0.8 1.7 g/L. Some waters contain H2S at concentrations as high as 20-30 mg/L (Fedorova, 1985) with rare values up to 115 mg/L (Tsjurupa, 1960). In local areas ClNa methane-bearing waters flow upward along faults from the underlying Cretaceous rocks and adjacent foredeep with TDS up to 7.5 g/L (Babinets and Tsapenko, 1960). These waters are concentrated in a narrow zone adjacent to the foredeep boundary where th e gypsum is absent and the sub-gypsum and supra-gypsum aquifers constitute a single hydrologic unit (Fedorova, 1985). The methane content in released gases

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.9 reaches 92%; hydrocarbon gas shows have been observed at many sulphur deposits (Aleksenko, 1967; Srebrodol'sky and Kachkovsky, 1973). Methane is believed to be the main source of organic carbon for sulphate reduction in the Miocene sequence. The supra-gypsum aquifer is 1 to 25 m thick with a highly varied but generally high porosity consisting of fractures, solution voids and conduits. Hydraulic conductivity normally ranges from 0.02 to 0.1 cm/s, with a maximum of 0.5 in karstified zones (Polkunov et al., 1979), which are mostly developed in the lower part of the aquifer, near the contact with the underlying gypsum. It is important to note that areas of the high hydraulic conductivity of the sub-gypsum lithothamnion limestones coincide with those in the supra-gypsum limestones (Babinets and Tsapenko, 1960). Waters are rich in calcium sulphate, or locally a calcium-sodium sulphate, with TDS of 1.0 3.6 g/L and SO4 = content of 1.5 2.0 g/L. High H2S (34 200 mg/L; up to 370 mg/L in places) and CO2 (120 to 167 mg/L) indicate sulphate reduction processes intensely operating within the aquifer (Babinets and Tsapenko, 1960; Tsjurupa, 1960). It is a widely believed that the GAS is an aquifuge that separates the sub-gypsum and supragypsum aquifers (Peres un'ko, 1960; Tsjurupa, 1960; Goleva, 1962; Aleksenko, 1967; Koltun et al., 1972; Polkunov et al., 1979; Fedorova, 1985,1986; Kushnir, 1988; Kutepov and Kozhevnikova, 1989 and others). Most of these works recognise that hydraulic connection between the two aquifers takes place only in "windows" where the GAS is absent. Only a few authors (Babinets and Tsapenko, 1960; But, 1962; Gajdin, 1983, Fedorova, 1985) referred to the GAS as a water-bearing stratum. Babinets and Tsapenko (1960) emphasise that the Miocene sequence consist of a single aquifer, but But and Fedorova describe zones of enhanced hydraulic conductivity (7.3x10-2 cm/s) in the gypsum/anhydrite that connect the aquifers above and below. According to Kutepov and Kozhevnikova (1989), the GAS can contain water in tectonically disrupted zones, where hydraulic conductivity can reach 3.8x10-2 cm/s. Only Gajdin (1983) clearly stated that the GAS is an aquifer in the vicinity of Pre-Carpathian sulphur deposits. The thorough analysis of available field data allows to argue that it is a misconception to threat the GAS as an aquifuge and a separating bed. This has led to misinterpretation of the water circulation pattern in the Miocene sequence, considerable inaccuracy of predictions of water inflow to opencut mines (Rozdol'sky a nd Jazovsky deposits), realisation of ineffective projects to prevent water inflow to quarries, but also to misinterpretation of the genesis of deposits. The latter will be discussed in the following chapters. The GAS prevents hydraulic connection between the two aquifers only in limited areas of some poorly fractured nonkarstified blocks. In karstified areas, often rather extensive, the GAS is highly permeable and allows both vertical and lateral groundwater circulation. Most authors considered the GAS to be an aquifuge a priori Evidence, if given at all, included references to monolithic cores from boreholes, contrasting water chemistry between the aquifers, and discontinuities in potentiometric levels. However, monolithic cores are commonly obtained in strata with high conduit permeability. Different chemical composition of water above and below the gypsum simply reflects differences in lithologies, recharge pattern, and flow conditions. Finally, field data show insignificant head differences between the sub-gypsum and supr a-gypsum aquifers. For example, at the Jazovsky deposit, under modern disturbed conditions, this difference is only 0.2-0.5 m; it is locally 1.5-2.0 m, but these data are from boreholes located several hundreds meters apart. Numerical modelling of area of the Jazovsky deposit has shown that a head difference of 10 m between the aquifers would require a zero hydraulic conductivity for the dividing GAS bed over an area of 35-40 km2 (A. Ishchouk, personal communication), which is not feasible. Additionally, data obtained from the Jazovsky deposit during the pre-quarrying stage showed that considerable variations of a head (up to 10-20 m) in the supra-gypsum aquifer simply reflected the highly heterogeneous nature of karst permeability in the aquifer system (Fedorova, 1985). The present author, along with A.Aksem, has conducted groundwater tracing in the Miocene sequence at the Jazovsky deposit for many years (30 injections, with tracers interception via boreholes). These experiments have shown welldeveloped fissure and conduit permeability in the GAS and a close hydraulic connection between karst waters in the GAS and those in the underlying aquifer. Tracers injected into the lower aquifer were frequently detected in the GAS, and vice versa. Regrettably, no active wells in the supragypsum aquifer were available during our tests. However, Fedorova (1985) reported that heads in two wells at the Jazovsky deposit, one completed in the GAS and the other in the sub-gypsum aquifer, immediately reacted to pumping of a third well completed in the supra-gypsum aquifer.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.10 A close connection between waters of the suband supra-gypsum aquifers is also shown by the water flow and chemistry measured during open-pit mining of the sulphur deposits. For example, it had been assumed for most deposits that major inflow to quarries would be a function of only the specific storage in the supra-gypsum aquifer (Tsjurupa, 1960). But during exploitation of the Rozdol'sky and Jazovsky deposits, consid erable and gradually increasing inflow occurred from the sub-gypsum aquifer and necessitated considerably greater pumping of water than predicted, which made it difficult to achieve the projected level of lowering. During the first years of water withdrawal, the sulphate content of the water in the supra-gypsum aquifer decreased, while the bicarbonate content increased. TDS decreased to 1.6 g/L, and H2S content rose to 2-70 mg/L (Peresun'ko, 1960). This demonstrates considerable increase of inflow from the sub-gypsum aquifer. Later on these changes have become even more pronounced because of induced recharge from the Quaternary aquifer, piracy of surface runoff, and activation of karst collapse (Gajdin et.al., 1991; Andrejchouk and Klimchouk, 1993). Many researchers explain the local variations of water chemistry in the supra-gypsum aquifer as the result of inflow from the sub-gypsum aquifer through "windows" in the GAS. Kushnir (1988) analyzed the water chemistry of wells in the vicinity of the Jazovsky sulphur deposit, comparing those where the GAS is present with those where it is absent. Where the "gypsum-anhydrite aquifuge" (Kushnir's term) is present, mean Ca++, Cl-, SO4 =, pH, and TDS are lower than those in "windows" areas, and the minimum values are considerably lower. Paradoxically, he considers this to be the result of greater inflow of fresh water from the subgypsum aquifer along fau lts that conduct water through the sulphates. In fact, such inflow occurs through cave systems in the GAS, as shown in the following section. All these interpretive problems about the GAS are remarkable examples of misconceptions resulted from overlooking of karst systems and poor understanding of their hydrogeological function, by geologists and hydrogeologists. The southeastern section of the GAS belt In this region the belt widens up to 80-100 km and extends far into the platform interior along the northern side of the Dniester valley (Fig. 9). A monoclinal artesian setting and slow water circulation within the Miocene strata have prevailed there since the early Pleistocene. Artesian flow to the southwest of the recharge area has been negligible due to the continuous spread of the Upper Miocene clay cover, which limits the discharge of water. Entrenchment of the protoDniester valley and its left tributaries during late Pliocene early Pleistocene allowed upward discharge into the valley bottoms, which activated artesian flow. Later deepening of the valleys breached the artesian confinement and allowed drainage of the Miocene stata (entrenched karst zone on Fig. 9). In the Podol'sky area the water table presently lies within the sub-gypsum strata in interfluve massifs. This aquifer receives point recharge through karst systems where the gypsum and clay coverbeds are present, and diffuse recharge where the clay has been removed (Klimchouk et al., 1985). Draining of the gypsum made accessible vast relict maze caves within it (Klimchouk, Andrejchouk, 1988; Klimchouk, 1992). The setting is similar also immediately south of the Dniester valley. In the Dniester-Prut interfluve area (Bukovinsky area) the GAS plunge conspicuously toward the foredeep along fault blocks. The Upper Miocene clays thicken in that direction, and the depth of erosional entrenchment diminishes. As a result, the water table is located within the GAS in the 3-15 km-wide belt that extends sub-parallel to the Dniester and Pre-Carpathian foredeep (subjacent karst zone, Fig. 9). Point recharge is favoured in this zone because of intense solution of the gypsum stratum and the presence of the clay cap. Groundwater flows north to the Dniester valley and south-southwest to the Prut valley. Flow becomes confined in the latter direction, near the foredeep, where the Tyrassky Formation lies at even greater depth (deep-seated karst zone, Fig. 9). Water from the confined aquifer discharges upward to the Prut valley. In some of the most uplifted blocks this valley has intersected the top of the Tyrassky Formation and has breached the artesian confinement. In the Kryvsky block this breaching took place during the final stage of the natural flow regime, when the water table established near the top of the GAS in the vicinity of the valley. Dewatering of the gypsum quarry since 1950s has lowered the water table, exposing the extensive Zolushka Cave, in which 92 km of passages have been mapped (Fig 10-E; Andrejchouk, 1988). This cave is representative of solution cavities within the present-day areas of confined flow.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.11 Karst in the Tyrassky Formation Karst in the sulphate/carbonate rocks of the Tyrassky Formation around sulphur deposits (in the artesian flow area) has been noticed by many investigators, although mainly in the Ratynsky (supra-gypsum) limestones. This is because close attention has been paid to this productive horizon and because its better av ailability for study in quarries, outcrops, and borehole cores. It is more difficult to study karst in the sulphates in the modern deep-seated (confined) karst zone. On the other hand, the well studied gypsum karst of the Podol'sky and Bukovinsky areas had long been considered the result of descending water circulation in unconfined hydrogeologic conditions similar to those of today; the concept which was clearly impossible to apply to the karst in the zone of artesian flow in the vicinity of the sulphur deposits. Podol'sky and Bukovinsky areas These areas contain the five longest gypsum caves in the world (Optimistichna 214 km, Ozernaya 117 km, Zolushka 92 km, Mlynki 28 km, Kristal'naya 24 km), as well as many smaller caves. They display rather uniform maze-like patterns (reticulate or polygonal network of passages) with a passage density up to 320 km km2 (see examples in Fig. 10). They are 2-4 storied systems that occupy areas of up to 2 km2 each. It was believed until recently that these caves were formed by water from sinking streams flowing laterally through the gypsum to nearby valleys, or by flow between entrenched valleys (Dubljansky and Smol'nikov, 1969; Dubljansky and Lomaev, 1980). This idea implied unconfined hydrogeological conditions and did not shed light on the origin of cavities encountered in drill holes in the artesian parts of the GAS, such as at the sulphur deposits. Recently a hypothesis of cave origin in artesian conditions has been developed, with special reference to the gypsum caves of the Podol'sky and Bukovinsky areas (Klimchouk, 1990; 1992; 1994). It was suggested that maze caves in the gypsum originated in a shallow, multi-storey artesian system as the result of upward water flow between aquifers, especially where discharge through the confining caprock was most concentrated, such as beneath large valleys. Solution by dispersed recharge from the underlying aquifer formed network mazes in the gypsum because enlargement of available fissures was rather uniform. Cave development was most active at areas of potentiometric low, where the confining caprock was thin due to erosional en trenchment, or (and) its permeability was enhanced by the presence of fault zones (Fig. 11). The multi-storey structure of caves was determined by a multi-storey occurrence of fissure networks, which occupy certain lithologic and structural intervals of the gypsum bed (Klimchouk et al., 1995). The pattern of fissures in networks does not coincide between adjacent horizons, allowing rather independent development of maze storeys at various levels. It is interesting to note that before speleogenesis evolved, the gypsum actually served as an aquifuge separating strata of higher permeability. As conduits enlarged, the gypsum gradually became a karst aquifer with inhomogeneous hydraulic conductivity that is very high in areas where maze caves developed. Fig. 10. Maps of some of the major gypsum maze caves in the Western Ukraine.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.12 Fig. 11. Conceptual model of artesian speleogenesis in the Western Ukraine. Deep river entrenchment in the Podol'sky area during the middle and late Pleistocene left the caves in a relict state above the piezometric surface. Only minor modification of cave morphology takes place today in areas of local sinking surface streams, and in the lower part of the gypsum where the water table is still above its bottom in some interfluve massifs. Within the slowly submerged part of the platform fringe (Bukovinsky area) the gypsum is still largely or entirely inundated, except in areas of artificial dewatering. Discovery of 92 km-long Zolushka cave by dewatering (Fig. 10-E) illustrates that cave origin under artesian conditions has proceeded until recently. Thus, the caves in the Podol'sky area have developed in largely the same conditions as those present today in the belt adjoining the foredeep, where the sulphur deposits are located. The data on known relict caves and the concept of artesian speleogenesis, described above, can be applied to the interpretation of artesian karst in the vicinity of the sulphur deposits. Northeast section of the GAS beltthe area of sulphur deposits Fedorova (1985, 1986) suggested that karstified and permeable areas in the GAS at the Jazovsky deposit are located where the gypsum is thinnest, especially below buried valleys. Kutepov and Kozhevnikova (1989) state that clay-filled and water-filled solutional cavities in the GAS at the Jazovsky deposit are most abundant where the GAS is thinnest, but that intense karstification is also present in places where the unit is rather thick. Drill holes in one area where the gypsum is 48 m thick, encountered 4-6 m high cavities about half filled with clay. Kutepov and Kozhevnikova (1989) suggest that dissolution is performed by ascending waters. Intense karstification of both the GAS and sulphur-bearing limestones was also mentioned by Bobrovnik and Golovchenko (1969), Merl ich and Dacenko (1976), Pertzovich (1969). Polkunov et al. (1979) characterize the GAS as an aquiclude, but they nevertheless present data about karst features within it and state that, although cavities occur throughout, they are most concentrated at the top and bottom of the GAS. Cavities encountered in drill holes have a maximum vertical range as much as 7.1 m at the Jazovsky deposit and 9.7 m at the Nemirovsky deposit. The author has analysed data from nearly 100 wells in the GAS at the Jazovsky sulphur deposit and at the Nikolaevsky clay quarry. Solution cavities were encountered in 23% of the wells (in places as high as 35%), most of them open (water filled). Cavities had the greatest vertical dimensions at the Jazovsky deposit (up to 12 m).

PAGE 13

Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.13 Our survey supports the observation of Polkunov et al. that cavities occur mainly in the lower and upper parts of the GAS. This corresponds well with the multi-storey structure of relict maze caves known in the PodolÂ’sky and Bukovinsky areas. At the Jazovsky deposit, tracers injected into the subgypsum aquifer were frequen tly detected at wells in the GAS, whether or not they intersect cavities. This indicates the upward pattern of groundwater flow through the gypsum/anhydrite. The characteristics of karst in the GAS in the vicinity of the Pre-Carpathian sulphur deposits are compatible with the pattern and secondary filling of known cave systems in the Podol'sky and Bukovinsky areas and with the artesian model of speleogenesis mentione d above. This model readily explains the geologic, hydrologic and hydrochemical peculiarities of the supra-gypsum and sub-gypsum aquifers and their relationships. Thus it is appropriate to use this speleogenetic model to explain karst development in the GAS in the vicinity of sulphur deposits. There is much more data published about karst features in the supra-gypsum limestones, the production horizon of sulphur. All researchers note the intense heterogeneous karstification of the limestone in the form of vugs and caverns up to 2030 mm in diameter. The presence of larger cavities is also frequently noted. Tkachouk and Koltun (1963) refer to extensive development of large cavities, similar to those in the GAS, at sulphur deposits and other areas where the supra-gypsum limestones are thickest. A large cave in sulphur ore at the Tcharkowy deposit (Poland) was described by Push (cited in Polkunov et al., 1979). The large Medovaya Cave, of clearly phreatic morphology, is known in the presently drained part of the Ratynsky limestones at Lvov. Polkunov et al. (1979) report intersection of cavities 3-5 m high in drill holes in certain areas of the Jazovsky deposit, where the GAS is absent and the sulphur-bearing limestones rest directly on lithothamnion limestones. This is explained as a "special case" of karst development below local base level due to action of water discharging under pressure from the lithothamnion limestones where these limestones are directly overlain by sulphur ores. These authors point out that at the Rozdol'sky deposit there is intense karstification of sulphurbearing and barren limestones throughout all areas where they directly overlie the lithothamnion bed, with karst commonly developed at two distinct levels (at both the bottom and top of the limestones). Collapse cavities are characteristic for the supragypsum limestones at all sulphur deposits. For example, during the geological mapping of mines, 247 collapse spots were noted within an area of 1.5 km2 of exposed limestones at the Rozdol'sky deposit. Karst collapses form continuous zones several hundreds meters wide extending 2-3 km (Polkunov et al., 1979). Similar features are also mentioned in the Jazovsky, Nemirovsky, Ljubensky, Zhydachevsky deposits, as well as for the Tarnobzeg deposit in Poland. When interpreting karst in the Tyrassky Formation, either in the GAS or in the supragypsum limestones, almost all researchers have invoked the traditional concept of a "descending" karst, which (as described by Polkunov et al., 1979) develops in the vadose zone, and also in phreatic zone where drainage to local rivers is possible. Such conditions are absent in the area of the sulphur deposits. Tkachouk and Koltun (1963) noted the difficulty of using such an interpretation where there has been wid espread development of karst beneath a thick clay cover and fluvial base level. Most authors assumed that karst formed during breaks in sedimentation in the pre-Ratynsky (Pertzovych, 1969) or pre-Kosovsky (Bobrovnik and Golovchenko, 1969) time. There are no reliable data supporting such assumptions. The presence of a sharp, gapping contact between the upper limestones and the underlying GAS, and the karst-like irregularity of the top of the GAS are used as the main arguments to support such breaks, besides the very fact that the Tyrassky Formation is highly karstified (which in "descending" karst would require exposure above fluvial base level). In fact such a contact can be formed without a break but under the artesian conditions by interstratal solution, and karstification does not necessarily require exposure of soluble rocks to the surface. All of the above peculiarities of karst in the Tyrassky Formation can be explained by the artesian speleogenetic model. Artesian cave origin by waters ascending from the underlying Lower Badenian aquifer is not a special case (as adopted by Polkunov et al., 1979) but is characteristic of the entire region of the GAS and supra-gypsum limestones. "Ascending" karst development in the supra-gypsum limestone is clearest where the limestone directly overlies the Lower Badenian member, but it can also occur where there are karst conduits within the GAS. In the latter case, aggressiveness of water toward calcite is enhanced by the presence of H2S and CO2 generated by reduction of sulphates. That the caves in the supra-

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.14 gypsum limestones do not display maze patterns like those in the GAS is explained by differences in the nature and distribution of initial fissure permeability. The model of artesian speleogenesis leads to a unified regional interpretation of karst in the Miocene strata. Moreover, it explains the important role of karst in governing the exchange of water between aquifers and in the formation of sulphur deposits. Biogeochemical processe s in the formation of sulphur deposits It is well accepted at present that all major sulphur deposits are epigenetic, formed by replacement of the parent gypsum/anhydrite with calcite and sulphur ores. This opinion is shared by most of the researchers that have studied PreCarpathian deposits (Aleksenko, 1967; Koltun, 1966; Koltun et al., 1972; Niec, 1992; Osmolski, 1973; Pawlowski et al., 1979; Sakseev, 1972; Sokolov, 1965, 1972; Kityk, 1979; Vinogradov et al., 1961). Replacement of sulphate rock with calcite and sulphur occurs in a process of redox reactions resulting in reduction of sulphate to sulfide and oxidation of organic carbon to CO2 and water. In low-temperature diagenetic environments these reactions are driven by microorganisms, as shown by microbiologi cal and geochemical (particularly isotopic) studies by Davis and Kirkland (1970), Feely and Kulp (1957), Ivanov (1964, 1972), Kirkland and Evans (1980), Ruckmic et al. (1979), Vinogradov et al. (1961), and others. Oxidation of reduced sulphur may occur either abiologically or (less common) microbially. The general processes of bacterial sulphate reduction and of sulphides oxidation are well known, but their stages and mechanisms are still controversial (Kushnir, 1988). The main processes and settings are briefly considered below. The most important process in epigenetic sulphate transformation is sulphate reduction driven by microbes, a heteroge neous assemblage of Desulfo-x The process can be described by the following simplified reaction: SO4 = + 2CH2O H2S + 2HCO3 [1] anaerobic bacteria Sulphate reducing bacteria are strictly anaerobic and require a reduced environment (Eh < -100 mV) for growth, either in the bulk environment or in microenvironments that may be present, or maintained by bacteria themselves, within an otherwise more oxidizing milieu. The bacteria generally cannot metabolize saturated hydrocarbons, such as methane, but require specific organic compounds, such as organic acids or aldehydes (Kushnir, 1988; Machel, 1992). Transformation of methane into more simple oxygen-bearing compounds can be provided by aerobic fermenting bacteria that are widespread under natural conditions. It is assumed that transformation of organic matter can take place in anaerobic environments by bacteria that co-exist with sulphate-reducing bacteria and provide nutrients for them (Kushnir, 1988). The solubility of H2S in water is rather low. At pH = 6.5-9.0, hydrosulphides predominant in solution. Most of the H2S exists as a gas phase. Calcium and bicarbonate commonly precipitates as CaCO3, utilizing the HCO3 generated in the process of sulphate reduction: Ca++ + HCO3 + OH! CaCO3 + H2O [2] For calcite to precipitate continuously during the process of sulphate reduction it is necessary that H2S be eliminated from the reaction zone, as its accumulation up to 500-700 mg/L would make the environment inappropriate for bacteria growth, and sulphate reduction would cease. Moreover, calcite is stable at pH generally higher than 7.0 (depending on the chemical environment). Excess H2S can be eliminated by water flow or by oxidation in situ to elemental sulphur. Kushnir (1988) hypothesized that anomalous sulphate reduction can be performed by bacteria if environmental conditions deteriorate; at this sulphate is reduced to thiosulphate (S2O3 =). In this way sulphate-reducing bacteria can partially consume low-activity organics, including methane. Isotopically light carbon in secondary calcite is one of the main points of evidence of its bioepigenetic origin. It inherits the isotopic composition of the initial organic compound. In the Pre-Carpathians, light carbon in epigenetic limestones of the Tyrassky Formation ( C13 = -32 to -65 ‰) clearly shows that the initial organic compound involved in sulphate reduction was methane. This origin is supported also by the presence of light carbon ( C13 = -32 to -42 ‰) in CO2 in the air of caves that have only recently been aerated, and where sulphate reduction and CO2 generation still occurs in the lower part of the GAS, or within the still-wet clay fill, such as in Zolushka cave (Klimchouk et al., 1984; Klimchouk and Jablokova, 1989). The CO2 concentration in the air of such caves can be as high as 4.0% Thus, although methane cannot be utilized immediately by sulphate-reducing bacteria, it still serves as the initial organic compound. Models of the origin of sulphur deposits should stipulate conditions for methane transformation into more active forms.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.15 Sulphate-reducing bacteria do not form elemental sulphur. Its formation is due to abiological or microbial oxidation of H2S or to polysulfide dissociation. According to Machel (1992), abiological oxidation of H2S is the single most important process forming epigenetic native sulphur deposits in low-temperature diagenetic environments: 2H2S + O2 2So + H2O [3] 2HS+ O2 + 2H+ 2So + 2H2O [4] Hydrogen sulfide (and/or HS-) reacts with O2 not only to elemental sulphur but also to sulfite, disulphate, sulphate, and polysulphides. The reaction varies with pH, concentration of reactants, and the presence of impurities. In general, the formation and precipitation of elemental sulphur is favoured by the following conditions (Kushnir, 1988; Machel, 1992): (a) relatively high concentrations of total reduced sulphur (S=); (b) relatively large S=/O2 ratios; and (c) relatively low pH (<6). With increasing pH, reduced sulphur is increasingly oxidized to sulphite, thiosulphate, and sulphate, rather then to So. This also explains the presence of barite and celestite as late mineral phases in many sulphur deposits, or in oxidized portions of these deposits. Moreover, at pH values between about 6 and 8.5, much if not most of the So that is generated by r eactions [3] and [4] forms polysulfides rather then precipitating as a separate phase. Thus, for the formation of native sulphur by these reactions, the environments must be nearneutral to slightly acidic, with relatively low concentrations of oxygen, and low rates of oxygen supply. At high partial pressure of oxygen, or where oxygen is supplied at large rates, hydrogen sulfide (and/or HS-) is oxidized to SO4 = (e.g. sulphuric acid) and no elemental sulphur is formed (Machel, 1992). Where there is an excess of SO4 =, oxidation of H2S and formation of colloidal sulphur can occur in the absence of oxygen (Feely and Kulp, 1957): 3H2S + SO4 = 4So (colloidal) + 2H2O + 2OH[5] It is assumed that such sulphur precipitates together with calcite, to form a calcite/sulphur ore containing uniformly scattered sulphur. The concept of the formation of sulphur ores by reactions [1] and [5] has been widely accepted until recently, but Kushnir (1988) has pointed out that it is as yet unproved, and that thermodynamic analysis shows it to be unlikely. In the presence of colloidal sulphur, the combined origin of calcite and sulphur occurs also in the following way: 2H2S + 2Ca++ + 2H2O Ca(OH)2 + Ca(HS)2 + 4H+ [6] Ca(HS)2 + 4S(colloidal) CaS5 (polysulfide) + H2S [7] CaS5 + CO2 + H2O 4S (crystalline) + CaCO3 + H2S [8] Polysulfides form as direct and indirect byproducts of bacterial sulphate reduction, or inorganically when H2S dissociates into HSand then reacts with elemental sulphur that is derived from reactions [3] and [4] (Machel, 1992): HS+ (n-1)/8 S8 Sn = + H+ [9] The concentration and speciation of polysulfide is strongly pH-dependent. In slightly to moderate alkaline aqueous solutions (pH 7-9) they are stable as anions. At pH 8 to 9 the polysulfides S4 =, S5 =, and S6 =, along with HS-, are the predominant sulphur species in aqueous solution. At neutral to slightly acidic pH (pH < 7-6), polysulfide anions rapidly dissociate: Sn = So + Sn-1 = [10] Thus, elemental sulphur is "consumed" in the formation of polysulfides at pH > 6-7, and polysulfides release So upon a decrease in pH below this range, which leads to the precipitation of elemental sulphur (Machel, 1992). Kushnir (1988) suggested the thiosulphate model of sulphur metasomatism, by which only anomalous sulphate reduction to thiosulphate occurs: SO4 = + organic matter S2O3 = + HCO3 + H2O [11] Thiosulphate then migrates to adjoining zones where it is reduced to H2S resulting in an increase of pH: S2O3 = + organic matter H2S + HCO3 + OH[12] At the boundary of the zones of different pH, calcite and sulphur precipitate: Ca++ + HCO3 + OH! CaCO3 + H2O [13] S2O3 = + 2H2S 3S + H2O + 2OH[14] It is assumed that this process leads to the formation of sulphur ores with uniformly scattered sulphur. The thiosulphate model is supported theoretically by thermodynamic analysis, but this is based on numerous assumptions and has not been proven by experiment.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.16 Mechanisms of replacement of parent rocks Soviet scientists have distinguished between two mechanisms of replacement of host gypsum/anhydrite with calcite or sulphur/calcite rocks: hydrogenic and metasomatic. Hydrogenic replacement occurs by precipitation of a new mineral in the place of a dissolved one. However, the mechanisms for metasomatic replacement is unclear, despite of the fact that sulphur/carbonate metasomatism is accepted by most researchers as a major ore-forming process. Sokolov (1972), who introduced and developed the concept of metasomatic origin of sulphur deposits, recognised that the conditions for metasomatism of sulphates to sulphur-carbonates, and even the process itself, are still poorly understood. Little progress has been made since then (Kushnir, 1988; Niec, 1992). With regard to the conditions for exogenous sulphur formation, metasomatism is usually referred to as a process of replacement of sulphate minerals with the almost simultaneous formation of calcite and sulphur, as a result of biochemical and abiotic reactions. Replacement occurs in watersaturated zones of microporosity, with the host rock retaining its original volume and solid state (Korzhynsky, 1955; Lazarev, 1972; Sokolov, 1972). It is believed that replacement of ions occurs directly within the crystal lattice without the lattice being disrupted (Polkunov et al., 1979). Not only is the process itself unclear, but also the mechanisms by which reactants are supplied to the metasomatic front and byproducts are removed. The process is postulated to occur in the microporous space of a rock, but studies have yet to evaluate this kind of porosity of a gypsum/anhydrite rock and matrix permeability. The chemistry and flow velocities of pore water in these materials are uncertain, and reaction rates are poorly understood. As the metasomatic front advances deeper into the sulphate rock, the inflow of reactants must, according to most schemes of ore formation, proceed through the newly-formed sulphur/calcite or calcite. As the result, the dynamics and chemistry of the pore solutions must change considerably, so that the conditions postulated for the process to start will change as well. In general, the existing models for the process imply one of the following: either simultaneous or sequential input of a reducer and oxidizer to the metasomatic front, or the simultaneous presence of reducing and oxidizing environments in close proximity. Such conditions are impossible, or at least highly improbable, in zones of microporosity, especially if metasomatic front has propagated through a considerable thickness of rock. Many authors actually invoke concurrent or conjugate dissolution and precipitation. It is quite characteristic that most of authors, who postulate the process of metasomatosis in the origin of sulphur ore, discuss hydrologic and hydrochemical settings of the formation of sulphur ore (supply of reactants and outflow of r eaction products) in terms of dynamics and chemistry of waters in the macrospace. In the light of the above problems, it is questionable that metasomatism is a major process in the formation of calcite and sulphur-calcite epigenetic rocks, at least under hypergenic temperatures and pressures. The importance of macroporosity within the parent gypsum/anhydrite should be stressed as a media of transmitting reactants to and away from the sites of metasomatosis. Also, hydrogenesis nearly always takes place in the formation of epigenetic calcite and sulphur ore and is probably the dominant replacement mechanism. As shown below for the Pre-Carpathians, when an adequate hydrodynamic model for the origin of sulphur is presented, most of the geologic data commonly used to argue in favour of metasomatism are more likely to support hydrogenic replacement. Water exchange in the Miocene sequence and the origin of sulphur deposits in the Pre-Carpathians The pattern of water exchange is critical in any proper genetic model for the sulphur deposits. Such pattern forms the spatial and temporal framework within which the processes take place; it controls geochemical environments, the migration of reactants and reaction products between them. It is also important to consider the evolution of this flow pattern as the result of internal factors (such as the development of karst permeability) and external factors (such as changes in recharge/discharge conditions during neotectonic and geomorphic development). As noted above, it is common in the regional hydrogeology of the Pre-Ca rpathians to consider the GAS to be an aquifuge and separating bed between the two limestone aquifers, preventing water exchange between them, except through windows where the GAS is absent. This concept has been part of all models suggested so far for the origin of sulphur deposits of the Pre-Carpathians. Several problems arise from this interpretation:

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.17 1. Large-scale sulphate reduction requires a stable supply of large amounts of dissolved sulphate into the reaction zone, which can not be provided if the GAS behaves as an aquifuge. The surface area of the sulphate contacting with water would be limited. If this scheme would be the case, then dissolution of sulphates would occur more or less equally in both the sub-gypsum and supragypsum aquifers. However, the concentrations of sulphate are normally small in the sub-gypsum aquifer and high in the supra-gypsum one. 2. If the GAS is impermeable, a uniform supply of hydrocarbons into the zone of sulphate reduction would be impossible. According to conventional hypotheses, hydrocarbons can penetrate the GAS only through tectonic faults and windows, which would produce localized zones of epigenetic calcite and sulphur-calcite. On the contrary, epigenetic calcite is distributed rather uniform even in the regional scale, and sulphur ores are uniformly distributed in the local scale of deposits. 3. The supra-gypsum aquifer is usually considered to be the zone of sulphate reduction. The generally accepted scheme also considers this aquifer to be the zone of H2S oxidation, which occurs concurrently or se quentially with sulphate reduction. Attempts to constrain all stages of the sulphur formation processes into a single aquifer have forced investigators to make rather doubtful assumptions about paleohydr ogeology. In other cases, the formation of sulphur is considered to be associated with windows into the GAS, or with marginal zones, where H2S-bearing waters and oxygen-bearing waters from two different aquifers can mix (Aleksenko, 1967). This idea obviously contradicts the presence of sulphur ores above the GAS at considerable distances from margins or "windows". Moreover, this hypothesis ignores the fact that the local thinning and windows within the GAS are in most cases the result of its epigenetic replacement with sulphur ores or with barren limestones. The GAS was initially present, and if itÂ’s low permeability is implied, then it would prevent mixing of water and the formation of epigenetic calcite and sulphur ores. 4. According to the "window" scheme, the supply of reactants to the zone of sulphate reduction comes to the GAS from above. As the replacement front propagates downward from the top of the "impermeable" GAS, newly formed epigenetic rocks are incorporated into the upper limestone. Obviously hydrodynamic and hydrochemical conditions behind the replacement front must change considerably during the process, so that the supply of reactants and the outflow of reactions products would likewise change. This would cause the processes and lithologic patterns at the reaction front to change. However, no such evidence has been observed in the Pre-Carpathian deposits. 5. If metasomatic replacement of the GAS had actually occurred, then the metasomatic front would have had to propagate deeply into the supposed "impermeable" GAS for 20-30 m. As pointed out previously, it is doubtful that metasomatism could do so without the presence of macroporosity. Thus, the notion that the GAS is impermeable not only contradicts the hydrogeologic data and ignores well developed karst in it, but it also casts doubt on some of the basic assumptions about sulphur genesis. The result has been 40 years of debate that continues even today. The artesian hypothesis of speleogenesis introduced in the previous section explains the major peculiarities of water circulation within the Miocene sequence and supports the karst model for the origin of sulphur deposits for the PreCarpathians. The major points of this model are presented below. 1. Ascending water flow though the GAS, from the sub-gypsum aquifer to the supra-gypsum one, occurs in potentiometric lows where sufficient fissure permeability is present. These lows coincide with valleys entrenched into the capping clay aquiclude. The whole aquifer system discharges upward into the bottoms of such valleys. By this process, karst permeability develops in the GAS in the form of twoto four-storey cave systems (Figs 11 and 12). Maps of known maze caves show typical delineation of karstified areas (Fig. 10). The association of these "ascending" karst systems with old valleys explains why sulphur deposits that form after cave systems are also associated with old valleys (the "paleogeographic criterion" of Otreshko). 2. Waters of the sub-gypsum aquifer are commonly high in oxygen and calcium bicarbonate, and low in TDS. Gases (predominantly methane) enter the aquifer along faults from the adjacent foredeep and disperse various distances from the points or lines of input. Under aerobic conditions within the aquifer, methane is transformed under microbial mediation to simple organic compounds that can be utilized by sulphate-reducing bacteria.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.18 Fig. 12 Karst model of the formation of sulphur deposits in the Pre-Carpathians. A = initial stage, B = mature stage. Quaternary sediments : 1 = sands and loams. Upper Badenian sediments : 2 = clays and marls; 3 = sandstones; 4 = Ratynsky limestones; 5 = epigenetic sulphur-bearing and barren limestones; 6 = gypsum and anhydrite. Lower Badenian sediments : 7 = lithothamnion limestones; 8 = sands and sandstones. Upper Cretaceous sediments : 9 = marls and argillaceous limestones. 10 = dissolu tional cavities; 11 = pattern of water flow in major aquifers; 12 = pattern of water flow through the gypsum/anhydrite. 3. Waters from the sub-gypsum aquifer ascend into karst fissures in the GAS and gain calcium sulphate composition. The TDS increases and the oxygen content and Eh decrease. The upper part of the GAS, near the contact with the overlying limestone, is the zone of ac tive sulphate reduction. When conditions are favourable, replacement of the gypsum/anhydrite rock by calcite or sulphur-calcite takes place there. Excess H2S escapes into the supra-gypsum aquifer and discharges along with water though the clay cap. 4. As sulphate reduction proceeds, the SO4 = content of the waters decreases, so that water at the upper contact again acquires (or increases) its solutional capacity with respect to gypsum. Dissolution of gypsum at the top of the GAS is concurrent with precipitation of calcite in the bottom of overlaying limestone (hydrogenic replacement). The newly formed epigenetic rock incorporates into the upper limestone horizon. In this way, the sharp, gapping contact is maintained between the GAS and epigenetic limestones, with clearly discerned dissolution features in the top of the GAS, which are peculiarities noted by many researchers. If sulphate reduction processes proceeds for a long enough time, or at high enough rate, the replacement front propagates downwards through the full thickness of the GAS, removing it entirely. Along the strike, as intrastratal water flow diminishes, epigenetic limestones may thin gradually with increasing distance from karst zones, from which reactants emanate. According to this model, reactant rise from below, so hydrochemical conditions remain rather stable as the replacement front propagates downward. 5. In the upper part of the GAS, replacement may occur also by metasomatic mechanism, encompassing the rock pillars between the conduits in dissolutional networks. The resulting features include relict structures inherited from the parent gypsum/anhydrite, abrupt transitions from one rock type to another along the strike, mineral complexes similar to those of the original rock.

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.19 6. Combinations of hydrogenic and metasomatic replacement of gypsum-anhydrite, as envisioned in the artesian karst model, explain virtually all major geological features in the supra-gypsum limestones. They also eliminate the seemingly contradictory aspects of structure and texture of epigenetic calcite and sulphur ores that have been the source of much argument by previous investigators. 7. If the vertical connectivity between fissure networks occurring in the lower part of the GAS and its higher horiz ons is poor, local conduit-fissure networks may form in the lower part of the GAS that allow back-circulation of water facilitated by natural convection. Thus water is able to flow through the GAS and return to the sub-gypsum aquifer (Fig. 12). This explains the local appearance of sulphate waters in the sub-gypsum aquifer, as well as local presence of epigenetic calcite and sulphur beneath the GAS. The latter have been one of the most contradictory points within previous models (Sakseev, 1966; Aleksenko, 1967; Kityk, 1979). 8. The karst model proposed here leads to three possible geochemical scenario of the origin of barren and sulphur-bearing limestones. That which of these situations takes place depends on parameters of local environment, which are controlled, to a considerable extent, by hydrodynamic conditions. Hence, they are dependent on karst systems development. (a) In some areas, oxygen-rich groundwater flows laterally into the supra-gypsum aquifer from the main recharge area. Additional oxygen is supplied by water rising from the sub-gypsum aquifer through windows in the GAS defined by facies or tectonic. Where oxygen-rich water in the supragypsum aquifer encounters karstified zones in the GAS, it mixes with ascending H2S-bearing water emerging from the GAS. If oxygen is supplied slowly and continually, then sulphur ore is formed by the "classic" mode of hydrogenic replacement (reactions [1], [3] and [4]). Zones of karst development are those of relatively low head toward which both sources of water are drawn. (b) Where discharge from the artesian system is slow, reducing conditions and H2S-bearing waters extend throughout the entire supra-gypsum aquifer above karstified zones in the GAS. High sulphate concentrations occur at the top of the GAS, where H2S is oxidized according to the scheme proposed by Feely and Kulp (1977) to produce sulphurcalcite ores with uniformly scattered sulphur (reactions [1] and [5]). The framework of the karst model allows also other mechanisms for the formation of sulphur described earlier. (c) Where conduit paths in the GAS are well developed, waters ascending through most transmissive conduits can still contain much oxygen, even in the upper part of the GAS. Mixing of these waters with H2S-bearing waters from peripheral areas of slow flow can lead to sulphur deposition by oxidation of reduced sulphur compounds (reactions [3] and [4]). 9. Spatial and temporal variations of hydrodynamics within the Miocene aquifer system cause the groundwater exchange pattern to be complex and variable. Controlling factors are both external (variations in recharge and discharge to the system) and internal (permeability variations caused by karst development). Changes in pH and Eh occur in different parts of the system, as well as the migrating boundaries between geochemical environments. Thus, certain mechanisms of the formation and secondary transformation of sulphur can start or stop at various times and places within the system. This explains the lithologic diversity of sulphur ores and their varied relationships with barren limestones. Regional pattern of karst evolution and the formation of sulphur deposits During the late Pliocene and early Pleistocene the most active uplift occurred in the Rostochje and Opolje (northeast flank of the northwestern section of the GAS belt; see Fig.2), where widespread erosion almost completely removed the clay cap and the Tyrassky Formation. Elsewhere the initial surface stream pattern was initiated at that time. The formation of the old Pre-Dniester alluvial plain (7th terrace) is dated as late Pliocene. Stream channels migrated widely over this plane (Gofshtain, 1979). Simultaneously, other valleys formed north of the Pre-Dniester plain (Fig. 7), as well as on the Shchiretsky depositional plain in the northwest part of the GAS belt. The 6th Dniester terrace (early Pleistocene) is narrower then 7th but still much wider than the modern Dniester valley. Thus, as erosional valleys began to entrench into the clay cap, the regional pattern of potentiometric highs and lows in the Miocene artesian system was established. This caused intensified water circulation and initiation of karst systems of the "ascending" type. In the interior of the platform fringe (Podol'sky area) inflow of hydrocarbons to the sub-gypsum aquifer and therefore sulp hate reduction was slow and limited to a few areas. Therefore, sulphate reduction in zones of as cending karst development was uncommon. Epigenetic limestones are limited

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Alexander B. Klimchouk / Speleogenesis and Evolution of Karst Aquifers, 2005, 3 (2), p.20 to certain zones where hydrocarbons were carried in along faults, allowing sulphate reduction to proceed. In the early Pleistocene, the uplift rate accelerated considerably in the Podol'sky area, causing the Miocene artesian system to rapidly lose its confinement. The resulting influx of oxygenrich water was unfavourable to the origin of sulphur. With further deepening of drainage cave systems became relict. In the northwest section of the GAS, within the Shchiretsky depositional plain, the uplift rate during Pleistocene was considerab ly lower than in the Podol'sky area. Distinct valley entrenchment into the clay caprock occurred only at the end of earlymiddle Pleistocene. Some entrenchments were caused by meltwaters from the waning Mindel glaciers, which covered this area (Gerenchouk et al., 1972). At this time, groundwater circulation through the Miocene sequence increased, as did karst development and sul phur deposition. In a few places (Rozdolsky deposit; Sokolov et al., 1969) erosional valleys entrenched as deep as the top of the Tyrassky Formation (Fig. 6), but in general they remained entrenched up to 30-40 m within the clay cap. Valleys were not deepened further, but instead were buried by late Pleistocene sediments of varied origin. Thus, in this region, conditions favourable to sulphate reduction and native sulphur formation in karst zones were maintained over quite a long time in rather stable hydrochemical environments. In many places such conditions persist even today. Conclusions 1. The development of sulphate karst is one of the major prerequisites for the formation of epigenetic sulphur deposits. It is the source of dissolved sulphates needed for large-scale sulphatereduction. This, likely, applies to most of epigenetic sulphur deposits associated with sulphate rocks, particularly to the Pre-Carpathian sulphur-bearing basin where the formation of sulphur deposits is related to speleogenesis in Miocene gypsumanhydrite stratum (GAS). 2. Certain processes of the sulphur cycle, particularly bacterial su lphate reduction, enhance karstification in sulphate rocks by removing sulphate-ions from solution and maintaining the dissolutional capacity of waters. 3. Field data described in this paper show that the widespread acceptance of the GAS as an aquifuge is erroneous. The GAS does not prevent water exchange between the artesian Miocene aquifers of the Pre-Carpathians. The concept of the “aquifuge” role of the GAS brought about severe contradictions in previous models for the origin of sulphur deposits in this region. 4. Within karst zones, often quite large, the GAS aquifer has high conductivity that allows both vertical and lateral wate r exchange. Dissolution conduit systems in these areas provide paths for water exchange between the Miocene aquifers in the multi-storey artesian system. 5. In such multi-storey artesian system, vast twoto four-storey maze cave systems in the GAS developed where erosional entrenchment into the cap clay caused groundwater to ascend across the stratum. The recent model of artesian speleogenesis, combined with field data in the vicinity of sulphur deposits suggest that relict maze caves in Podol'sky and Bukovinsky areas share the same origin as the dissolutional cavities in sulphur deposits, in the area of present artesian flow. 6. The proposed model for the genesis of sulphur deposits hinges on the ascent of groundwater from the sub-gypsum aquifer to the supra-gypsum aquifer through karst systems in the GAS. Such flow pattern provides the spatial and temporal framework within which the processes of the sulphur cycle take place, as well as it controls geochemical environments, the migration of reactants and reaction products between them. This model is well-compatible with the hydrogeologic settings of the sulphur deposits and with accepted biogeochemical processes of the sulphur cycle. It also explains why sulphur deposits are concentrated around buried valleys and karst zones and resolves many contradictions in the interpretation of geological features of the Tyrassky Formation. The major episode of sulphur-ore formation was from the early to middle Pleistocene, although in some areas it persists even today. 7. The karst model of the origin of sulphur deposits provides new criteria for predicting and prospecting for sulphur, as well as solutions to some of the engineering-geologic problems caused by mining of the deposits. Acknowledgements I would like to thank Kimberley I. Cunningham of the U.S. Geological Survey for his kind and generous assistance in translation of this paper to English. Josef E.Crawford of the Pennzoil Sulphur Company, Pecos, Texas, and Carol A.Hill of Albuquerque, New Mexico, added helpful suggestions. I am deeply indebted to Arthur N. Palmer of the State University of New York for his valuable editorial improvements to the paper.

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