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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. 1 (2003)

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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal www.speleogenesis.info The genesis of the Tennengebirge karst and caves (Salzburg, Austria) Ph. Audra (1), Y. Quinif (2) and P. Rochette (3) (1)* quipe Gestion et valorisation de l'environnement (GVE), UMR 6012 “ESPACE” du CNRS, Nice Sophia-Antipolis University, 98 boulevard douard He rriot, BP 209, FRANCE 06204 NICE Cdex. E-mail: audra@unice.fr (2) Centre d'tudes et de recherches appliques au karst (CERAK), Facult polyt echnique de Mons, 9 rue de Houdain, 7000 MONS, BELGIUM, quinif@fpms.ac.be (3) CEREGE, UMR 6635 CNRS, Aix-Marseille University, Europle mditerranen de l’Arbois, 13545 Aix-enProvence Cedex 04, Fran ce, rochette@cerege.fr Corresponding author Re-published by permission from: Journal of Cave and Karst Studies 64 (3), pp.153-164 Abstract Research has been carried out in the Tennengebirge Massif (Salz burg, Austria) with specific attention to karst morphology, cave systems, and sediment s. This study reveals the genesis of the karst and the underground systems of the Tennengebirge, since the Oligocene. Large horizontal systems, which date back to the Mi ocene, were studied through the example of the caves Hornhhle an d Eisriesenwelt, which respectively repres ent Ruinenhhlen (“cave ruins”) and Riese nhhlen (“giant caves”) The Cosa-Nostra Bergerhhle System is typical of a mostly vertical large high-relie f, alpine cave. The main charact eristic of this network is m ajor development in the vadose zone. The shafts' morphology is in “stairs beneath a faulted roof.” At greater depth, they connect to a perched epiphreatic zone, which is typi cal of a dammed karst. The main undergr ound sediments are of paleoclimatic and hydrodynamic significance, corresponding to hot, stable, or unstable environments (flowstones, rewo rked weathered rocks) and co ld environments (carbonate varves, glacial pebbles). A preliminary study of the Te nnengebirge sediments reveals significant information about its evolution th roughout Pliocene-Quaternary time. Keywords: Tennengebirge, Austria, cave evolution since Miocene, epiphreatic zone, cave sediments 1. Introduction This article describes the main results of research carried out in the Tennengebirge Massif of Austria, specifically in the Cosa-Nostra Bergerhhle cave system. This research concerns the surface karst morphology and, especially, the cave morphology. The area contains 3 types of cave systems including unroofed caves ( Ruinenhhlen ) located on the plateau surface and huge dry systems ( Riesenhhlen ) like Eisriesenwelt, both are related to Miocene conditions, and alpine sytems, like Cosa-Nostra – Bergerhhle, developed during the Plio-Quaternary and reaching considerable depth. Certain extensive underground sediment types are examined, each having specific paleoclimatic significance. The history of the cave system’s evolution is linked to the local physical setting, which includes the altitude of the massif, its shaping during successive uplifts, and its current position in relation to base level. This research has enabled the tracing of the main pattern of karst development in relation to the regional evolution of the massif since the beginning of the Tertiary, and also the establishment of successive phases and conditions of speleogenesis, especially since the late Tertiary (Audra, 1994). The Tennengebirge, located 30 km south of Salzburg, is part of the limestone high Alps, dominating the Danubian Piedmont. This high triangular plateau (20 x 10 km) rises to 2431 m altitude (Fig. 1). Nearly vertical escarpments and abrupt slopes encircle it. Thus it dominates the surrounding valleys by about 1500 to 2000 m, especially the Salzach Valley to the west, which contains one of the region's main waterways. It is one of the richest and best known Austrian massifs for the number and size of its caves. It contains five systems >1000 m deep and two >30 km long.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.2 Fig. 1. Topographic map of the Tennengebirge (contour in terval 500 m). Dashed lines correspond to water tracing. The local carbonate strata extend from the Triassic to the Jurassic (Tichy, 1985). At its base is a thick series of Triassic dolomites, which are overlain by the 1000-meter-thick Dachstein Limestone (Fig. 2). The Jurassic is represented by alternating reddish limestones ( Rtlichen Knollenkalken ), and shale. Above this is an Oligocene quartz conglomerate cover, the Augensteine which caps the massif unconformably. The Tennengebirge consists of an overthrust sheet, in which the normal flank has resulted in a gently inclined plateau, slightly higher in the south. The frontal saddle accounts for the sudden northerly slope. This outcrop pattern has determined the erosional history of the different strata. The dolomites are exposed in the southern part. The karstified Dachstein limestones make up the bulk of the plateau. It is locally sprinkled with gravel veneers of weathered Augensteine Finally, the Jurassic layers have been preserved from erosion at the northern foot of the saddle. 2. The morphology and genesis of different levels of karstification 2.1. The Ruinenhhlen (“cave ruins”): the Hornhhle Example A karst landscape of Miocene heritage (Fig. 3). Hornhhle is located at about 2200 m altitude in a landscape consisting of large cones and depressions, containing destroyed karst forms such as kettleshaped dolines, large grikes, arches, discontinuous tunnels and unroofed corridors called Ruinenhhlen (Lechner, 1949; Goldberger, 1951, 1955). The mainly horizontal tubular galleries whose rock walls Fig. 2. The Cosa Nostra Bergerhhle system and the Tennengebirge (see also text). To the left (3), relationship between cave passages altitude and old karst leve ls (after Klappacher & Knapczyk 1985) Karst development began during the Oligocene beneath Augensteine (1). During the Miocene, horizontal systems developed with alpine water inputs (2), showing different levels (3) relate d to successive phases of stability: Ru inenhhlen (4) and Riesenhhlen (e.g. Eisriesenwelt 5). Following Pliocene uplift, alpine systems developed (e.g. Cosa Nostra – Bergerhhle 6). Entrance horizontal tubes correspond to a Miocene level (7). Shaft series (6) connect to horizontal tubes from Bergerhhle-Bierloch (8), corresponding to Pliocene base level (9). Present water-table at -700 m (10) poors into Brunnecker Cave, which connects to Salzach base level (11).

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.3 Fig. 3. Diagram showing the evolution of the Windischriedel karst, from the Late Tertiary (top) to the Pleistocene (bottom). Fig. 4. Reworked weathered rocks, mainly composed of quartz gravels (light) and iron oxides (dark). are sculpted by scallops and the grain-size distribution of the clastic sediments show that these were once partly flooded, with slow current, and developed close to the water table. They are presently perched above current base level. Augensteine sediments. Hornhhle contains some relatively young flowstones that overlie older flowstone debris broken by neotectonic activity. Most of the older stalagmites are corroded and partly destroyed, bearing witness to their old Miocene age. Fragments ar e also located at the entrances as well as on the surrounding land surface. This abnormal distribution of the flowstones is proof of slope erosion that intersected the cave. Clastic sediments are represented by homogenous sands, bedded and hardened, consisting of quartz and iron oxides, as well as hardened clays. The latter sometimes contain large gravels consisting of rounded ironoxide nodules and Augensteine (Fig. 4). These argillaceous, ferrous, quartzose sediments come from the reworking of the weathered superficial rocks ( Augensteine ). A major karst level, known as the “Hochknig level” (Seefeldner, 1961), developed close to a former base level. This evolution was triggered by the partial erosional removal of the Augensteine from the Lower Miocene onward (Tollmann, 1968). At that time the karst had rivers flowing through it from the insoluble igneous and metamorphic rocks in the central Alps. These helped to feed the caves, as well as erode away the sedimentary cover, part of which is still trapped in the caves. This karst was then uplifted up to 20002200 m. These tectonic movements, along with Quaternary glaciation, accelerated superficial erosion. In such an unstable environment the sedimentary cover was largely cleared from the surface, with reworked remnants preserved in the cave sediments. Glacial erosion of the karst was limited due to the fact that the area was located close to the crest of the massif. However, some carbonate rock was removed, exposing older caves and transforming the weathered low areas into kettle-shaped dolines, once the infilling of weathered rock was partially removed. The karst is still evolving as a result of periglacial activity (frost and snow action): kettle-shaped dolines act as snow-pits, and bare be drock is sculpted by karren and covere d with debris. 2.2. The Riesenhhlen (“Giant Caves”): the Eisriesenwelt example The Eisriesenwelt is one of the world's most famous caves. Its underground ice formations extend through nearly one kilometer of a system that is 42 km long, which attracts thousands of tourists every year. It opens onto the western face

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.4 of the Tennengebirge about half way up the steep cliff that dominates the Salzach Valley (Fig. 2). It is a dry cave consisting of vast sub-horizontal galleries filled with debris. Less common are tubular passages with scallops. The main level, whose initial shape has b een obscured by rock debris, varies between altitudes of 1650 and 1750 m. Lateral tubular labyrinths are located in the range of 150 metres both above and below the main cave axis. Links between different levels consist of steeply inclined tubes. A variety of interesting sediments can be found in the talusfloored galleries. Allogenic fluvial sediments Clastic sediments carried in by vadose flow are represented by sands and pebbles trapped in potholes (in the Fuchsgang and Gerade Kluft). Fine s ilts left by phreatic flow are abundant within all the tubular conduits, which they sometimes choke entirely. Sometimes these can form rounded, case-hardened bodies, known as Krapfen (“donuts”). They contain weathered residue from the local limestone and old karst cavities (limestone grains, quartz gravels, and iron oxides). The fluvial sediments composition gives evidence for a more remote origin, as was noticed by early researchers (Lehmann, 1922; Pia, 1923). Their content of quartz, mica, tourmaline, sphene, zircon, garnet, magnetite, and mica schist fragments suggests an input from the metamorphic zones of the central Alps. These sediments are usually rounded as the result of transport by turbulent flow. Relatively recent flowstone is rare but older flowstone is common. Examples of the latter are very large, e.g. stalagmites in the Steinerne Wald (“Petrified Forest”), and thick, partly eroded flowstone floors (Pia, 1923). Many of them consist of transparent calcite, which shows that their growth was under wood covering (Maire, 1990). Their surfaces are commonly corroded by flowing water. Scallops in the dissolved profile of a stalagmite even show the direction of the paleocurrent that sculpted it. An Upper Miocene cave related to fluviokarst. The Eisriesenwelt was fed by sinkholes located between 1650 and 1750 m, which is about 1000 meters above the present Salzach Valley base level. As demonstrated by minerals within detrital sediments, the runoff came from the central Alps. Sediment from the central Alps was carried mainly into poljes where it disappeared into ponors. Large underground rivers developed, similar to those in the tunnel-like caves presently found in tropical climates. Considerable discharge flowed through these galleries, which reach 50 m in width, creating networks with ma ny ramifications. Eroded flowstone and scallops confirm the direction of flow from the central Alps (SE to NW). Steep tubes and looping profiles demonstrate that the conduits were partially flooded and located close to the water table (Audra, 1994; Huselmann et al in press). This system evolved during periods when the base level was stable, correlative with low-gradient sections of valley slopes. These comprise the “Gotzen level” and “level I” (Seefeldner, 1961), which developed from the end of the Miocene to the beginning of the Pliocene (Tollmann, 1968). The presence of eroded flowstone shows alternating wet and dry phases. These flowstones, dating from the active period of the Eisreisenwelt, are of Pliocene age, as suggested by Lehmann (1922). The most recent flowstones, which are not eroded, probably date from the lower or middle Pleistocene, during a warm interglacial period (Trimmel, 1992). As uplift continued, the Salzach Valley became entrenched. With the relative lowering of base level, the Eisriesenwelt became perched, drained and intersected by scarp retreat. Neotectonic activity is responsible for the boulder chokes. Thus the Eisriesenwelt is a good example of the Riesenhhlen level of karst development, which is recognized throughout the limestone high Alps (Bauer and Ztl, 1972). The study of perched caves sediments clarifies the evolution of former karst environments. In Hornhhle, the sediments are linked to the erosional removal of the Augensteine cover at the beginning of the Miocene. In the Eisriesenwelt, the transition to fluviokarst can be demonstrated by the less-weathered sediments, which were carried in by rivers from the central Alps. The mineralogical evolution of the sediments with altitude (and thus with the age of the caves), can also be found in the plateaus of the Hagengebirge and Steinernes Meer, located farther west (Langenscheidt, 1986). Their composition is comparable to that of piedmont sandstones and conglomerates, also dated from the Miocene (Fchtbauer, 1967; Lemke, 1984; Tollmann, 1968). Each level of karstification includes a link between superficial and deep forms, each of which contains diagnostic sediments. Hornhhle is a Ruinenhhle containing highly weathered sediments, developed in a karst cone in the upper level of the Hochknig. Lower down, the Eisriesenwelt is a Riesenhhle containing fluvial deposits linked to large poljes. Thus, the surface karst features correspond to the vertical arrangement of caves.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.5 Fig. 5. Cross section of the Cosa Nostra Bergerhhle System. 2.3. A typical alpine cave-system: Cosa-Nostra Bergerhhle Shafts and dry tubes. The Austrian limestone high Alps are known as much for their large vertical cave systems as for their cave systems with large galleries. The Cosa-Nostra Bergerhhle System combines these two aspects, giving rise to a vast Pliocene-Pleistocene network, under the joint influence of the last phase of uplift and glaciation. The Cosa-Nostra Bergerhhle System is located in the northwestern Tennengebirge, extending from 2300 m at the Wieselstein summit all the way down to the Salzach at the Lueg Pass (470 m), the low point of the massif (Figs. 1 and 5). The entrance of the Cosa-Nostra-Loch is at 1965 m, 350 m below the Wieselstein summit. However, other as-yet unconnected shafts extend right up to the high point (e.g. Flohschacht, with a depth of –502 m) and probably constitute the upstream end of the system. Lower down, a dozen linked entrances to the Platteneck ice caves (Platteneck Eishhle), are located between 1400 and 1600 m beneath the Platteneck summit. These two caves are the uppermost entrances to the Cosa-Nostra-Loch Bergerhhle and are linked to the main part of the system by sub-vertical conduits. Beyond the entrance series, the Cosa-Nostra-Loch consists of a vertical profile known as a “staircase beneath a faulted roof” (Fig. 6). At -600 m, the Rivire des Incorruptibles (“River of the Incorruptables”) appears. The gradient decreases because of the presence of dolomite beds, which are only slightly permeable (Fig. 5). Beyond -1000 m the river can no longer be followed. The conduit, even though a dry tube, has a nearly horizontal profile, which in places is broken by sudden level changes. A kilometer beyond, this passage connects with the Fig. 6. Profile showing “Stairs beneath a faulted roof” of Cosa-Nostra Loch, a predom inant aspect of vertical systems.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.6 Bergerhhle-Bierloch System. This consists of a labyrinth of about 25 km of dry tubes, partly horizontal but containing many vertical sections. A number of streamlets have entrenched the floors of the tubes into steeply sloping canyons interspersed with shafts, which lead to the present phreatic zone at 700-750 m. There are two entrances at this level of horizontal galleries, Bergerhhle and Bierloch, which are located in the glacial hollow below Platteneck summit, at an altitude of about 1000 m. The lowest part of the system is Brunneckerhhle. Its geometry is completely different from that of the Bergerhhle-Bierloch. During periods of high water, the phreatic zone, normally at 700 m at the sump in the Schotter Galerie (“Pebble Gallery”), forms a torrent that pours into a canyon consisting of ramps and cascades that eventually joins the Salzach level. Although the upper and lower parts of the system are still not connected by exploration, the whole system has a vertical range of >1500 m. In all, it contains about 30 km of surveyed passages. Tertiary and Quaternary sediments. The CosaNostra-Loch contains hardly any sediment. From 1000 m downward, varved carbonate sediments begin to appear that are typical of the Bergerhhle network, covering the cave floor in a thick layer. Some sections include older sediments preserved beneath the varves. There are two successive sequences of sediments (Fig. 7). The lower sequence consists of a variety of pebbles (e.g. in Bierloch). These are overlain by a first generation of large speleothems. The pebbles come from the Jurassic strata overlying the Dachstein limestones. These strata have practically disappeared from the plateau. Their erosion was probably very long ago because there is no eviden ce of their presence in the sediments of the plateau. Today these strata remain only on the northern flank of the massif several hundred meters below the cave system. The pebbles were rounded by fluvial action before being trapped in the karst, and for this reason they are known as “fake cave pebbles” (Schauberger, 1961). They appear to provide evidence for old waterways draining across the Jurassic beds along the northern part of the massif at about 1000 m altitude, and thus about 500 m above the current level of the Lammer River. This suggests that the water inputs for the initial development of the system could have been allogenic. These high levels of fluvial erosion are apparently Pliocene (Toussaint, 1971). The highly varied grain size of this sediment (clays with cobbles), indicates violent discharges with h eavy loads. This suggests a climate with abundant, irregular, and occasionally heavy precipitation. Similar sediments, in the form of pebble conglomerates, which also contain Jurassic components, can be found on the neighbouring Dachstein massif in the Hierlatzhhle (Schauberger, 1983). In the same way, these are linked to former periods of intense surface erosion, combined with torrential subterranean through-flows. The first generation of flowstone after the pebble deposits represents a cessation of stream flow, perhaps as the result of lowering of base level. Their micro-morphology (micrites, alternating with transparent calcite containing fine reddish layers that include considerable clastic material) indicates that they were deposited in an unstable environment, with sparse soil in the process of being removed. The upper sequence consists of varved carbonate sediments locally covered with more recent flowstone. These varved carbonates were deposited in a glacial environment. Thus they correspond to one or more glaciations of the middle or upper Pleistocene. This glacial phase is noted above all for its filling of karst voids at all levels. Even the intercrystalline pores of the ancient weathered flowstones of Bierloch are filled by calcite. Erosion is only superficial, as can clearly be seen on the flowstones in Bergerhhle. In the Hierlatzhhle (Dachstein massif), the glacial varves reach 5 m in thickness in the Lehmtunnel (“Clay Tunnel”), at an altitude close to those of Bergerhhle (Schauberger, 1983). This study noticed a black film covering the varves in Hierlatzhhle, in the Schwartzhalle and Schwartzgang (“Black Hall” and “Black Passage”), and in the Dachstein-Mammuthhle, which gives a dull and sinister look. This can also Fig. 7. Generalized stratigraphic section of the sediments in the Bergerhhle and Bierloch system. Black circles = normal paleomagnetic orientation; white circles = reversed paleomagnetism.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.7 underground soot deposited after a huge fire that ravaged the Dachstein massif during the Atlantic interval (about 4.5-7.5 ka), when the forest was widespread on all the high plateaus. However, the presence of this film in other areas undermines this unconvincing hypothesis. The film is probably simply due to superficial oxidation in contact with air. The end of the upper sequence has sparse speleothems, with some still active in the lower wooded areas of the karst above the Bierloch Brunnecker System. These speleothems give evidence for a return to a biostatic environment (development of thick soils during a period of tectonic stability, when biotic activity was at a maximum). Several samples have been dated. Paleomagnetism has been applied to stalagmites, flowstones and varves. The natural remnant magnetizations (NRM) were measured with a rotating remanometer (JR5A Spinner Magnetometer) during demagnetization in an increasing alternating field (AF), up to 100 mT (Table 1). The measured intensity in flowstones was weak, which raised some difficulties in interpreting the results, as intensity became too low. However, varves show a strong magnetic intensity. Samples BR 4, BR 5, BH 4 (Fig. 8) have strong directional stability during demagnetization with characteristic directions showing normal polarities pointing toward deposition in the Brunhes period (<780 ka). Samples BH 2, BH 3 (Fig. 8) have a poorly defined behavior due to low intensity, nevertheless their direction seems to show reverse polarities that could be related to Matuyama period (>780 ka), but these values are not completely reliable. Varves show a horizontal inclination, linked to sedimentation mechanisms. Magnetic anisotropic susceptibility measurements from BH 4 and BR 5 samples do not show any preferential axis, sedimentation occurring by decantation, without any current. This also proves primary magnetizations, acquired during sediment deposition (Audra and Rochette, 1993). Fig. 8. Orthogonal Zijderveld plots showing the NRM vector evolution during AF demagnetization up to 100 mT (BH 2, BH 4 samples). TABLE 1. Paleomagnetic results. Specimen NRM intensity, characteristic directions obtained from AF demagnetization, magnetostratigraphic interpretation. Sample Sediment type NRM Int. (mA/m) Decl. () Incl. () Polarity Age (ka) 3 BR5 a 0,001 342 59 Normal < 780 ka 3 BR5 b 0,002 340 55 Normal < 780 ka 3 BR5 c Stalagmite Brunnecker 0,002 342 60 Normal < 780 ka 1 BR5 0,02 12 0 Normal < 780 ka 2 BR5 Varves Brunnecker 9,5 334 4 Normal < 780 ka BR4 Stalagmite Bru. 0,4 340 68 Normal < 780 ka BH4 a1 11 21 12 Normal < 780 ka BH4 b1 11 28 12 Normal < 780 ka BH4 a2 11 22 12 Normal < 780 ka BH4 b2 Varves Bergerhhle 13 29 18 Normal < 780 ka BH2 a 0,1 94 -75 Reverse > 780 ka BH2 b 0,2 176 -59 Reverse > 780 ka BH2 c Flowstone floor Bergerhhle 0,3 176 -59 Reverse > 780 ka BH3 a 0,02 131 50 Reverse? > 780 ka? BH3 b 0,02 194 -73 Reverse? > 780 ka? BH3 c Stalagmite Bergerhhle 0,008 148 -47 Reverse? > 780 ka?

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.8 TABLE 2 Geochemical data and sample radiometric age. Sample [U] ppm 234U / 238U 230Th / 234U 230Th / 232Th 234U / 238U t = 0 Age (ka) BL 2 0,038 ( 0,002) 1,254 ( 0,072) 0,795 ( 0,043) 4,8 ( 0,4) 1,395 157,3 (+ 23,5 / -18,5) BR 2 0,228 ( 0,012) 1,357 ( 0,073) 0,047 ( 0,007) 2,2 ( 0,6) 1,362 5,2 (+ 0,9 / 0,8) Two speleothems were dated using U-series (alpha-counting). The weak 230Th / 232Th ratio might show a contaminated system opening, so calculated ages are unreliable (Table 2). However, BR 2 Holocene age is quite certain and concords with sediments stratigraphy. For BL 2, calculated age corresponds to isotopic stage 6, the early beginning of last glacial stage “Riss”. A low probability of speleothem development in a glacial context, indications of system opening and incoherence with sediments stratigraphy that suggest an older age, hence this date can not be taken at face value. These data show the difficulties in establishing long-term chronology in karst environment, linked firstly to the lack of widespread dating methods for the ancient periods and secondly by stratigraphic discontinuities that often hinder correlations. Nevertheless, these initial datings of an Austrian cave system give evidence for two distinct periods (Fig. 2). !" Bergerhhle-Bierloch lower sequence old age is confirmed, its upper part corresponding to flowstones showing signs of reverse polarities, older than 780 ka. Pale omagnetic data is compatible with an early Pleistocene or Pliocene age. The conduits located at a passage level controlled by local runoff, probably allogenic. Lowering of base level interrupted this process, allowing extensive calcite deposition to take place. All of this occurred in an unstable environment with extremely violent high waters, corresponding alternatively with either surface erosion bringing clays or the regrowth of woodland cover giving rise to calcite deposition. This stage can be linked both to the progressive cooling of the climate, as well as to continued uplift of the massif. Additional research is necessary to confirm this hypothesis. !" Interpretation of the second depositional period is based on more solid evidence. Brunnecker Pleistocene upper sequence seems to have developed during a recent normal period, being younger than 780 ka. It is still difficult to attribute its origin to one or more identified glaciations, and to precisely identify the Salzach valley entrenchment steps. Advancing glaciers reactivated the deep karst, causing repeated flooding to heights of >600 m, partially choking the conduits with sediment. After this phase, the system drained and speleothems developed. However, the correlation of late Tertiary karst phases is not clear on a regional scale. This cave system has revealed its potential; the fairly similar layout of Pliocene-Quaternary cave systems of the high limestone Alps makes it possible to compare them and gives solid evidence for this evolutionary model. An accurate chronology, based on a detailed stratigraphy combined with dating, would be of interest. 3. Parameters determining the current morphology and hydrology 3.1. Conduit Morphology in vadose systems Parameters that determine the pattern of vadose conduits. A study of all kinds of vadose conduits leads to a model that incorporates 3 parameters: slope, jointing, and discharge (Fig. 9). The origin of a meander involves a number of factors. Jointing must be moderate (Fig. 9/1). If the slope is gentle, sinuous meanders will evolve into an angular system of joint-controlled pseudomeanders (Fig. 9/2). If the slope is steep, shafts will dominate by capturing the runoff (Fig. 9/3). In addition, the discharge must remain moderate. Otherwise the conduits widen and become rather straight canyons. Depending on the amplitude of the initial floods, the original tubes will either be hardly affected or will alternatively form a “keyhole.” In the Tennengebirge, meanders are rare because of intense jointing and steep conduit gradients. Canyons (in the context of this paper) are underground galleries higher than they are wide, with abrupt rock walls entirely dissolved in limestone. Their main characteristic is their size, which is several meters wide and several dozens of meters to >100 meters high. This type of canyon is generally not sinuous. The condition necessary for the establishment of such a canyon is great discharge (several hundreds of L/sec to several m3/sec). These flows can move large clasts, such

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.9 Fig. 9. Predominant conduits in the vadose zone, in relation to slope, jointing, and discharge. as cobbles several decimetres in diameter, as seen at Brunnecker. If the overall slope is small (Fig. 9/4), then the canyon will consist of a series of basins with channels linking them (e.g. Skocjan Cave in Slovenia). With increasing slope, small potholes appear. At the foot of each vertical drop, mechanical erosion will form potholes with the aid of suspended pebbles (Fig. 9/5). This is the classic morphology of tropical mountains, where highrelief shafts are fed by inputs from large perched basins (e.g Mexico and China; Zhang et al, 1991). Brunnecker includes a canyon that consists of a series of inclined ramps interspersed with cascades of pits and potholes, where the high-water discharge can reach several cubic metres per second. For there to be shafts there must be sufficient topographic potential. In most places, vertical routes are developed along joints, which enable rapid penetration of the limestone layers (Fig. 9/6). Thus, the large vertical systems, with successions of pits, are more likely in highly jointed carbonate rocks. The Austrian limestone high Alps have all these characteristics, and the large vertical shafts are a common feature particular to this area (some being >400 m deep). The intersection of a vertical fault and an inclined fault of 70 or 80 is typically responsible for the “staircase beneath a faulted roof” (Fig. 6). The inclined fault provides a steeply sloping roof, which guides the conduit. Beneath this fault roof a succession of pits develops, each pit separated by narrow windows, forming a giant stairway that can extend over a vertical range of several hundred metres. Jointing and vertical cross-sectional shape of the systems. As with detailed conduit morphology, jointing is a determining factor in the organization of vadose networks. Two ki nds of vadose networks can be distinguished, depending on how well they are adapted to the geologic structure. Where jointing is moderate, condu its have a gentle slope, with meandering canyons predominating over shafts (Fig. 10, top). These conduits have many high-order tributaries and are fed by large drainage basins. The whole network has a slightly concaveupward profile. This makeup is common in gently sloping plateau karst (such as in the Vercors of France). On the other hand, where jointing and topographic potential are very strong, as in massifs with overthrusts, it is typical to have shaft series with a steep descent to the horizontal conduits (Fig. 10, bottom). Each one drains a relatively small basin, in places less than one hectare. The result is a great number of catchments, each with small discharge, leading to the low-gradient conduit system with only a very slight hierarchical arrangement. The whole network has a very concave-upward profile. These are common in the Austrian high alpine karsts. Fig. 10. Morphology of passage in relation to jointing intensity. Above, with little join ting, profile is slightly concave and strongly hierarchical (e.g. Vercors). Below, in highly jointed rock, profile is strongly concave and only slightly hierarchical (e.g. Tennengebirge). 3.2. A dammed karst Karst drainage is determined by several parameters, such as structure of the karst aquifer, position of base level, and duration and phases of karst development. The great homogeneity of the Tennengebirge strata provides no significant differences on scale of the entire massif. The base-

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.10 level position, having continually dropped since the late Tertiary, is a determining factor in the evolution of the karst and the nature of the runoff. The vertical speleogenetic potential, presently more than 1500 m, was developed before the Pleistocene, thus all of the major cave systems were developed before the Pleistocene. All of these variables determine the organization of underground circulation in the northwestern Tennengebirge, including that of the Cosa-Nostra Bergerhhle system. The vadose zone The vadose zone is typified by large shaft systems. As in all high mountain karsts, discharge is highly variable, and peak flows are very powerful. Intense jointing allows the drainage area to be partitioned into many small basins of several hectares each. Thus it is not uncommon for two neighbouring shafts to penetrate deep into the limestone without interconnecting. In the CosaNostra-Loch, the River of the Incorruptibles enters at -600 m and then leaves the passage again at 1073 m. Most of the runoff comes from the Wieselstein area. Its discharge can vary from a few liters to hundreds of liters per second. This water feeds the phreatic zone beneath the BergerhhleBierloch system at about 700 to 750 meters altitude. The phreatic zone. The phreatic zone, whose top is at an elevation of about 700 m, drains most of the northwestern part of the Tennengebirge. The known tributaries come from the Wieselstein (Cosa-Nostra-Loch), Platteneck and from the area above Bierloch. The latter passages are captured by the large N-S Bierloch fault, which during low water acts as a drain toward the north and as a lowpermeability dam to water from the east. This phreatic zone originates from the Jurassic beds of marly schist, which have been preserved from erosion at the northern foot of the massif (Fig. 11). These act as a dam for water in the Dachstein Limestone aquifer. Outlets are located at low points in the geologic structure, which correspond to areas where the Jurassic layers have been eroded the most. The geological map shows these hollows to be located at the foot of glacial valleys (e.g. Winnerfall Spring), or at the outlet of the Salzach Gorge (Klappacher and Tichy, 1986). Glaciers seem to have produced the localized erosion that determined these emergence points, and as a result they have controlled the organization of the entire phreatic zone. According to Toussaint (1971), this level of 700 m, which is clearly marked by horizontal galleries and levels of springs, corresponds to a former static base level. This would correspond to the “level V”, dating from the Lower Pleistocene. Its development is linked to glacial activity, which supports the idea of evolution during this period. The only place where this structural dam has completely disappeared is at the level of the Salzach Gorge, where the erosional power of the glacier was much greater, and where the plunging saddle of Dachstein limestone is dissected by the valley. As a result, this potential outlet of the aquifer below 500 m altitude has led to the development of the Brunnecker Cave System, which forms a link between the 700-meter-high phreatic zone and the present Salzach base level. During low water, low discharge occurs (Fig. 12, top). The entire phreatic zone seems to be drained by the Kuchl Creek Spring (Kuchlbachquelle), which emerges at 670 m altitude in the Infang Meadows (Infangalm), through the slightly karstified Jurassic beds. Their weak transmissivity is enough, however, to transmit moderate discharges. Thus, the “lithological dam” is only a relative concept. Fig. 11. Impermeable Jurassic strata dam the Dachstein Limestone aquifer (not to scale). Spring outlets are located at structural lows where glaciers have eroded the Jurassic beds.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.11 The epiphreatic zone. The level and amplitude of the epiphreatic zone fluctuates seasonally according to the amount of runoff. During high water resulting from storms or snow melt, the input can be 100 or even 1000 times greater than its usual flow, to the point that runoff from many sources reaches the water table more or less simultaneously (Fig. 12, bottom). The Kuchlbach Spring is unable to handle such a large input, even when its overflow outlet, the Infang Wasserloch, becomes active. As a result, the top of the phreatic zone in the aquifer rises several tens of meters, for example about 50 m for Bierloch and the Winnerfall. This rise activates higher-level springs with great discharges. One of these outlets, located about 80 m deep in the Bierloch phreatic zone, bypasses the damming influence of the Bierloch fault, discharging several cubic meters per second into the Brunnecker sump. This runoff pours into the Brunnecker canyon, filling the bottom of the cave system, resulting in flooding to depths up to 50 m. In July 1991, after 10 days of heavy rainfall, the Brunnecker spring discharge was about 5 m3/sec. Similarly, farther east, following a 50meter-high flood, the Winnerfall Spring became active. However large these floodings in the epiphreatic zone may seem, at no time do they reach the 600 m mark, which is the level reached during Pleistocene glacial phases. The emergences. The phreatic zone drains through a series of outlets, of which the main ones are (from east to west) the Dachserfall, Tricklfall and Winnerfall. These are the three largest springs in the massif (Fig. 1). Directly connected to the phreatic zone, these are located at about 700 m altitude and are perche d about 150 m above the low parts of the Lammer valley. The average discharge is several hundreds of L/sec. The Brunneckerhhle, located at the Salzach level (500 m), has grown by capturing the phreatic zone at 700 m, which still remains perched above the cave (Fig. 11). This phreatic zone extends along the northern edge of the massif. It has an unusual drainage pattern. As the inflowing vadose water arrives, these infeeders do not have a preferential direction but instead diverge toward the springs, some of which are far apart from each other. Dye introduced in the Western part emerged from both Winnerfall and Brunnecker (Toussaint, 1971; Fig. 1). Thus the behavior of the phreatic zone depends on the hydraulic conditions. At low water, the drainage is oriented along a south-to-north axis, and the phreatic water is able to drain through the structural dam. In contrast, during high water the structural dam serves as a significant choke, and so the bulk of the discharge is oriented east-to-west, where it escapes by overtopping the structural dam. The runoff then uses better-organised drains at a higher level, which are more capable of evacuating the water. Fig. 12. Nature and direction of the high water (top) and low water (bottom) flow in the 700 m phreatic zone.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.12 4. The karst sediments: paleoclimatic and hydrodynamic markers The deep karst is an environment well suited to preservation. It harbors old sediments that are no longer at the surface because of erosion. Their study is of prime importance to the understanding of karst genesis and paleoclimates (Audra, 1995). 4.1. Typical sediments of hot or temperate environments Reworked weathered rocks. During the Tertiary, the warm, wet climates were responsible for intense chemical erosion of the rocks that produced the clastic covers of the Augensteine, the thick residuum of weathered rocks that were reworked and transported into the karst (Fig. 4). These reworked weathered rocks are mainly made up of Augensteine, clays, and iron oxides (Weingartner, 1983). These make up the oldest deposits in the studied caves, some being certainly linked to the first phases of the cave development. The weathered rocks were trapped in the cavities after the removal of the surficial cover. This erosion was linked to the change of precipitation patterns in the late Tertiary, as well as the climatic degradation at the end of the Pliocene. In any case, this climatically induced erosion was enhanced by tectonic uplift. The clearing of these covers, followed by their trapping in the inside of the karst, seems to be a common characteristic of the Pliocene-Quaternary evolution of temperate mountain karsts (Fernandez Gibert et al 1994; Maire, 1990). Speleothems of warm or temperate periods. The growth of flowstone is a function of vegetation cover, and thus of climate (Quinif, 1992). As a result, speleothems do not normally develop in high alpine karst areas. Their complete absence in the Cosa-Nostra-Loch is c onsidered evidence for a cold mountainous environment. In contrast, flowstone developed in the nearby perched horizontal galleries of late Tertiary age, such as Hornhhle, indicating warm paleoclimates. Today speleothems are able to form at lower altitudes as soon as there is a vegetation cover (e.g. Bierloch and Brunnecker caves; Table 2). The crystalline fabric of the flowstone gives indications of the environment of deposition (Maire, 1990). Transparent calcite indicates growth within a biostatic environment, w ith thick vegetation cover that blocks the descent of clastic sediment into deep openings and promotes supersaturation of infiltrating water, causing rapid speleothem growth. This was especially true for the firstgeneration speleothems of Bergerhhle and Eisriesenwelt, of both Miocene and Pliocene age. These characteristics can also indicate age, because such biostatic conditions have not existed in these mountains for a considerable length of time. Conversely, red and brown impurities and a succession of micro-layers indicate deposition in an unstable environment, when detrital material was carried in from overlying soil, with numerous interruptions of calcite deposition. This sudden acceleration of geomorphic processes, shown by the flowstones in old Tennengebirge cavities, appears to correlate with droughts, as was often the case during the late Tertiary, or periods of widespread cooling, as occurred several times during the Plio-Pleistocene. The flowstone surfaces also recorded events following their development. An eroded surface means that the cave has been reactivated by increased discharge following calcite deposition. Most flowstones in Tertiary caves contain such features, indicating complex climatic cycles during the late Tertiary or due to Pleistocene glaciation. 4.2. Glacio-karstic sediments Carbonate-rich varves in the epiphreatic zone. Glacial abrasion of the limestone massifs pulls off rock particles, which are then brought underground by sub-glacial streams. These particles are composed of calcite flakes (amounting to a total of 35-62% CaCO3) as well as angular quartz grains. Seasonal glacial melting releases large amounts of runoff that overwhelms the underground systems, causing them to flood to heights of several hundred meters. The calcite particles are transported in a “uniform suspension” (Riviere, 1977) that extends throughout the flooded conduits. Later the runoff decants during the slow winter draining of the system. The resulting sediments form alternating light and dark laminae that correspond to the successive hydrologic phases. Thus, these seasonal lacustrine deposits of glacial origin can be considered varves (Maire, 1990), which give unquestionable evidence for past glaciation In the absence of current and with calcite supersaturation, the erosive capacity of the meltwater in the epiphreatic and phreatic zones is very weak. Their speleogenetic effect is mainly to seal the cave systems. There is much evidence for their inability to enlarge caves, such as the preservation of older flowstones, whose surfaces are slightly smoothed. Therefore, this suggests that these deep networks are in many case pre-glacial and that they have

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.13 been affected by a variety of runoff and environmental conditions (Bini et al., 1998). Coarse glacial deposits in the vadose zone. Vadose conduits close to glacial meltwater streams contain easily identifiable fluvio-glacial sediments derived from the erosion of moraines. Their varied petrography consists of numerous crystalline and metamorphic phases carried in from high alpine areas by valley glaciers. These subterranean streams are very competent and can transport cobbles several dozens of centimeters in length, for example at Brunnecker spring (Fig. 13). Fig. 13. Wassergang (“Water gallery”) in Brunneckerhhle, showing deposits of allogenic moraine-derived pebbles. (Photograph by S. Caillault.) 5. Conclusion: evolution and genesis of the Tennengebirge karst The Tennengebirge is distinguished by its great limestone thickness and steep local dips. Due to this thickness, along with alternating periods of uplift and stability, the massif retains a very clear record of all the different karst stages, exemplified by distinct passage levels (Fig. 2). This evidence will benefit future research on this subject, especially as more dates are obtained on deposits. Despite their altitude, the glacio-karstic features are very widespread, and their long evolution dating from the end of the Paleogene has left a strong geomorphic evidence for their Tertiary heritage. The karst is strongly linked with allogenic inputs, including: !" a widespread fluviokarst during the Miocene, !" a fluviokarst limited to the northern slope during the Pliocene, !" a Quaternary glacio-karst, fed by allogenic glaciers, which also blocked the outlets of deep cave systems. The contribution of these external factors, especially the major allogenic water inputs, is the main explanation for the large size of the cave systems. Similar characteristics are well developed in the northern pre-Alpine massifs of France (Audra, 1994). During this long developmental history, the hydrologic function of the karst depended on variations in the nature of the water inputs. High-water periods can be linked to “tropical” precipitation patterns or glacial melting. Concentrated infiltration during these brief periods caused sediment chokes and reactivation of higherlevel passages and perched drains (Audra, 1997). This sort of coincident activation of drains is still quite apparent today. Thus, not only does the Tennengebirge have all the characteristics of high alpine karst, but it also contains evidence for a long karstic evolution, which many other more relatively recent evolved alpine massifs do not have. For example, the Tennengebirge contrasts with certain massifs of Savoy (France) and Switzerland, where the exposure of carbonates by erosion of their insoluble cover has occurred much more recently (Maire, 1990). Acknowledgement I would like to thank P. Wilson for the translation of this article, and A. N. Palmer and the two anonymous reviewers for their helpful comments.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.14 References Audra, Ph. 1994. Karsts alpins. Gense de grands rseaux souterrains. Exemples: le Tennengebirge (Autriche), lÂ’Ile de Crmie u, la Chartreuse et le Vercors (France) [Alpine karsts. Genesis of large cave systems]. Thesis. Grenoble 1 University, France. Karstologia Mmoires (5), 280 p. Audra, Ph. 1995. Signification des remplissages des karsts de montagne [Signification of mountain karst infillings]. Karstologia (25), 13-20. Audra, Ph. 1997. Le rle de la zone pinoye dans la splogense [The epiphreatic zone part in the speleogenesis]. Proceedings of the 12th International Congress of Speleology, La-Chaux-de-Fond. International Union of Speleology / Swiss Speleological Union, la Chaux-de-Fonds. Vol. 1, 165-167. Audra, Ph. and Rochette P. 1993. Premires traces de glaciations du Plistocne infrieur dans le massif des Alpes. Datation par palomagntisme de remplissages la grotte Vallier (Vercors, Isre, France) [First traces of lo wer Pleistocene glaciations in the alpine massif. Datation by paleomagnetism of the fillings in the Vallier cave (Vercors, Isre, France)]. Compte-rendu lÂ’Acadmie des sciences, S. 2, 317, (11), 1403-1409. Acadmie des sciences, Paris. Bauer, Fr. and Ztl J. 1972. Karst of Austria. In: Herak, M. and Stringfield (Ed.), Karst, the important karst regions of the Northern Hemisphere. Amsterdam, Elsevier, 225-265. Bini, A., Tognini, P. and Zuccoli, L. 1998. Rapport entre karst et glaciers durant les glaciations dans les valles pralpines du Sud des Alpes [Karst and glaciations in the Southern pre-alpine valleys]. Karstologia (32) 7-26. Fernandez Gibert, E., Palo mars, M., Rossi, C. and Tortosa, A. 1994. Analisis de procedendia en arenas karsticas: evidencia de una cobertera permotriasica erosionada en el macizo occidental de Picos de Europa [The origin of karst sands: evidence of Permo-Triassic cover erosion in Western Picos de Europa massif]. Actas del 1 congreso internacional sobre Picos de Europa, Oviedo 1991. Asturian Speleological Federation, Oviedo, 1-14. Fchtbauer, H. 1967. Die Sandsteine in der Molasse nrdlich der Alpen [Molasse sandstones in the Northern Alps]. Geologische Rundschau (56) 266300. Goldberger, J. 1951. Reste abgetragener Hhlen auf dem Hochknig [Relics of eroded caves in Hochknig]. Die Hhle (1) 9-11. Goldberger, J. 1955. Die Altlandschaft auf dem Hochknig [Inherited landscapes in Hochknig]. Mitteilungen der sterreic hischen Geographischen Gesellschaft (97) 183-191. Huselmann, Ph., Jeannin, P.-Y. and Monbaron, M. Genesis of caves a new, comprehensive model. Zeitschrift fr Geomor phologie (submitted paper). Haseke-Knapczyk, H. 1989. Der Untersberg bei Salzburg [The Untersberg near Salzburg]. Doctoral Thesis University of Innsbruck, Wagner. 224 p. Klappacher, W. and Knapczyk, H. 1985. Salzburger Hhlenbuch [SalzburgÂ’s caves book]. Vol. 4 (Tennengebirge). Landesverein fr Hhlenkunde, Salzbourg, 557 p. Klappacher, W. and Tichy, G. 1986. Geologische Karte des Tennengebirges [Tennengebirge geological map]. Salzburger Hhlenbuch. Vol. 4 (Planbeilagen). Landesverein fr Hhlenkunde, Salzbourg. Langenscheidt, E. 1986. Hhlen und ihre Sedimente in den Berchtesgadener Alpen [Caves and their sediments in the Berchtesgaden Alps]. Forschungsbericht (10). Berchtesgaden National Parc, 95 p. Lechner, J. 1949. Neue ka rst und quellengeologische Forschungen im Toten Gebirge [New karstic and hydrogeologic researches in Totes Gebirge]. 3. Vollversammlung der Bundeshhlenkommission. Vienna, 32-38. Lehmann, O. 1922. Die groe Eishhle im Tennengebirge (Salzburg). (Eisriesenwelt) [The large ice-cave in Tennengebirge (Salzburg). (Eisriesenwelt)]. Spelologisches Jahrbuch (2), 52121. Lemke, K. 1984. Geologische Vorgnge in den Alpen ab Obereozn im Spiegel vor allem der deutschen Molasse [Geological conditions in the Alps from Upper Miocene compared to the german molasse]. Geologische Rundschau (1), 371-398. Maire, R. 1990. La haute montagne calcaire [The calcareous high mountain]. Karstologia Mmoires (3). Thesis Nice University. 731 p. Pia, J. 1923. Die groe Eishhle im Tennengebirge (Salzburg), (Eisriesenwelt). Geologische Beobachtungen [The large ice-cave in Tennengebirge (Salzburg). (Eisriesenwelt). Geological observations]. Spelologisches Jahrbuch (2), 48-65. Quinif, Y. 1992. L'apport des mthodes de datation absolue: la mthode Uranium / Thorium [Contribution of absolute dating methods: the U/Th series method]. In: Splo-club de Paris (Ed.): Journes Pierre Chevalier, Grenoble, 248-260. Riviere, A. 1977. Mthode granulomtrique. Techniques et interprtations [Granulometric method. Technic and interpretations]. Masson, Paris. 167 p. Schauberger, 0. 1961. ber falsche Hhlenschotter [The fake cave pebbles]. Die Hhle (12) 12-14.

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Ph.Audra, Y. Quinif and P. Rochette / Speleogenesis and Evolution of Karst Aquifers, 1(1) January 2003, p.15 Schauberger, O. 1983. Geologische und morphologische Beobachtungen in der Hierlatzhhle (Dachstein) [Geologic and morphologic observations in Hierlatzhhle cave (Dachstein)]. Schiftenreihe des Heimatmuseums, “Ausserland” (4) 1-21. Seefeldner, E. 1961. Salzburg und seine Landschaften: eine geographische Lande skunde [Salzburg and its landscapes: a geographic contribution]. BerglandVerlag, Salzbourg 573 p. Tichy, G. 1985. Geologische bersicht [Geological overview]. In: Klappacher, W. and Knapczyk, H. (Ed.): Salzburger Hhlenbuch. Vol. 4 (Tennengebirge). Landesverein fr Hhlenkunde, Salzbourg, 27-45. Tollmann, A. 1968. Die palogeographische, palomorphologische und morphologische Entwicklung der Ostalpen [Paleogeogaphy, palaeomorphology and morphologic evolution of the Eastern Alps]. Mitteilunge n der sterreichischen Geographischen Gesellschaft (I-II), 224-244. Toussaint, B. 1971. Hydrogeologie und Karstgenese des Tennengebirges (Salzburger Kalkalpen) [Hydrogeology and karst genesis of the Tennengebirge (Salzburg calcareous Alps]. Steirische Beitrage zur Hydrogeologie (23), 5-115. Trimmel, H. 1992. Quelques remarques sur le dveloppement des grottes des Alpes orientales au Plistocne et l’Holocne [Some remarks about Eastern Alps cave evolu tion from Pleistocene to Holocene]. In: Samon J.-N. and Maire R. (Ed.), Karst et volutions climatiques, hommage J. Nicod. University Press, Talence, 285-292. Weingartner, H. 1983. Geomorphologische Studien im Tennengebirge [Geomorphological studies in Tennengebirge]. Arbeiten aus dem Institt fr Geographie der Universitt Salzburg (9), 205 p. Zhang, Sh., Audra, Ph. and Maire, R. 1991. Les systmes karstiques et les cavits [Karst systems and caves (in China)]. Karstologia Mmoires, Gebihe 89, karsts de Chine (2) 150-161.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal www.speleogenesis.info Basic processes and mechanisms governing the evolution of karst Wolfgang Dreybrodt1 and Franci Gabrovšek2 1 University of Bremen, Germany. E-mail: dreybrod@physik.uni-bremen.de 2 Karst Research Institute ZRC SAZU, Postoj na, Slovenia. E-mail: gabrovsek@zrc-sazu.si Re-published by permission from: Gabrovšek, F. (Ed.), Evolution of karst: from prekarst to cessation. Postojna-Ljubljana, Zalozba ZRC, 115-154. Abstract Models of karstification based on the physics of fluid flow in fractures of soluble rock, and the physical chemistry of dissolu tion of limestone by CO2 containing water have been presented during the last two decad es. This paper gives a review of the basic principles of such models, their most important results, and future perspectives. The basic element of evolving karst systems is a single isolated fracture, where a constant hydraulic head drives calcite aggre ssive water from the input to the output. Non linear dissolution kinetic s with order n = 4 induce a positive feedback by which dissol utional widening at the exit enhances flow rates thus increasing widening and so on until flow rates increase dramatically in a breakth rough event. After this the hydraulic head breaks down and widening of the fracture proceeds fast but even along its entire length un der conditions of constant recharge. The significance of modelling su ch a single fracture results from the fact that an equation fo r the breakthrough time specifies the parameters de termining the processes of early karstifi cation. In a next step the boundary condi tions for isolated fractures are varied by including different lithologies of the rock, expressed by diffe rent dissolution kinetics. This can enhance or retard karstification. Subterranean sources of CO2 can also be simulated by changing the equilibrium concentration of the solution at the point where CO2 is injected. This leads to accelerated karstification. At the confluence of solutions from two isolated tubes into a third one, mixing corrosion can release free carbon dioxide. Its eff ect to solutional widening in such a system of three condui ts is discussed. Although these simple models give interesti ng insights into karst processes more rea listic models are required. Combining singl e fractures into two-dimensional networks m odels of karst in its dimensions of lengt h and breadth under constant head conditions are presented. In first steps the Ford-Ewers' high-dip and low-dip models are simulated. Their results agree to what one expects fr om field observations. Including varying lithologies produces a variety of ne w features. Finally we show that mixing corrosion has a str ong impact on cave evolution. By this effect micro climatic cond itions in the catchment area of the cave exert significant influenc e. A common feature in the evolution of such two-dimensional models is the competition of various possible pathways to achieve breakthrough first. Varying conditions in lithologies, carbon dioxi de injection or changing hydrological boundary conditions ch ange the chances for the competing conduits. Karst systems developing at steep cliffs in the dimensions of length and depth are ch aracterized by unconfined aquifers with co nstant recharge to the water table. Modelling of such systems shows that dissolution of limestone occurs close to the water table. The widening of the fractures there causes lowering of the water table until it becomes stable when base level is reached, and a water table cave grows headwards into the aquifer. When prominent deep fractures with large aperture widths are present deep phreatic loops originate below the water table. A river or a lake on a karst plateau imposes c onstant head conditions at this location in addition to the cons tant recharge from meteoric precipitation. In this case a breakthrough cave system evolves along the water table ke pt stable by the constant head input. But simultaneously deep phreatic loops arise below it. In conclusion we find that all cave theories such as those of Swinnerton (1932), Rhoades and Sinacori (1941), and the Four-stat emodel of Ford are reconciled. They are not contradictory but they result from the same physics and chemistry under different bo undary conditions. Keywords: evolution of karst, karst aqui fers, speleogenesis, karst modelling 1. Introduction Modelling of the evolution of karst aquifers requires reduction of extremely complex systems to highly idealised simple principles. Fig.1 shows such an idealised concept of a karst system. It represents a limestone terrain with a cliff and a plateau. The limestone is dissected into blocks separated by narrow fissures and fractures. Some of those are prominent (1,2) exhibiting large aperture widths and connect water inputs at the plateau to base level at

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.2 the river flowing along the foot of the cliff. They may be located along a major bedding plane parting or a master joint. Water also infiltrates from the plateau as evenly distributed seepage from meteoric precipitation and/or from a lake or stream located at the plateau. Due to such water supplies a water table is build up, where continuous dissolutional widening of fractures is activated by water containing CO2 from the atmosphere or from CO2 in the vegetated soil covering the plateau. Fig. 1. The basic elements of a karst aquifer discussed in this work. 1: A single fracture under constant head conditions. 2: 2D fracture network under constant head conditions. 3: Vertical section of an unconfined aquifer with a constant recharge and constant head conditions applied to it. As shown, a dense network of fine fractures and a coarse network of prominent fractures are superimposed to simulate the multiple porosity character of karst aquifers. The thick dashed line WT represents the position of the water table. The hydraulic heads at the inputs and outputs are denoted as hin, h*in and hout. Modelling the evolution of such a system in space and time needs several basic ingredients: 1. How does water flow under such conditions through the aquifer? To answer this question we must know the most simple, basic concepts of laminar and turbulent flow through single fractures or conduits. 2. Dissolutional widening depends on the dissolution rates which give the amount of limestone removed from a given area in a known time. These rates usually measured in mol cm-2s-1 can be easily converted to retreat of the wall in cm/year by a factor 91.1710 !(Dreybrodt, 1988). They depend on many parameters. First of all on the chemical composition of the 223"" HOCOCaCO solution. But also on the type of flow and on the widths of the conduits, as we will show later. 3. Dissolutional widening changes the hydrological properties of the aquifer and alters the flow rates. Therefore flow and dissolutional widening must be coupled, to obtain the evolution in time. These basic ingredients should be combined as a first step in the most simple way by just considering one single conduit. Once the evolution of this basic element of the aquifer is known, combination of many single fractures into a complicated network becomes reasonable. Such networks are not just the sum of their ingredients. They exhibit more complex properties and these give insights into the overwhelming varieties of karst evolution We will lead from simple principles to more complex models, and we will try to present the most important processes active in karstification. 2. Basic Ingredients 2.1 Flow in fractures Fig. 2 illustrates an idealised fracture with plane parallel walls separated by an aperture width ao. The width of the fracture is bo and its length is L. At the left hand side a hydraulic head h injects water into the fracture which leaves at the exit at hydraulic head zero. Fig. 2. Uniform fracture with aperture width a0, width b0 and length L Calcite aggressive water is driven through it by the time-independent hydraulic head h The goal is to calculate how do the aperture widths and flow rates evolve in time due to dissolutional widening. For laminar flow, when each particle in the fluid follows a stable pathway, which never crosses the pathway of any other fluid particle, the flow rate Q [cm3s-1] and the hydraulic head h [cm] are related by / # QhR (1) a relation very similar to Ohm's law in electricity, which relates electric current I to voltage V by a resistance R The resistance R for hydrodynamic

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.3 flow is well known from the equation of HagenPoiseuille (Beek and Muttzall, 1975) and is given by 3 0012 $ % #! L R gabM (2) $ is the kinematic viscosity of water, % its density, and g earth acceleration constant. M is a geometrical factor in the order of 1 and depends on the ratio a0/b0. For wide fractures a0>>bo, M=1 For circular conduits one has a0 = b0 and M = 0.3. After dissolution has changed the profile of a fracture such that the aper ture width changes with distance from the entrance, R is then given by a summation over short tubes of length &x, where a(x) does not change in a significant way. This can be expressed by an integral as 3 012 ()()() $ % # !'Ldx R gaxbxMx (3) When flow rates exceed a threshold of discharge Q flow becomes turbulent. In this case the motion of each water particle shows large fluctuations from its average flow path and also eddies will occur. As we will see later this has a significant impact on dissolution rates. The threshold when flow becomes turbulent is given by the Reynold's number Re/ %$ # av, where v is the flow velocity in the conduit. For smooth fractures and tubes flow becomes turbulent for Re > 2000 Then the relation between head and flow rate is no longer linear and the DarcyWeissbach equation has to be applied (Dreybrodt, 1988). It reads 22 #!! gAdhh Q fLh (4) where A is the cross-sectional area and d is the hydraulic diameter of the conduit. f is a friction factor given by 10 312.51 2log 3.71 2/ $ % () *+ #", *+ -. r d f gdhL (5) r is the roughness of the wall. It is important to note that change to turbulent flow, especially in nonuniform tubes and in nets of tubes alters the distribution of heads and this puts the evolution of karst to a new stage. 2.2 Dissolution of limestone Under conditions of karst where the pH of the solution is about 7 limestone dissolves by the reaction (Plummer et al. 1978) 22 2332 /" 0,,, HOCaCOCaCOHO (I) If no carbon dioxide is present in the solution saturation is at about 10-4 mmol/cm3. If CO2 is present the following slow reaction, enhances the solubility of calcite 223 ,",/, HOCOHHCO (II) This process delivers a proton which removes the carbonate detached from the mineral by the reaction 23 3 ",",/ COHHCO (III) By this way the ion activity product 121222 3 ",COCa is kept below the solubility constant Kc of calcite. Reactions I to III can be summarized by 2 32232,",,/, CaCOHOCOCaHCO (IV) Thus stoichiometrically for each CaCO3 released from the rock one molecule of CO2 is consumed by conversion to 3 "HCO. Fig. 3 shows the equilibrium concentration of Ca++ with respect to calcite for a solution in equilibrium with a given partial pressure 2COp. Under such open conditions, where each CO2 consumed is replaced by one CO2 entering from the surrounding atmosphere into the solution the equilibrium concentration is given by 23 310.75(10.139) 34 #" 56 78eqCOmmol cTp cm (6) T is the temperature in C and 2COp is in atm. This relation is valid for 24310"9!COp atm. Mostly dissolution in early karstification proceeds under conditions closed to a surrounding atmosphere and CO2 is not replaced. In this case the CO2concentration in the solution drops by the relation :;:;22 ,,34 #" 78iCOCOCa (7) The straight line in Fig. 3 presents this chemical evolution towards equilibrium. Eqn. 6 is depicted by the lower curve in Fig. 3. The intersection between these two curves represents the equilibrium concentration, when the initial solution was free of Ca und in equilibrium with an initial 2i COp.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.4 Fig. 3. Chemical pathways of solutions in the open and closed system. The thick line represents the CO2 -Ca2+ equilibrium. Dashed lines represent the pathway of solutions in the open and closed system. For each Ca2++ released one molecule of CO2 is needed. For the open system it is replaced from the CO2 – atmosphere and PCO2 stays constant. In the case of the closed system one CO2 molecule is consumed from the solution. At the intersections between the pathways and the equilibrium curve, the corresponding equilibrium concentrations can be read. The thin solid lin es point to these calcium equilibrium concentration. Dissolution of limestone in undersaturated water is controlled by three mechanisms: 1. The detachment rate at the surface of the mineral. For karst waters Plummer et al. (1978) experimentally found the following rate 2 343()(),"#"sssFkkCaHCO (8) k3 is a constant depende nt on temperature, and k4 depends also on the CO2-concentration in the solution. The first term k3 represents dissolution, the second one is the back reaction, which depends on the activities of (Ca2+)s and (3 "HCO)sat the surface of the mineral. 2. The ions released must be transported away from this surface into the bulk of the solution, otherwise if they would accumulate there, dissolution would stop. This transport is affected by molecular diffusion. As a consequence concentration gradients build up and the concentrations at the surface are different from those in the bulk. 3. Each 3 "CO detached from the mineral requires one molecule of CO2 to be reacted to 3 "HCO. Mass conservation requires that the flux of Ca2+ from the surface must be equal to flux of Ca2+ transported into the bulk and equal to flux of CO2 towards the mineral surf ace. The surface dissolution rates are high, but CO2-conversion or mass transport may be rate limiting. CO2-conversion is a slow process. For pH between 6 and 8 it takes times up to one minute until CO2 has come to equilibrium with 3 "HCO. If water of volume V dissolves limestone from a surface area A mass conservation requires 2[] !#! dCO VAF dt (9) If the ratio V/A becomes small, due to the slow change of [CO2], which does not depend on V or A the rates will be limited by CO2-conversion. Note that for water flowing in a fracture with aperture width 2< the ratio V/A =<. On the other hand if the aperture widths 2 < becomes large the diffusi onal resistance can also limit rates. Fig. 4 shows dissolution rates for laminar flow in an aperture under closed system conditions. The numbers on the curve give the value of < in 10-3 cm. The rates are shown as a function of Ca, how they develop in free drift, when the solution approaches equilibrium. For small / <# VA, e.g. 10-4 cm the rates are very low. This is the region where CO2 conversion is rate limiting. When < increases they first increase linearly with < provided the Ca-concentra tion is kept constant. With increasing < the rates approach a limit, where they become almost independent of < in the region of 3510"! cm to 10-1 cm. This is of high relevance since this region covers the dimension of initial fracture aperture widths in the early evolution of karst. All the curves in Fig.4 can be reasonably well approximated by -21mol cm () ="34 #" 78eqFscc (10) The kinetic constant =[cms-1] is in the order of 10-5 cms-1 and is listed elsewhere (see Dreybrodt 1988, Buhmann and Dreybrodt 1985). If <>1cm mass transport becomes rate limiting and the rates are given by lim lim()() 13 = = =< #"#" ,eqDeqFcccc D (11) where lim= is the kinetic constant at the limit and D is the constant of diffusion for Ca2+ (>10-5 cm2s-1) With 5 lim310 =">! cms-1 the rates are reduced by a factor of 2 for 0.3 <> cm.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.5 Fig. 4. Dissolution rates from the theoretical model for a free drift run under the conditions of a system closed to CO2. The numbers on the curves denote the values of <, i.e, half the distance between the parallel calcite surfaces in 10-3 cm. For 3510 <"#! cm and 2110 <"#! cm the curves are identical. The uppermost curve gives the rates for fully turbulent motion and 1 < # cm. The insert in the upper right depicts the geometry of the fracture. Fig. 5. Dependence of dissolution rates on saturation ratio c/ceq. In the early state of karstification flow is laminar, and as we will see later, after short distances of flow away from the entrance the solution comes very close to equilibrium. Close to equilibrium as has been shown experimentally (Svensson and Dreybrodt 1992, Eisenlohr et al, 1997, Dreybrodt and Eisenlohr, 2000) natural calcite carbonates exhibit inhibition of dissolution rates due to impurities in the limestone (e.g. phosphate or silicates). Then the dissolution rates drop by orders of magnitude to a non-linear rate law. The dissolution rates for limestone are therefore given by 11(1/) for = #"?#eqseqFkcccckc (12) n(1/) for #"9nneqsFkcccc (13) n varies between 3 and 6 and cs between 0.7ceq and 0.9ceq. It should be noted here that gypsum rocks follow a similar rate law (Jeschke et al, 2001) and that gypsum karst therefore can be modelled the same way as karst in limestone. In the following we use as representative numbers, 1121 1410 """#! kmolcms, n = 4, cs = 0.9ceq, and 821410 """#!nkmolcms for limestone. It must be stressed here that the values of 1= #eqkc are constant only for aperture widths from about 3510"! cm to 1 mm. According to Eqn.11 they drop for 110 <"9 cm, as long flow stays laminar. The rate constants kn are properties of the mineral’s surface, solely. Due to inhibition the non linear surface rates close to equ ilibrium are so low that they become rate limiting. Fig. 5 represents Eqns. 12 and 13. The vertical line separates the region of linear kinetics (n=1) from that of the non linear kinetics (4) # n. The dotted line extends the rates of linear kinetics into the non linear region. This visualises the steep drop of inhibited non linear rates in comparison to the linear kinetics. When the flow becomes turbulent the bulk of the solution is mixed by the eddies, such that concentration gradients are levelled out. The completely mixed bulk is separated from the surface of the limestone by a diffusion boundary layer (DBL) of thickness @. Mass transport from the mineral’s surface into the bulk and vice versa is affected by molecular diffusion through this layer. The thickness of DBL depends on the hydrodynamic conditions of flow and is given by / @# aSh (14) Sh is the dimensionless Sherwood number given by (Incropera and Dewitt, 1996), 2/3(/8)(Re1000) 112.7/8(1) # ," fSc Sh fSc (15) Re is Reynolds number, fis the friction factor and Sc is the Schmidt number /() $% # ScD. For water 1000 > Sc. The resistance to mass transport through the boundary layer is determined also by conversion of CO2. When the diffusion length of 2 3 "CO-ions

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.6 (i.e. the distance a 2 3 "CO-ion travels from the mineral’s surface until it is converted to 3 "HCO ( 0.02 A> cm )) is small compared to @ then diffusion is rate limiting and the effective rate constant k drops with increasing @. If @A B than CO2 conversion becomes rate limiting because it is affected mainly in the boundary layer. Therefore k becomes constant. Detailed numbers are given by Liu and Dreybrodt (1997). The thickness of the boundary layer Eqn. 14 in our calculations was in the order of several tenth of a mm. Therefore values of 112 1410/"#! kmolcms were used. 3. Early karstification of confined aquifers 3.1 One-dimensional condu its at constant head As the first basic building block to understand early karst processes, we go back to Fig. 2, which represents some fractures, e.g. a bedding parting. Fig. 6 represents a part of this fracture between x and ,& xx, where x is the distance from the entrance. Fig. 6. Mass conservation in the part of the fracture between x and x+dx The widening rate can be calculated by use of Eqns. 12 and 13 everywhere in the fracture if the concentration c(x) along the fracture is known. This can be found as follows. The amount of calcium dissolved from the walls per time unit must be equal to the difference of the amount of calcium entering at x to that leaving at x + dx, also related to one time unit. From this we obtain the mass balance equation (())()()() #!# FcxPxdxvxAxdcQdc (16) where A(x) is the cross-sectional area at x, P(x) the corresponding perimeter and v the velocity. Due to the continuity of flow the constant flow rate Q through the fracture is given by ()() !# vxAxdcQdc. Eqn.16 can be solved if one uses the dissolution rates given in Eqns.12 and 13. (Dreybrodt 1996; Dreybrodt and Gabrovšek 2000; Gabrovšek 2000). The solutions are (17) 11s 1 /(1) 4s()(1)exp(/) for x
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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.7 active all along the entire fracture thus maintaining a profile of parallel planes. For this we can use the analytical solutions of Eqn.16. Widening in the fracture is even and given by 2(,) G #! da FLt dt (18) with (19) /(1) 3 0 3 ,0(,)11 () A"" #() CH DD #!"!, *+ EI *+ DD FJ -.n nn s n eqntcLa FLtk cat (2) For most relevant cases, when the inflowing solution has a calcium concentration below xs the summand 1 in the bracket can be safely neglected and Eqn. 18 can be integrated after inserting Eqn. 19. The result is 1 21 0()(1/)" ,#"n n BatatT (20) with 011 221(,0) G #!! ,Ba n T nFL (21) Inserting the equation for F(L,0) we obtain the relation: (22) 12 21 2 1 1 1 1 011124(1) 221 $ G%, " "() () "" #!! *+ *+ *+ -. -.n n n n n Bn eqnLn Tk naghc The evolution a(t) is illustrated by Fig. 8. Initially a slow increase is observed. But increasing a(t) creates increasing dissolution rates at the exit and vice versa. This positive feedback loop then finally creates the steep increase in a(t) and correspondingly in flow rate Q through the fracture. One point is of utmost importance. The breakthrough time TB in Eqn. 22 contains the parameters which determine the time scale of early karstification. TB increases drastically with decreasing aperture width (exponent (2n1)/(n-1) ,), it increases also with L (exponent 2/1 nn). The dependence on head h is less drastic (exponent /(1) "" nn). A most important chemical parameter is ceq. For bare karst areas, when rainwater with atmospheric 2COp enters into the fractures ceq under closed condition is at about 10-4 mmol/cm3, in contrast to 3210"#!eqc mmol/cm3 in vegetated areas. This increases breakthrough times by at least one order of magnitude. Fig. 8. Evolution of fracture apertu re widths at the exit as given by Eqn.14. The dotted vertical line represents the pole of the function a ( t). The above considerations are only approximative. Numerical solutions of the equations are more exact. Fig. 9 shows the result of digital modelling of our standard case. Its parameters are listed in Table 1. TABLE 1 Parameters used for the model of the evolution of a single fracture Description Name Unit Initial or Standard value Aperture width a0 cm 0.02 Fracture length L cm 105 Fracture width b0 cm 100 Hydraulic gradient i 0.05 Order of non linear kinetics n 4 Linear kinetics constant k1 2/ molcms 11410"! Non linear kinetics constant kn 2/ molcms 8410"! Concentration of calcium c 3/ molcm Switch concentration cs 3/ molcm 61.810"! Equilibrium concentration ceq 3/ molcm 6210"! Viscosity of the solution $ / gcms 21.210"! Density of the solution % / gcms 1

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.8 Fig. 9a in a logarithmic plot depicts as a function of time the flow rate through an initially plane parallel fracture with initial parameters as given in Table 1. The flow rate exhibits a slow increase that is enhanced in time until it is drastically accelerated to such an amount that they exceed the water available at the surface. At this breakthrough time TB the hydraulic head breaks down, and the initial phase of laminar flow through the fracture is terminated. Fig. 9b represents the evolution of the aperture widths along the fracture for various times depicted by points 1 to 9 in Fig. 9a. The widening rates (2 G F) are shown by Fig. 9c. In the beginning where the fracture widths at the entrance (2-4) are below 1mm the rates are given by Eqns. 12 and 13. They are maximal at c=0 Later when the entrance widens, molecular diffusion becomes rate-limiting (cf. Eqn. 11) and the rates close to the entrance drop. But they rise downstream as the aperture widths decrease. Finally when 4th order nonlinear kinetics becomes active rates are determined by Eqn.13 and they drop again. After breakthrough when flow is turbulent (dashed lines) the rates become high and even along the conduit. Details of this dynamic behaviour are given in the literature. (Dreybrodt, 1988, 1990, 1996; Dreybrodt and Gabrovšek, 2000). Fig. 9d shows the saturation ratio c/ceq along the fracture. At the beginning the ratio rises steeply until c=cs where the higher kinetics become active causing a further slow increase. With increasing flow rate the region of first order kinetics ( c
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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.9 becomes shorter. Consequently flow velocity increases and dissolution rates at the exit are enhanced. After breakthrough flow becomes turbulent and the head distribution changes (dashed curves) and the decline along the fracture becomes more even. This is also important for 2 or 3dimensional nets, because after breakthrough this new distribution of heads determines the further evolution of the aquifer. So far we have considered that the inflowing solution is allogenic water far away from equilibrium with respect to calcite. We now ask the question: What happens if the inflowing solution is closer to equilibrium. To answer this we consider Eqn. 17, which gives the dissolution rates along a plane parallel fracture. When cin approaches ceq, the penetration length An approaches infinity and the rates become almost constant along the entire fracture, close to the maximal value at the input. Further increase in flow, will further increase the value of An without significant influence to the rates. The reason is that the second term in the bracket of Eqn.17 becomes small with respect to the summand 1. Therefore one expects that the feedback loop gradually looses influence when cin approaches ceq. Fig. 10. The evolution of flow rate in the standard fracture for various values of c0/ceq as denoted in the figure. Fig. 11. Conceptual model of a heterogeneous fracture with change of kinetic properties ( n,kn) or equilibrium concentration at x=KL. This is shown by Fig. 10. It depicts the flow rates versus time for a single conduit with differing input concentrations. Curves 1 to 5, with input concentrations equal or below 0.95ceq, clearly exhibit breakthrough behaviour. When cin comes closer to ceq as in curve 6 ( cin = 0.975ceq) flow rates increase steadily because dissolutional widening becomes even along the entire tube at maximal possible rate (1/) !n nineqkccduring the entire time of evolution. In that case karstification becomes very slow, but it may create horizons of increased permeability, which later, when geological conditions change may be utilised for conduit growth. Such horizons have been suggested by Lowe and Gunn (1997). 3.2 The heterogeneous fracture In the examples discussed up to now all chemical parameters ceq kn n and k1 were assumed as constant along the entire fracture length. This is highly idealistic. In nature karst conduits may have sections of varying lithology and therefore with varying values of n and kn. Subterranean sources of CO2 could enter into fractures and increase the solubility of limestone reflected by an increase of ceq. In this section we will shortly discuss the influence of such new conditions on breakthrough time (Gabrovšek et al., 2000, Gabrovsek, 2000). Fig. 11 depicts the heterogeneous standard fracture, which exhibits different properties, such as different dissolution kinetics or differing ceq, in its entrance part for ? xKL and the exit region for 9 xKL. First we assume differing dissolution kinetics with 182410/"#!nkmolcms and 14 # n in the first half of the fracture (K = 0.5) and 262410/"#!nkmolcms and 26 # n in the remaining part. Fig. 12 shows the numerical results for the standard fracture with 14 # n and 26 # n (a,b,c) and for the reverse case, when 16 # n and 24 # n (d,e,f). Both fractures show breakthrough behaviour, but at quite different time scales. For n1 < n2 the dissolution rates drop drastically when the solution encounters the border of changing lithology. Since the solution is close to saturation the dissolution rates, shown by Fi g. 12b, remain constant from then on. After the time of 45ky the first half of the fracture has opened up to such an extent that the hydraulic head remains close to the input head and the head loss from h to base level zero is along the second half. With increasing widths of this half

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.10 Fig. 12. a) Evolution of flow rates in time for the standard fracture with n1 = 4, n2 = 6 and K = 0 5 ( b,c) Profiles of dissolution rates and aperture widths for n1 = 4, n2 = 6 and K = 0 5 plotted at 0.1, 45.8, 452.1, 599.4, 649.1, 665.5, 670.9,672.5, 673 and 673.1ky marked from 1-10 respectively. d) Evolution of flow rate in time for the standard fracture with n1 = 6, n2 = 4 and K = 0 5. e,f) Profiles of dissolution rates and aperture widths for n1 = 6, n2 = 4 and K = 0 5 at 0.1, 3.3, 5.8, 7.3, 8.2, 8.7,9, 9.2, 9.3 and 9.4 ky marked from 1-10 respectively. flow rates increase such that the concentration at the boundary becomes sufficiently high to switch on the feed back loop causing breakthrough. Breakthrough time is greatly influenced by the low dissolution rates at the exit of the fracture and is substantially longer than in the homogeneous standard fracture (cf. Fig. 9). For 16 # n and 24 # n (Fig. 12 d-f) the dissolution rates are boosted up when the solution meets the border. Due to the significantly lower dissolution rates in the entrance half the concentrations remain sufficiently high such that breakthrough is dominated by the kinetics with 24 # n in the exit part. Compared to the first case the initial rates at the exit are higher by two orders of magnitude. Therefore breakthrough time is only 10ky, even less than that of the standard fracture. These examples show that differing lithologies have a strong impact on karstification and cannot be neglected. More details on this issue are reported by Gabrovšek (2000), and by Dreybrodt and Siemers (2000). We now turn to the case where ceq changes step like by injection of subterranean carbon dioxide. Fig. 13 shows the numerical results of the dissolution rates along the standard fracture for two cases. a) K = 0.25 and b ) K = 0.75. At these positions 6210"#!eqc mol/cm3 increases by 6210"&#!eqcmol cm-3. With an increase of eqc the rates are enhanced, where CO2 is injected into the fracture and consequently dissolutional widening becomes faster there. The fracture profiles are depicted by Fig. 14 for both cases. Fig. 15 illustrates the breakthrough behaviour for the cases K=1 (standard fracture), K = 0.25 (Fig. 14a), K = 0.75 (Fig. 14b), and K = 0.9 In all cases of CO2 injection breakthrough times are reduced. Further details are published by Gabrovšek et al. (2000). It should be noted here that also injection of solutions with different chemical composition can increase ceq e.g. mixing corrosion, and reduce breakthrough times.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.11 Fig. 13. Dissolution rates for the point CO2 input at 0.25 # K (dashed lines) recorded at 0.1, 6.5, 9.4, 10.13, 10.35, 10.39 ky (labeled from 1a to 6a respectively) and the for the input at 0.75 # K (full lines) recorded at 0.1, 6.93, 10.19, 11.07, 11.27, 11.31 ky (labeled from 1b to 6b respectively). Fig. 14. a) Profiles of aperture widths for the input of 2CO at 0.25 # K, recorded at the times given in Fig. 13, labeled from 1-6. b) Profiles of aperture widths for the input of 2CO at 0.75 # K, recorded at the times given in Fig. 13, labeled from 1-6. Fig. 15. Breakthrough times for the standard fracture with 2CO-input at to # xKL for various values of K. 3.3 Evolution of caves in 2-dimensional nets of fractures To approach further to reality we consider a confined aquifer consisting of a limestone bed dipping downward, which is dissected by narrow initial fractures. This is shown schematically by the model domain in Fig. 16. The bed dips from left to right. The left hand side has input points, where the head and the chemical composition of the inflowing water can be defined individually for each input Fig. 16. Modelling domain of its 2-dimensional fracture network. The length of the domain is 2km, the width 500m. Fracture spacing is 10m x 5m. Each fracture has uniform initial aperture width, which is determined by a log-normal distribution over the entire set of fractures. There are three inputs on the left side of the domain at h=50 m and a series of outputs on the right hand side at h=0.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.12 point. The left hand side is at base level, i.e. h=0. The aperture width of each fracture can be assigned individually. By this way it is possible to model the statistical properties of fracture widths. In the following we use a log-normal distribution of fracture widths with average 00.02 # acm and 0.01 K # cm. It is also possible to use bi-modal distributions to simulate a coarse net of wide fractures embedded into a dense net (continuum) of narrow fractures to account for the double-porosity properties of karst aquifers (Sauter, 1992). To model the evolution of conduits in such nets one has to calculate the distribution of hydraulic heads at all the nodes in the net and from these the flow rate through each fracture. Furthermore one must know the chemical composition of the solution at each node to couple the flow to the chemical dissolution model in each single fracture. Technical details of this procedure are described by Siemers and Dreybrodt (1998) and by Gabrovšek (2000). The parameters of the model domain are 00.02 # acm, 0.01 K # cm, h = 50 m, the length L of the domain is 2km and its width 500 m. The upper and the lower boundaries are impermeable.We first apply our modelling efforts to the high-dip model of Ford and Ewers (see Ford and Williams, 1989). Fig. 17 shows the evolution with 3 inputs with equal chemical composition and equal heads. Fig. 17a depicts the situation after 8000 years. Several conduits have grown downstream. At approx. 17590 (Fig. 17b) years the upper has reached base level and breakthrough occurs. Therefore the constant head boundary condition at this input is replaced by a constant flow rate at that input. Thus the hydraulic head in this conduit drops to values close to base level, as can be seen from the pressure isolines in the figures. This directs flow from the conduits which have not reached base level towards the master conduit. Therefore they integrate into a common system. When such a conduit reaches the leading one breakthrough occurs (Fig. 17c and d) and again the constant head condition at the input is changed to constant flow. The distribution of flow rates within the network is illustrated by Fig. 19. The rates are depicted in units of Qmax which is the maximal flow rate through a fracture which occurs in the net. Note that Qmax increases in time. As one visualises from Fig. 19a flow is widely distributed over the net initially. As the conduits grow down head the flow is radiating into the net from their tips. After breakthrough flow is concentrated in the mayor conduits (Figs. 19b-d). The evolution of the flow rates through the aquifer is depicted by Fig. 18 (curve1), which also shows the subsequent break through events. This scenario shows a competition Fig. 17. Distribution of aperture widths and hydraulic head isolines at various times for the domain presented in Fig.16. Hydraulic head at the left hand side is set to 50m and 0 at the right-hand side. Water at all inputs is in equilibrium with pCO2=0.05atm. All fractures smaller than 0.05 cm have been omitted. The bar code indicates the numerical values. between the growing conduits. Prior to breakthrough each of these propagates down head. The breakthrough time for such single 1-dimensional channels is therefore similar as in Eqn. 22, where a0 has to be replaced by th e average initial aperture

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.13 width along the conduit and L must be replaced by its total length. Since these values, due to the statistical distribution of fracture widths are different for each of the evolving conduits the one with shortest breakthrough time wins the race. Siemers and Dreybrodt (1998) have shown by a sensitivity analysis to the various parameters that Eqn.22 is applicable for the breakthrough time in twodimensional nets under consta nt head conditions. It should be noted here, that the breakthrough time for a defined conduit, where length and average a0 are exactly known, can be significantly shorter, than that of an isolated one-dimensional conduit. The reason for this is that a conduit embedded into a net of fractures looses fluid into the net (see Fig. 19a,b). Therefore under constant head conditions more aggressive solution enters into this fracture and enhances dissolutional widening (Romanov et al. 2002a). As a next example we discuss the low-dip-model by Ewers and Ford. In this case two additional inputs are placed in the central section of the aquifer, both with equal heads and equal chemical composition of the solutions as those at the left hand side of the domain. This is illustrated by Fig. 20. After 2000 years (a) channels from these inputs have propagated downstream. No conduits can grow from the inputs at the left hand boundary, because these have only negligible head difference to the central region. After 2470 years (b) the conduits exhibit a breakthrough and integrate together similar as in the high dip case. This changes the hydraulic head distribution and conduits from the inputs at the left hand side grow towards these channels. Each row of inputs behaves similar as the high-dip model, and an integration of their conduit systems grows upwards (Figs. 20 c,d). Fig. 18. Evolution of total flow through the network in time for all 2-dimensional cases. Insert shows the sequential breakthroughs in the high-dip model (Figs. 17 and 19). Fig. 19. Flow pattern in the network at the same times as in Fig. 17. Line’s thickneses in the bar code represent the values of Q/Qmax, where Qmax is the flow rate through the fracture with highest flow. At each breakthrough event the constant head condition at the corresponding input is replaced by constant flow. This causes changes in the distribution of hydraulic heads, which direct growing channels towards conduits of constant flow rate, or in other words head zero. The evolution of flow rates is depicted by Fig. 18, curve 2.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.14 Fig. 20. Aperture widths and hydraulic heads at various time for the low-dip model. Same settings as in Fig.17, but two inputs at h=50m are added to the middle of the domain. During the early evolution of a karst aquifer prior to breakthrough the concentrations of the solutions within the karst aquifer are very close to saturation (cf. Fig 9c, single fracture). If the chemical composition of the waters at the inputs are different with respect to initial 2COp and these waters mix somewhere within the aquifer mixing corrosion will boost up dissolution rates. To show the impact of mixing corrosion to the evolution of karst we have changed the chemical composition at the central input in the high-dip model shown by Fig. 17 and have left everything else unchanged. Fig. 21 shows the result. In contrast to the evolution with equal input chemistry in Fig.17, where the outer conduits grow simultaneously, now the middle conduit has advanced significantly. The outer conduits have stopped growth at 12000y. The reason for this is that during the early phase of evolution water from the two outer inputs injected into system of narrow fractures mixes with water from the central input. Therefore dissolutional widening by mixing corrosion increases the average fracture widths to such an extent that the central conduits gain sufficient advantage to reach base level first. After breakthrough growth of the outer channels is directed towards the central one. The evolution of flow rates is shown in Fig. 18, curve 3. This example shows that changes in the chemical parameters are of utmost importance. Changes in vegetation can alter the chemical composition of the input waters and can divert the evolving cave patterns to a high extent. By this way climate can effect cave evolution. Details on the effect of mixing corrosion are reported by Gabrovšek and Dreybrodt (2000) and Romanov et al. (2002b).

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.15 Fig. 21. Aperture widths and hydraulic heads at different times for the case with mixing corrosion. Same settings as in Fig. 17, but water at upper and lower inputs has reduced content of CO2 (pCO2=0.03) PCO2=0.03atm PCO2=0.05atm PCO2=0.03atm

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.16 Fig. 22. High-dip model (Fig. 17) with a CO2 input. pCO2 in the region marked in Fig. 22a is set to 0.05atm.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.17 Fig. 23. High-dip (Fig. 17) model with changing lithology. As marked in Fig. 23a, the kinetic order changes within the network from 4 to 6 and back to 4.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.18 When subterranean sources of CO2 either by volcanic origin or by microbial activity are present in the aquifer dissolutional widening is enhanced in the regions neighbouring such inputs. Therefore these regions favour the evoluti on of conduits growing to their direction. Fig. 22 shows such an example. Here a point source of CO2, which raises ceq by ten percent is located in the middle of aquifer. From this input points due to the enhanced dissolution rate a channel grows down head, which is joined by the upper conduit. Comparison of the evolution compared to that in Fig. 17 shows the strong influence of subterranean CO2 sources. This is also reflected by the evolution of the flow rates shown in Fig.18, curve 4. Further details have been publis hed by Gabrovšek and Dreybrodt (2000). Varying lithology can change the parameters of the dissolution kinetics (Eisenlohr et al. 1999). Fig. 23 shows the evolution of the aquifer, when its middle part is composed of limestone with 626,410/"##!nnkmolcms whereas the outer parts remain unchanged with n=4. Due to the altered kinetics dissolutions rates in the middle part drop, when they are encountered by the solution and channels grow only in the input part with n=4. When such a channel has reached the border between the two lithologies channels grow down head until a first reaches base level and breakthrough occurs. The flow rates for this case are shown by Fig.18, curve 5. 4. Modelling unconfined aquifers in the dimension of length and depth So far we have discussed confined aquifers subjected to constant head conditions. By use of this modelling concept the evolution of caves in the dimension of length and breadth can be studied. Now we project one dimension to depth to investigate cave evolution in length and depth. In this projection of reality new boundary conditions must be envisaged. The aquifer need no longer be confined and a water table may be present. Furthermore input conditions of constant recharge must also be considered. In the following we will present the conceptual frame for such models. Fig. 24 shows a vertical section of a limestone plateau down cut by a valley on its right-hand side. The massif is dissected by fractures of various initial aperture widths. Prominent fractures are shown in the massif. The enlargement shows a net of fine fractures that are evenly distributed throughout the aquifer. Fig. 25 shows a simplified version of the cross-section representing the modelling domain discussed in this chapter. Th e model plateau is rather small, 200m long and 30m high. The hydraulic conductivity represented by the rectangular net of fine fractures with aperture widths 0.005 > acm is about 10-7cm/s Prominent fractures with aperture widths in the order of few tenth of a millimetre can be incorporated into the system of narrow fractures. A recharge of 450mm/year is evenly distributed at the surface of the plateau and ”offered” to the aquifer. The left-hand side, the base and the lower right-hand side are assumed impermeable. This is marked by thick solid lines in Fig.25. The hydraulic head hout at base level is set to zero. The cliff on the right represents a seepage zone where the water leaks from the aquifer. Constant head conditions can be applied additionally on the top, e.g. an allogenic river flows over the massif. Table 2 presents the parameters used in the model and their typical values. Fig. 24. Schematic representation of a cross-section through the limestone plateau. At the top a combination of constant recharge and constant head conditions can be applied. Thick grey lines show impermeable boundaries. The lower picture is an enlargement showing the parameters of the fracture systems.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.19 TABLE 2 Parameters of unconfined model Description Name Unit Standard or Initial value Aperture width a0 cm 0.006 Aperture width of prominent fractures ap cm 0.02 Length of vertical and horizontal fractures Lv, Lh cm 50, 200 Hydraulic conductivity K ms-1 10-7 Annual recharge Q mm/year 450 Fig. 25. The model domain with its boundary conditions. Our model implies an unconfined aquifer, i.e. an aquifer with a water table (WT) which divides a saturated phreatic and an unsaturated vadose zone (see Fig. 1). Recharge is infiltrating through the surface and the vadose zone down to the phreatic zone at the WT. The position of the WT depends on recharge. To obtain the flow through the fractures, the position of the WT must be known, since it defines the boundary conditions for flow and separates the saturated zone from the unsaturated one. The position of the WT and the height of the seepage face is calculated by the following procedure: 1. An initial guess for the WT is assumed, f.i. the surface of the plateau. 2. A recharge defined by precipitation is equally distributed to the points of the assumed WT. 3. The heads of all the net-points below and at the assumed WT are calculated with the boundary conditions defined by the assumed WT and seepage face, i.e. h=z. 4. The heads of the points on the WT are checked for the boundary condition. Their head must be, within a given error, equal to their elevation. If this condition is valid for the point, the WT is kept there, otherwise the WT is either shifted to the point above, if h > z or to the point below if h < z Thus a new approximation for the WT is obtained. 5. Procedure 1-3 is iterated until all the points on the assumed WT fulfil the condition h = z. Once the WT and the seepage zone are obtained, the flow through the fractures in the phreatic zone is calculated and the transport-dissolution model is applied. This is done in the same way as described for confined networks. Chemical parameters used are those given in Table 1. During percolation through the vadose zone, the solution already attains some saturation state. This is taken into account by taking c0 between cs and 0 97 ceq, so that the initial concentration rises linearly with the depth of the water table. The choice of the parameter c0 is rather arbitrary. It influences the evolution of an aquifer, but does not change the results conceptually. A broad sensitivity analysis has not yet been done. Details are given in Gabrovšek and Dreybrodt (2001) and in Gabrovšek (2000). 4.1 The evolution of fine fracture systems. No prominent fractures: Constant recharge The aquifer consists of system of narrow fractures. No prominent fractures are present in the modelling domain. We assume a constant recharge of 450mm/year evenly distributed to the surface of the plateau. Fig.26a shows the situation at the onset of karstification. The phreatic zone is indicated by fat fracture lines, and the vadose zone by thin interrupted ones. In this way the position of the WT is clearly presented. The colours show the fracture aperture widths increasing from dark blue to red as denoted in the figure. a0 is the initial width of the vertical fractures. Fig.26b shows the situation after 5ky. The WT has dropped due to the increasing fracture widths in the aquifer. After 10ky the WT reaches the lowest output fractures. This is presented in Fig.26c. By continuous dissolution along the base level a conduit develops and grows headwards (Fig.26d) until it reaches the left boundary after 20ky. Inspection of the colours in Fig.26d reveals

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.20 Fig.26. Evolution of an aqui fer with evenly spaced fine fractures and constant recharge. Distribution of fracture widths after 50y ( a ), 5ky ( b ), 10ky ( c ) and 15ky ( d ). The colours represent the widths of the fractures in units a ( t ) /a0, where a0 is the initial width of vertical fractures. Fractures designated by full squares represent the phreatic zone, those by open angles the vadose zone. By this way the water table is clearly presented. that the hydraulic conductivity increases by about 2 orders of magnitude, leaving a highly permeable vadose zone as is observed in nature. Dissolutional widening is most active close to the water table at all times, since the solution quickly approaches equilibrium when penetrating into the net. Therefore close to the WT a narrow region of higher permeability is established which attracts flow. In Fig.26a a small light blue fringe indicates this zone. Fig. 27. a) Flow rates in units of Q/Qmax at 5ky. The fracture with the highest flow rate has a ratio of Q/Qmax equal to 1. Most of the flow is concentrated to the permeable fringe at the top of the phreatic zone. Confer to Fig.26. b ) Dissolution rates in units F/Fmax where Fmax = 122410/"! molcms corresponding to bedrock retreat of few 310/"cmyear. The maximal dissolution rates are active close to the water table and drop rapidly with depth. Later the fringe becomes wider and is composed of fractures with apertures up to 16 a0. With increasing time the water table drops, leaving behind the vadose region of increased conductivity. The phreatic zone below still has low hydraulic conductivity. In such an aquifer most of the flow is directed along the water table. After 15ky the conduit at the WT has reached an average width of 1cm. The fracture widths in the vadose zone above are between 0.01-0.1cm, with the wide fractures close to the final WT. Fig.27a shows the flow rates through the fractures at 5 ky and clearly illustrates (green, yellow, red fractures) that the flow is r estricted close to the water table. Dissolution rates shown in Fig.27b also exhibit a maximum close to the water table and drop significantly below. As the water table drops dissolution becomes active in the lower parts of the aquifer. Once the water table has reached a stationary position dissolution stays active close to it and large conduits can grow. This corresponds to the ideal water table cave in the model of Swinnerton (1932) and Rhoades and Sinacori (1941) which requires high and even fissuring of the rock.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.21 Combined conditions: constant recharge and constant head Often constant head boundary conditions and those of constant recharge exist simultaneously. This for example is the case when allogenic rivers are present. Fig.28 shows such a case where a constant head equal to the elevation at the upper left boundary is imposed (see Fig. 28a). Other parameters are equal to those in Fig.26. Constant recharge is offered to the aquifer everywhere else. Fig.28a shows the water table and the distribution of the fracture aperture widths 1ky after the initial state. In the region of constant recharge the water table drops towards the seepage face, whereas in the condition of constant head it coincides with th e surface at the platform. Dissolution occurs only in a small banded region close to the WT as illustrated by Figs.28a, b, c. In the constant head region, the water table cannot drop below the surface, theref ore the head difference along the WT increases. A permeable fringe along the WT offers an effective pathway draining the water from the constant head region to the output. The feedback mechanism along this pathway leads to the breakthrough at 2.1ky. A wide zone of high conductivity has been created which carries flow from the constant head area. Fig.29 shows the total discharge as a function of time. The arrows indicate the flow rates at the time steps from Fig.28. This resembles a typical breakthrough behaviour such as observed for onedimensional conduits or for nets under constant head conditions. To obtain some information on the widths of the fractures Fig. 30 depicts the aperture width profiles along a vertical and horizontal crosssection as indicated by the arrows in Fig. 28c. 4.2 Aquifers with a net of prominent fractures To create a more realistic karst aquifer we now superimpose a net of prominent fractures to the system of fine fractures. The following procedure is used: LM Divide the net of fine fractures into a coarse net of 5 by 5 dense fractures LM With a random procedure assign to each fracture of this coarse net the aperture width ap, of a prominent fracture. If the random number chosen for each fracture is smaller than p the fracture has an aperture width ap, else its aperture is that of the fine net fractures. Fig. 28. Evolution of a fine fracture system with combined constant recharge and constant head conditions at the top. The constant head region is marked in figure a. The WT is fixed due to the constant head. Dissolution is mainly active in the fringe close to the water table. The arrows on figure c denote the position of the profiles presented in figure 30. Fig. 29. Evolution of flow rate through the network in Fig.28. Arrows indicate times when aperture distributions in Fig.28 are presented.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.22 Fig. 30. Distribution of the fracture widths for a horizontal and vertical cross-section of the aquifer as indicted by arrows in Fig. 28c. The numbers on the curves indicate the time of the aquifer evolution in ky. a) Widths of the vertical fractures in the ver tical cross-section for various times indicated in ky. The region of maximal widths corresponds to the red area in Fig.28. b) Widths of the horizontal fractures in the horizontal cross-section for various times. The small peak at about 30m corresponds to the vertical channel, which develops at the border between constant head and constant recharge regions. Note that the increase of ap erture widths accelerates in time. This initial scenario is similar to the approach of a double continuum model of Clemens et al. (1996, 1997, 1999) but it avoids calibration parameters, which are difficult to specify. It is also close to the approach of Kaufmann and Braun (1999, 2000) whomodel the initial aquifer by a superposition of a prominent fracture net within a rock matrix with homogeneous initial conductivity. The most important difference in our approach is that dissolutional widening is regarded in both parts of the aquifer, whereas Clemens et al. and also Kaufmann and Braun disregard dissolution in the dense fractures or in the matrix respectively. Boundary conditions of constant recharge As pointed out dissolution in our model is active in the prominent fractures as well as in the continuum of narrow fractures. The question arises, where does permeability arise in the competition of gaining flow from the surface. This is shown by Fig.31. It depicts an aquifer with a network of prominent fractures (p=0.8) with aperture widths of 0 04 cm To get a more pronounced pattern, recharge is increased to 700 mm/year Fig.31a represents the fracture widths after 30k y Fig.31b depicts the flow rates and consequently the flow path at that time. In addition to the conduit growth along the water table, deeper phreatic loops form along prominent fractures below the water table. Constant head and constant recharge conditions The boundary conditions for the case in Fig.32 are the same as in Fig.28: constant head at the left-hand upper side and constant r echarge on its right hand side. Fig. 32a shows the widths close to the onset of the evolution after 200 years. After 1ky (Fig. 32b) a complex net of conduits has developed along the prominent fractures. The region of constant head becomes connected to the area of constant recharge by increasing hydraulic conductivity, caused by both, widening of the fine fractures and also connection to the prominent ones. Consequently the water table rises. Close to the water table a region of higher conductivity connects the prominent fractures to the seepage face. This change of conductivity and hydraulic heads enhances the evolution of the conduits along the prominent fractures. Fig. 31. Aquifer with a percolating net of prominent fractures ( ap = 0 04 cm ). Annual precipitation is 700 mm/year. a) Distribution of fracture widths after 30ky. Conduits grow along the base level and along the phreatic loops. Note the change in the colour code with respect to other figures. The widths designated by red are above 0.5cm. The dashed line depicts the initial WT. b) Distribution of flow rates at 30ky.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.23 Fig. 32. Evolution of an aquifer with combined constant recharge and constant head boundary conditions and a percolation network of prominent fractures appended to the dense fracture network. The initial aperture width of prominent fractures is 0.02cm (light blue). Other parameters are the same as in Fig. 28. a) 0.2ky, b) 1ky, c) 1.2ky. After 1.2ky breakthrough occurs (Fig. 32c), with prominent fracture aperture widths on the order of a few millimetres. To illustrate the distribution of fracture widths as they evolve in time, Fig. 33 depicts these along a horizontal cross-section as indicated by an arrow in Fig. 32c. The widening of the fracture accelerates by feedback and consequently the discharge through the aquifer shows the characteristic breakthrough behaviour. This event terminates the early evolution of the aquifer. The constant head condition breaks down and must be replaced by constant recharge. Flow becomes turbulent. Nevert heless, the complicated pattern of vertical and hor izontal conduits and a high permeability region close to the spring will design the future structure of the mature karst aquifer. It should be stressed at that point, that constant head conditions are crucial for the evolution of such complicated structures as shown in Fig.32. Fig. 33. Profiles of horizontal fractures along a crosssection as indicated by an arrow in Fig. 32c. Another point of interest is that two processes proceed simultaneously: The evolution of conduits along the prominent fractures and creation of increasing permeability in the continuum, as shown by the red region in Fig 32c. Such behaviour cannot be obtained if dissolutional widening is restricted to the prominent fractures. 4.3 Time dependent boundary conditions: down cutting of the cliff In nature the boundary conditions are changing during the evolution of a karst aquifer. The precipitation rate Q the chemical parameters of the inflowing solution and also the hydrological boundary conditions may alter. All these variations can be applied to the model presented. We present a case where the base level of an aquifer is down cut during the evolution. In a first scenario we assume that a ”sudden” incision of a valley lowers the base level. This is presented in Figs.34 a and b. The model is the same as used in Fig.31 but the position of base level is kept at 15m during the first 10ky and than it is down cut to 25m immediately. In Fig. 34a, which shows the situation at 9.8ky, a water table cave has developed and as in Fig.31 the system of conduits evolves below the base level. After the down cutting the WT adopts to the new base level in a short time. This is presented in Fig.34b. After 11ky a new water table cave is already evolving. Between both base levels the hydraulic conductivity is relatively small since the WT has dropped fast due to the phreatic loops which have evolved prior to down cutting. Probably more realistic than the step down cutting is a gradual down cutting. Such a scenario is shown on Fig.35. The initial base level is almost at the top of aquifer and is being lowered in steps of two nodes (1m) every 5ky. The mechanisms are similar to those

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.24 Fig. 34. Evolution of an aquifer with a net of prominent fracture under constant rechar ge condition with step down cutting of the base level. Fig. 35. Evolution of an aquifer with a net of prominent fracture under constant rechar ge condition with gradual down cutting. for the step down cutting. The formation of phreatic conduits below base level forces the WT to adapt promptly to the temporary base level. Fig.35a and b show the aquifer at 6ky and 12ky, respectively. The vadose zone exhibits a relatively high permeability since the region of fast widening at the WT was gradually lowered together with the slow down cutting. A slower down cutting produces higher permeability in the vadose zone. 5. Conclusion The basic element in modelling the early evolution of karst is a single 1-dimensional fracture which is widened by dissolutional attack of calciteaggressive water. Under c onstant head conditions flow increases slowly at the beginning but then is enhanced dramatically and breakthrough occurs. The breakthrough times depend on the initial aperture width, the length and the hydraulic head acting on the entrance of the fracture. But they also depend on chemical parameters ceq, kn, n, which determine the dissolution kinetics. When these chemical parameters vary within a single fracture, f.i. ceq by injection of subterranean carbon dioxide or kn and n by varying lithology of the rock comprising the fracture, breakthrough times are changed significantly. Thus even for one-dimensional systems a great variety of boundary conditions determines the time scale of karst evolution. A further step to model karst is the combination of one-dimensional fractures into a two-dimensional net either on length and breadth or on length and depth. By this way two-dimensional projections of karst aquifers can be modelled. In such nets due to the interaction of flow one-dimensional fractures can inject flow into this net and therefore enhance dissolutional widening by increased input flow of calcite aggressive solution. Consequently a prominent fracture embedded into a net of narrow fractures can exhibit breakthrough much earlier than if it were isolated. Furthermore boundary conditions in networks become more complex. Differing chemical compositions of input waters at various input regions cause mixing corrosion, where those waters mix, deep in the aquifer. This creates increased permeability which attracts further flow and directs conduits towards such regions. Therefore solely by chemical boundary conditions of the input waters karst aquifers with or without mixing corrosion (i.e. with or without differing chemical composition of the inflowing water) will develop entirely different. A large impact to the evolution of cave systems is caused also by different rock lithologies within the aquifer. Finally the hydrological boundary conditions, i.e. constant head inputs and their location and/or constant recharge to the surface of the aquifer exert significant influence. Due to our lack of knowledge on all these various boundary conditions and their change in time it is not possible to explain a specific cave system by modelling. The aim of karst modelling is to understand the processes operating in the evolution of karst aquifers.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.25 References Beek, W.J. and Mutzall, K.M.K. 1975. Transport Phenomena. New York, Wiley. Breznik, M. 1998. Storage reservoirs and deep wells in karst regions. Rotterdam ,A.A. Balkema, 251. Buhmann, D. and Dreybrodt, W. 1985. The kinetics of calcite dissolution and precipitation in geologically relevant situations of karst areas: 2. Closed system Chem. Geol. 53, 109-124. Clemens, T., Hckinghaus, D., Sauter, M., Liedl, R. and Teutsch, G. 1997. Modelling the genesis of karst aquifer systems using a coupled reactive network model. In: Hard Rock Geosciences. Colorado, IAHS Publ. 241, 3-10, Clemens, T., Hckinghaus, D., Sauter, M., Liedl, R. and Teutsch, G. 1996. A combined continuum and discrete network reactive transport model for the simulation of karst development. In: Calibration and Reliability in Groundwater Modelling. Colorado, IAHS Publ 237, 309-318. Clemens, T., Hckinghaus, D., Liedl R. and Sauter, M. 1999. Simulation of the development of larst aquifers. The role of epikarst. Int. Journal of Earth Sciences, 88, 157-162. Dreybrodt W. 1988. Processes in karst systems Physics, Chemistry and Geology. Springer Series in Physical Environments 4, Berlin New York, Springer, 288 p., Dreybrodt, W. 1990. The role of dissolution kinetics in the development of karstification in limestone: A model simulation of karst evolution. Journal of Geology 98, 639-655 Dreybrodt, W. 1996. Principles of early development of karst conduits under natural and man-made conditions revealed by mathematical analysis of numerical models. Water Resources Research 32, 2923-2935. Dreybrodt, W., and Gabrovšek, F. 2000. Dynamics of the evolution of a single karst conduit. In: Klimchouk, A., Ford, D.C., Palmer, A.N. and Dreybrodt, W., (Eds.), Speleogenesis: Evolution of karst aquifers. Huntsville, Nat. Speleol. Soc., 184-193. Dreybrodt, W. and Siemers, J. 2000. Cave evolution on two-dimensional networks of primary fractures in limestone. In: Klimchouk, A., Ford, D.C., Palmer, A.N. and Dreybrodt, W., (Eds.), Speleogenesis: Evolution of karst aquifers. Huntsville, Nat. Speleol. Soc., 201-211. Eisenlohr, L., Madry, B. and Dreybrodt, W. 1997. Changes in the dissolution kinetics of limestone by intrinsic inhibitors adsorbing to the surface. In: Proceedings of the 12th Int. Cong. of Speleology La Chaux de Fonds, Switzerland, Vol II, La Chaux de Fonds, Switzerland, 81-84, Eisenlohr L., Meteva, K., Gabrovšek, F. and Dreybrodt, W. 1999. The inhibiting action of intrinsic impurities in natural calcium carbonate minerals to their dissolution kinetics in aqueous H2O-CO2 solutions. Geochimica et Cosmochimica Acta 63, 989-1002. Ford, D.C. and Williams, P.W. 1989. Karst geomorphology and hydrology. London, Unwin Hyman, 601 p. Gabrovšek, F. 2000. Evolution of early karst aquifers: From simple principles to complex models. Ljubljana, Zaloba ZRC, 150 p. Gabrovšek, F., Menne, B. and Dreybrodt, W. 2000. A model of early evolution of karst conduits affected by subterranean CO2 sources. Environmental Geology 39, 531-543. Gabrovšek, F. and Dreybrodt, W. 2000. The role of mixing corrosion in calcite aggressive H2O-CO2CaCO3 solutions in the early evolution of karst aquifers. Water resources research 36, 11791188. Gabrovšek,F., and Dreybrodt, W. 2001. A model of the early evolution of karst aquifers in limestone in the dimensions of length and depth. J. Hydrol 240, 27-34. Incropera, F. and DeWitt, D. 1996. Fundamentals of Heat and Mass Transfer. 4rd edition, New York, John Wiley & Sons, 912 p. Kaufmann, G. and Braun, J. 1999. Karst aquifer evolution in fractured rocks. Water Resources Research 35, 3223-3238. Kaufmann, G. and Braun, J. 2000. Karst aquifer evolution in fractured, porous rocks. Water Resources Research 36, 1381-1391. Liu Z. and Dreybrodt W. 1997. Dissolution kinetics of calcium carbonate minerals in H2O-CO2 solutions in turbulent flow: The role of the diffusion boundary layer and the slow reaction 223HO +COHHCO,N,. Geochimica et cosmochimica acta 61, 2879-2889. Lowe, D.J., and Gunn, J. 1997. Carbonate speleogenesis: An inception horizon hypothesis. Acta Carsologica 26 (2), 457-491. Plummer, L.N., Parkhurst, D.L., and Wigley, T.M.L. 1978. The kinetics of calcite dissolution in CO2 systems at 25oC to 60oC and 0.0 to 1.0 atm CO2. American Journal of Science 278, 179-216.

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W.Dreybrodt and F.Gabrovšek / Speleogenesis and Ev olution of Karst Aquifers 1, January 2003, p.26 Rhoades, R. and Sinacori, M.N. 1941. The pattern of ground-water flow and solution. Journal of Geology 49, 785-794. Romanov, D., Dreybrodt, W. and Gabrovsek, F. 2002a. Interaction of fracture and conduit flow in the evolution of karst aquifers. In: Martin, J.B., Wicks, and Sasowsky I.D. (Eds.), Proceedings of the Symposium on Karst Frontiers: Florida and Related Environments. KWI Special Publication No. 7. Romanov, D., Gabrovšek, F. and Dreybrodt, W. 2002b. The impact of hydrochemical boundary conditions on the evolution of karst aquifers in limestone terrains. Submitted to Journal of Hydrology. Sauter, M. 1993. Double porosity models in karstified limestone aquife rs: field validation and data provision. In: Gltekin, G., Johnson, I.A., (Eds.), Hydrogeological Processes in Karst terraines. IAHS Publication 207, 261-279. Siemers, J. and Dreybrodt, W. 1998. Early development of karst a quifers on percolation networks of fractures in limestone. Water resources research 34, 409-419. Swinnerton, A.C. 1932. Origin of limestone caverns. Geological Society of America Bulletin 34, 662693. Svensson, U. and Dreybrodt, W. 1992. Dissolution kinetics of natural calcite minerals in CO2 -water systems approaching calcite equilibrium. Chem. Geol. 100, 129-145.



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Speleogenesis and Evolution of Karst Aquifers Th e Virtu al Scien tific Jou r n al ww w.s p el e oge nesi s.i n f o Karst breakdown mechanisms from observations in the gypsum caves of the Western Ukraine: Implications for Subside nce Hazard Assessment Alexander Klimchouk (1)* and Vjacheslav Andrejchuk (2) (1)* Senior Scientst, Institute of Geological Scien ces, Natl. Academy of Sciences of Ukraine, P.O.Box 136, Kiev-30, 01030 Ukraine. E-mail: klim@speleogenesis.info (2) Professor, Department of Earth Science, University of Silesia, ul. Bedzinska 60, 41-200 Sosnowiec, Poland. E-mail: geo@wnoz.us.edu.pl Corresponding author Abstract The term karst breakdown is employ ed in this paper to denote the totality of processes and phenomena of gravitational and/or hydrodynamic destruction of the ceiling of a karst cavity and of the overlying sediments. It refers not only to the existence o f a surface subsidence (collapse) feature but, first of all, to the "internal" (hidden in the subsurface) structures that precede d evelopment of a surface form. This study reports and discusses the results of direct mapping and examination of brea kdown structures in the gypsum karst of t he Western Ukraine, at the level of their origin, i.e. in caves. The accessibility of numerous laterally extensive maze cave syste ms in the region provided an excellent opportunity for such an approach which made it possible to examine the relationship between breakdown structures and particular morphogene tic or geological features in caves, and to reveal stages of breakdown developmen t. It is found that breakdown is initiated mainly at specific speleogenetically or geologically "weakened" localities, which class ify into a few distinct types. The most of breakdowns, which are potent to propagate th rough the overburden, relate with the outlet cupolas/domepits that represent pl aces where water had discharged out of a cave to the upper aquifer during the period of trans verse artesian speleogenesis. Distribution of brea kdown structures does not correlate particularly well with the size of the master p assages. Several distinct mechanisms of breakdown development are revealed and most of them proceed in several stages. They are guided by speleogenetic, geological and hydrogeological factors. The study confirms that a speleogenetic approach is indi spensable to the understanding of breakdown pre-requisites and mechanisms, as well as for eventual subsidence hazard assessment. Direct observations in caves, aimed both at speleogenetic investigation and breakdown charac terization on regional or si te-specific levels, should be employed wherever possible. Keywords: gypsum karst, speleogene sis, karst subsidence, subsidence hazard assessment, Western Ukraine 1. Introduction The term karst breakdown is used in this paper to denote the totality of processes and phenomena of gravitational and/or hydrodynamic destruction of the ceiling of a karst cavity and of the overlying sediments. Use of this more general concept avoids potential misconceptions that commonl y arise from the ambiguous use of terms "collapse" and "subsidence" in the literature. It has an additional advantage in that it does not refer to the existence of a surface subsidence (collapse) feature and includes "internal" (hidden in the subsurface) processes and phenomena that precede the appearance of a surface form. Karst breakdown is complex, consisting of a number of processes, with components developing in various combinations, either simultaneously or sequentially. Some components may dominate during certain stages of the breakdown development, whereas others may occur throughout the entire process. The karst breakdown mechanism is understood here as a combination of specific component processes in a regular sequence, and their development in time and space. An understanding of the karst breakdown mechanisms is crucial to subsidence hazard assessment, prediction and management in karst terraines. A set of component agencies and a shifting of the breakdown process proper (i.e. breakdown mechanism) depends on many factors and conditions, a combination of which is referred to here as settings. Analysis of the available literature on the subject suggests that the mo st important factors that determine settings are: 1) th e presence and structure of the overburden, 2) lithological (geotechnical) properties of individual units in the cover, 3) hydrogeological conditions (especially piezometric levels and hydraulic gradients), and 4) degree of karstification and characteristics of the primitive initiating cavities.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.2 Fig. 1. Location of the gypsum karst of the Western Ukraine (A) and zonation of the region according to evolutionary types of karst (B). Zones of different karst types are labeled by Roman numbers: I = the gypsum is entirely denuded, II = entrenched karst, III = subjacent karst, IV = deep-seated (confined) karst. Numerous accessible and laterally extensive cave systems in the Western Ukrainian gypsum karst provide excellent opportunities for direct examination and mapping and examination of breakdown structures at the level of their origin, i.e. in caves. Such observations and surveys are indispensable for an adequate understanding of conditions favorable to breakdown initiation and of mechanisms favorable to their development. 2. Geological and hydrogeological background to gypsum karst development The Miocene gypsum sequence is widespread on the southwestern edge of the eastern European platform, along the Carpathian Foredeep, where it occupies over 20,000km 2 Gypsum stretches from the northwest to southeast for more than 300km as a belt ranging from several kilometers to 40 to 80km wide (Fig. 1A). It is the main component of the Miocene evaporite formation that girdles the Carpathian folded region to the northeast, from the Nida river basin in Poland across the Western Ukraine and Moldova to the Tazleu river basin in Romania. Most Miocene rocks along the platform margin rest on the eroded terrigenous and carbonate Cretaceous sediments. The Miocene succession comprises deposits of Badenian (Tortonian) and Sarmatian age. The Lower Badenian unit, beneath the gypsum, includes mainly carbonaceous, argillaceous and sandy beds (30-90 m thick) adjacent to the foredeep, and these grade into rocks of calcareous biohermal and sandy facies (10-30 m thick) towards the platform interior. The overlying gypsum bed is variable in structure and texture. Most commonly it grades from microcrystalline massive gypsum at the lower part through variably grained bedded gypsum in the middle to giantocrystallline rock in the upper horizon. A layer of evaporitic and epigenetic limestone, locally called "Ratynsky", commonly overlies the gypsum, ranging from half a meter to more than 25 m in thickness. The gypsum and the Ratynsky limestone comprise the

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.3 Tyrassky Formation which is overlain by the Upper Badenian unit represented either by argillaceous and marly lithothamnion limestones and sandstone beds or, adjacent to the foredeep, by marls and clays of the Kosovsky formation. The latter grades upward into the Lower Sarmatian clays. The total thickness of the capping marls and clays ranges from 40-60 m in the platform interior to 80-100 m and more in the areas adjacent to the regional faults that separate the platform edge from the foredeep. There is a distinct trend in the depth of the gypsum occurrence, position of the overall denudation surface within the Miocene succession and the depth of erosional entrenchment in the direction across the gypsum belt, from the platform interior towards the foredeep. The Tyrassky Formation dips 1 to 3 o towards the foredeep and is disrupted by block faults in the transition zone. To the south and south-west of the major Dniester Valley, large tectonic blocks drop down as a series of steps, the thickness of clay overburden increases, and the depth of erosional entrenchment decreases. Along the tectonic boundary with the foredeep the Tyrassky Formation drops down to the depth of 1000 m and more. This variation, the result of differential neotectonic movement, played an important role in the hydrogeological evolution of the Miocene aquifer system and resulted in the differentiation of the platform edge into the four zones (Andrejchuk, 1984, 1988; Klimchouk et al, 1985; Klimchouk and Andrejchuk, 1988; Klimchouk, 1996, 2000). The gypsum was entirely removed by denudation within the 1-st zone, but other three zones represent the distinct types of karst: entrenched, subjacent and deep-seated (Fig. 1-B). The gypsum bed is largely drained in the entrenched karst zone, is partly inundated in the subjacent karst zone and remains under artesian confinement in the deep-seated karst zone. In hydrogeologic terms the region represents the southwestern portion of the Volyno-Podolsky artesian basin (Shestopalov, 1989). The Sarmatian and Kosovsky clays and marls serve as an upper confining sequence. The lower part of the Kosovsky Formation and the limestone bed of the Tyrassky Formation form the original upper aquifer (above the gypsum) and the Lower Badenian sandy carbonate beds, in places along with Cretaceous sediments, form the lower aquifer (below the gypsum), the latter being the major regional one. The hydrogeologic role of the gypsum unit has changed with time, from initially being an aquiclude, intervening between two aquifers, to a karstified aquifer with well-developed conduit permeability (Klimchouk, 2000). Regional flow is from the platform interior, where confining clays and the gypsum are largely denuded, toward the large and deep Dniester Valley and the Carpathian foredeep. In the north-west section of the gypsum belt the confined conditions (zone IV) prevail across its entire width. In its wide south-east section the deeply incised valleys of Dniester and its left tributaries divide the Miocene sequence into a number of isolated deeply drained interfluves capped with the clays (Podol'sky area). This is the entrenched karst zone (zone II) where most of the explored, presently relict maze caves are located. To the south-southeast of the Dniester (Bukovinsky area) the gypsum remains largely intact and is partly inundated (the subjacent karst zone III). Further in this direction, as the as the depth of the gypsum occurrence below clays increases and entrenchment decreases, the Miocene aquifer system becomes confined (the deep-seated karst zone IV). In this zone the groundwater flow pattern includes a lateral component in the lower aquifer (and in the upper aquifer but to a lesser extent) and an upward component through the gypsum in areas of potentiometric lows, where extensive cave systems develop as evidenced by numerous data from exploratory drilling. 3. Speleogenesis Fourteen large caves over 1km in length are known in the region. Most of these caves are presently relict. They are located north of the Dniester, within the 2nd zone (entrenched karst). Two other large caves, Zoloushka and Bukovinka, are in the Bukovinsky sub-region, near the Prut River, generally in the area of artesian flow within the Miocene aquifer system (4th zone) but within local, exceptionally uplifted blocks, where entrenchment into the upper part of the gypsum caused unconfined (water table) conditions to be established during the Holocene. All the large gypsum caves in the region are mazes arranged into laterally extensive multi-storey networks, which have developed along vertical and steeply-inclined fissures. Interconnecting passages form lateral twoto four-storey systems that extend over areas of up to 1.5km 2 Such areas, termed here cave fields, are defined by drawing an arbitrary boundary closely enclosing the passages on a cave map. Significant morphological parameters of the caves are summarized in Table 1. Figs. 5A, 10, 11 and 12 illustrate some typical cave patterns. Optimistychna Cave, with more than 214km of surveyed passages, is the longest gypsum cave and the second longest cave of any type known in the world. The Western Uraine contains the five longest known gypsum caves in the world, accounting for well over half of the total known length of gypsum caves on the Earth. By area and volume the largest caves are Ozernaja (330,000m 2 and 665,000m 3 ) and Zoloushka (305,000m 2 and 712,000m 3 ), followed by Optimisticheskaja Cave (260,000m 2 and 520,000m 3 ). The absolute parameters of cave systems change as exploration progresses. Specific parameters are more informative. Specific volume (the cave volume/length ratio, which is in fact the average area of passage cross-section) characterizes an average size of cave passages in a cave system. For the caves of the region this parameter ranges from 1.7 (Gostry Govdy Cave) to 8.0 (Zoloushka Cave) m 3 /m. The average value for the region is 3.9 m 3 /m. Passage network density is characterized conveniently by using the ratio of cave length to a unit area of the cave field (km/km 2 ). This parameter varies within the region from 118 (Verteba Cave) to 278 (Jubilejnaja Cave) km/km 2 with an average value of 164 km/km 2

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.4 TABLE 1 Parameters of large caves and cave fields in the Western Ukraine No Cave name* Development, m Specific volume m 3 /m Density of passages, km/km 2 Areal coverage, % Cave porosity, % The Podolsky sub-region 1 Optimistychna 214000 2.8 147 17.6 2.0 2 Ozerna 111000 6 150 44.6 5.0 3 Mlynki 25000 3.3 141 37.6 3.4 4 Kristalna 22000 5.0 169 29.2 6.0 5 Slavka 9100 3.7 139 27.6 3.4 6 Verteba 7800 6.0 118 34.7 12.0 7 Atlantida 2520 4.5 168 30.0 4.0 8 Ugryn 2120 3.8 177 33.3 5.7 9 Jubilejna 1500 2.3 278 37.0 4.0 10 Komsomolska 1240 2.1 177 24.3 3.0 11 Dzhurinska 1130 2.4 126 17.8 2.0 The Bukovinsky sub-region 12 Zoloushka 92000 8.0 142 48.4 3.8 13 Bukovinka 2400 2.5 120 21.5 4.4 14 Gostry Govdy 2000 1.7 270 17.5 4.0 Totals 493820 Averages 3.9 164 29.5 4.5 The names are given here according to the Ukrainian spelling. In many other publications Russian spelling is common, where most of names ended here with "a", end with "-skaja" or "aja". The availability of detailed morphometrical data on caves and host rock bodies allows calculation of areal coverage and cave porosity parameters (fractions of the total area and the volume of the rock within a cave field occupied by passages). The areal coverage varies from 17.5 to 48.4 %, the average value being of 29.5 %. Cave porosity varies from 2 to 12 %, with an average value of 4.5 %. Maze caves in the region have been developed (and are presently developing in the 4th zone) under confined conditions, due to upward transverse groundwater circulation between the sub-gypsum and supra-gypsum aquifers (Klimchouk, 1990, 1992, 1996, 2000). Such a flow pattern is characteristic of potentiometric low areas, related to topographic lows (valleys),which commonly coincide with zones of enhanced fluid conductivity created within the capping clays by tectonic or stratigraphical discontinuities. Overall discharge from artesian aquifer systems occurs in such areas. Under conditions of transverse circulation in a multi-storey artesian system, all available fissures in the gypsum, which hold similar positions within analogous flow paths, enlarge at comparable rates because of the availability of dispersed aggressive recharge from below and suppressed hydraulic competition due to constrained outflow. This behavior generally favors the development of maze cave structures, but the actual conduit arrangement in any given locality depends upon the initial fissure pattern. Three major components can be distinguished in the cave systems based on shape, arrangement and hydrologic function of cave mesoforms during the main (artesian) speleogenetic stage (Figs. 2 and 3): 1. Feeder channels, the lowermost components in a system: vertical or sub-vertical conduits through which water rose from the sub-gypsum aquifer to the master passage networks. Such conduits are commonly separate but sometimes they form small networks at the lowermost part of the gypsum, along the top of the underlying bed. The feeder channels join master passages located at the next upper level and are scattered rather uniformly through their networks. 2. Master passages: horizontal passages that form laterally extensive networks within certain horizons in the middle part of the gypsum bed. They received dispersed recharge from numerous feeder channels and conducted flow laterally to the nearest outlet feature. 3. Outlet features: domes, cupolas and vertical channels (domepits) that rise from the ceiling of the master passages to the bottom of the overlying bed. They discharged water from cave systems to the overlying aquifer.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.5 Fig. 2. Main morphogenetic features of maze cave systems in the Western Ukraine shown at their hydrologic functionality. 1 = feeder channels, 2 = master passages, 3 = outlet features The Western Ukrainian maze caves provide the most outstanding and unambiguous evidence for the transverse artesian speleogenetic model. Artesian speleogenesis in the Podolsky sub-region took place during the Late Pliocene through Early Pleistocene when the overall maze structure of caves became established. Breaching of artesian confinement and further incision of the valleys during the Middle Pleistocene caused substantial acceleration of groundwater circulation within the Miocene artesian system. The majority of passage growth, as well as breakdown formation, probably occurred during this transitional period. Where the water table was established in the gypsum for a prolonged time, further widening of passages occurred. Eventually, with the lowering of the water table below the lower gypsum contact, cave systems in the entrenched karst zone became entirely fossilized. Cave development under confined or semi-confined conditions continues today within the zones of deep-seated and subjacent karst (the 4th and 3rd zones). 4. Speleological observations of the breakdown formation and development: methods and criteria The accessibility of numerous laterally extensive cave systems in the Western Ukrainian gypsum karst provides an excellent opportunity for direct mapping and examination of breakdown structures at the level of their origin. This allows almost all breakdown structures, which have evolved within a cave field, to be mapped, including those that are still hidden within the coverbeds and not manifested on the surface. Such mapping makes it possible to investigate the relationship of breakdown structures with particular morphogenetic and geologic features in a cave and to reveal stages of breakdown development. The state (quasi-equilibrium or non-equilibrium) of a breakdown structure can be judged and a degree of its propagation toward the surface through the cover (the height of a breakdown column the depth of a migrating void below the surface) can be determined. Fig. 3. Examples of typical morphogenetic features in the caves: 1 = feeder channels, Mlynki and Ozerna caves; 2 = master passage, Dzhurinska cave; 3 = outlet features, Slavka and Optimistycha caves. Photo by A.Klimchouk. Together with detailed data on lithostratigraphy, thickness and hydrogeology of the overburden, this reveals the breakdown mechanisms and facilitates subsidence hazard assessment for the respective areas with a precision and certainty unachievable by the approaches of conventional engineering geology. Such investigations make it possible to test the validity and adequacy of various indirect approaches to subsidence hazard assessment and the assumptions on which such approaches are based.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.6 The following features were identified and mapped as breakdown structures in the caves (Fig.4): 1) Any outlet features (domes, cupolas and domepits in cave passages) indicating considerable breakout in the vault. Breakdown is identified by predominantly gravitational morphology of the vault and by the presence of disarticulated fragments of bedrock and coverbed materials beneath it. 2) Breakdown taluses in cave passages consisting of the fallen bedrock and coverbed material. Depending on the initial dome diameter and the distance of upward stoping, such taluses can plug an access to the breakout cupola and separate the migrating void from the cave. Breakdown structures in the Western Ukrainian gypsum karst normally develop in a number of stages through a prolonged period of time. The multi-stage development is determined by the stratified nature of the overburden which has varying lithological, geomechanical and hydrogeological properties of individual units. The stages are identified from a position of a cupola or a migrating void within the cover that, if not directly observed, can be inferred in most cases from the size, shape and composition of breakdown taluses. The state (quasi-equilibrium or non-equilibrium) of a breakdown structure can be additionally determined from the presence or absence of signs of recent activity in a breakdown talus (water seepage or flow, dampness of sediments, signs of creep or extrusion, etc.). Fig. 4. Breakdown structures in the caves. B reakout cupolas: A and B = in Mlynki cave, C = in Slavka cave. Breakdown taluses: D = in Kievljanka cave, E = in Mlynki cave, F = in Zoloushka cave, G = in Verteba cave. Photo by A.Klimchouk.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.7 Such surveys were performed in several caves developed in different geological and hydrogeological settings and representing different morphological character: in Zoloushka Cave (subjacent karst settings), Mlynki, Slavka and Verteba caves (entrenched karst settings). 5. Breakdown development in the subjacent karst zone: Zoloushka Cave and the Dankivsky Collapse Zoloushka Cave (Fig. 5A) is the third longest gypsum cave in the world with 92km of passages mapped since 1976 when few entrances were opened in the face of an active gypsum quarry. The cave area lies generally in the confined karst zone, although in some of the more uplifted tectonic blocks (where the gypsum was partially entrenched by the nearby major Prut Valley during the Holocene) the groundwater surface is some 2 to 3m below the gypsum top. The quarry operation and accompanied groundwater withdrawal since 1950s caused the water table to further drop 17 to 19m below the gypsum top and brought about considerable transformations in the karst system. The cave was thoroughly studied in various aspects (Andrejchuk, 1984, 1988, 1999; Andrejchuk and Korzhik, 1984) and provides an excellent playground for examination of karst breakdown mechanisms. 5.1. Local settings Local geomorphological and geological settings are depicted on Fig. 5, B and C. The gypsum in the cave area has a thickness of 23 to 25m, being overlain by the microcrystalline grey light brown Ratynsky limestone, up to 1m thick. The Kosovsky Formation, 5 to 60m in thickness (depending on the local relief), spreads over the cave area. It comprises mainly argillaceous sediments of grayish-blue color, with some minor sandstone and limestone beds in its lower part. The clays consist predominantly of montmorillonite (up to 38%) and hydro-illite (25 to 30%), and are massive in the lower part of the formation and thinly-bedded in the upper part. The main geotechnical characteristics of the clay are as follows: natural humidity 17 to 18%, plasticity index 28, density 2.1 g/cm 3 skeletal volume weight 1.77 g/cm 3 porosity 35.6%. Above the Kosovsky Clays the Quaternary alluvium of the upper (III to IV) Prut terraces is present, comprising sandy-gravel (immediately above the Kosovsky Clays) and loam sediments. The loams, ranging from a few to 19m in thickness, are light and porous. The soil layer, 0.5 to 1.2m thick and rich in humus (2.7 to 6.4%), lies on the top. The gypsum rests on the sands and marls of the Lower Badenian (3 to 4m), which in turn overly the eroded Cretaceous limestones and sandstones. Together they form the presently unconfined aquifer, which also includes the lower part of the gypsum. Under natural conditions the aquifer discharged to the Prut River through the terrace sediments. During the quarrying stage a depression cone due to water withdrawal from the quarry deformed the groundwater surface in the cave area. The Quaternary aquifer is also present, being perched on the Kosovsky Clays, although in the cave area it is increasingly drained by breakdown structures that disrupt the clay succession. Fig. 5. A = The map of Zoloushka cave (courtesy of the Chernovitsky Speleological Club), B = Geomorphological map of the area, C = Geological cross-section across the cave field. 5.2. Dewatering of the cave The quarry that opened the cave started at the end of 1940s. Since then, groundwaters have been continuously abstracted from the quarry and a drawdown cone has formed around it. In the beginning the withdrawal rate was rather modest amounting about 20 to 50m 3 /hour. When the quarry had deepened up to 8 to 10m, the pumping rates increased to 100 to 500m 3 /hour. Since the mid-1960s, with the cutting of the third quarry bench to the depth of 18 to 22m, groundwater inflow reached 700

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.8 to 800m 3 /hour and this rate was maintained until nowadays. Before the quarry, the groundwater level had been situated at about 2 to 3m below the gypsum top, some 1 to 2m below the ceiling of Zoloushka's upper storey passages. The groundwaters circulated slowly toward the Prut River and discharged through the alluvium. They contained considerable amount of H 2 S and dissolved solids (3.0 to 4.5 g/L). With the start of the operations, the quarry became the drainage focus. Within the drawdown cone that expanded up to several kilometers in diameter, groundwater flow changed to radial, with a considerable increase of flow rates, decrease of TDS content (up to 1.9 to 2.6 g/L) and H 2 S degassing. The lowering of the piezometric surface and dewatering of the cave had progressed most during the 1960s. In the first period of the cave exploration (1976 to 1978) passage bottoms were covered by "fresh" wet slippery clay, progressively desiccating and shrinking in the following years with overall decrease in volume and the formation of characteristic crack patterns. The floor level in most passages has lowered by 1 to 2m and the volume of passages has increased by 25 to 35% since the time of the first exploration. This apparently contributed to an activation of the pre-existing breakdown structures that rested on the cave fill. The lower storey of the cave remains inundated, being located below the water table. 5.3. Cave morphology Zoloushka Cave is a labyrinth of horizontal passages occurring in two storeys. The upper storey consists predominantly of large passages (average width and height are respectively 2.8m and 3.0m; specific volume is 8.0m 3 /m) with ceilings located 1 to 3m below the gypsum top (Fig. 6A and B). Their cross-sections are oval, rhomb-like or hemispherical. Numerous solution domes (1 to 5m in diameter) in the passage ceilings expose the overlying Ratynsky limestone bed. Such domes were outlets for the water to the overlying aquifer during the period of transverse artesian speleogenesis. In some areas large closely spaced passages coalesce laterally, with only small pillars remaining in between them (Fig. 6C). This is due to horizontal notching by preferential dissolution at the water table during the Holocene (Fig. 6B and C). In this way some quite large (15,000 to 30,000m 3 ) chambers were formed. In areas where the level of clay filling lowers, it is possible to observe 3 to 10m-deep rift-like extensions in the bottoms (Fig. 6D), otherwise obscured by the filling. Thus, the entire cross-sections commonly have "keyhole" shapes, with the width of the rift part from 0.3 to 3.0m. The lower storey of the cave, still inundated and explored only in fragments, lies along the bottom of the gypsum. It is connected with the upper level through large pits (feeders), whose morphology indicates "ascending" hydraulic communication during the formation period. The cave map (Fig. 5A) displays only the upper storey passage network. Sixteen morphological regions are distinguished in the cave, according to characteristic passage size and the structural peculiarities of the patterns. The differences in passage sizes are illustrated by the specific volume parameter varying between regions from 5.1 to 16.1m 3 /m (Table 2). Fig. 6. Zoloushka cave morphology. See the text for explanations. Photos by A.Klimchouk and V.Kisseljov. 5.4. Breakdown structures About 70% of the maze has been covered with a special mapping of breakdown structures (BS). At least 700 breakdown structures were found in the cave, over 630 of which were mapped and documented according to the criteria outlined in the previous section. This gives an average density of breakdown structures for the whole cave field of about 1800 per km 2 Breakdown initiation. A great majority of breakdown structures initiate and develop where solution domes and cupolas have exposed the bottom of the overlying Ratynsky Limestone bed to the cave. Speleogenetically, such domes and cupolas represent the outlet features through which the water discharged from the cave during the period of transverse artesian speleogenesis. The Ratynsky bed is less than 1m thick and is normally rather densely fissured and brecciated. It falls readily when exposed from below by the outlet features, giving rise to the formation of BS. In places where the Ratynsky bed is coarsely fractured and exposed by occasional block fall-ins, it provides an effective support for the ceiling. It is likely that most of BS in the cave was initiated during the period of transition from confined to unconfined conditions, due to the loss of buoyant support. Mechanisms of the breakdown development. Among the structures examined, none was found to display signs of a single massive collapse of the cave roof and overburden. All BS demonstrate prolonged multi-stage development. This is determined mainly by the stratified nature of the coverbeds. Five to six distinct stages are distinguished in the breakdown formation (Fig. 7).

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.9 TABLE 2 Morphometric characteristics of the regions of Zoloushka Cave, and breakdown distribution by region Averages Name of the region Length, m Width, m Height, m Passage area, n1000m 2 Passage volume, n1000m 3 Specific volume, m 3 /m Area of cave field, m 2 Number of BS BS density per km 2 1 Privkhodovoj 5,600 2.7 2.2 15.2 33.8 6.0 39.2 105 2,679 2 Zabludshikh 4,330 2.4 2.7 10.5 28.7 6.6 30.3 42 1,386 3 Perspectiv 1,237 3.5 3 4.4 13.3 10.8 8.7 31 3,580 4 Chernovitskij 3,919 3.7 3.2 14.4 45.7 11.7 27.4 68 2,479 5 Majsky 1,424 2.1 2.7 3 8.1 5.7 10.0 15 1,505 6 Central'ny 7,880 2.6 2.9 20.3 59.7 7.6 55.2 14 254 7 Zapadny 5,015 2.8 2.7 13.6 38.4 7.7 35.1 26 741 8 Anakonda 3,891 2.7 2.9 10.5 30 7.7 27.2 50 1,836 9 Vesely 5,317 2.2 2.3 11.8 27 5.1 37.2 59 1,585 10 Metropoliten 2,337 3.7 3.7 8.6 37.6 16.1 16.4 5 306 11 Ozerny 4,228 3.1 3.8 12.9 49.1 11.6 29.6 38 1,284 12 Gotichny 4,091 3 4.5 12.4 56.2 13.7 28.6 63 2,200 13 Vostochny 4,769 3.1 3.6 14.7 53.1 11.1 33.4 60 1,797 14 Dal'nevostochny 2,414 2.3 3.4 5.6 19 7.9 16.9 21 1,243 15 Kamtchatka 1,084 2.4 2.4 2.6 6.3 5.8 7.6 40 5,271 16 Geochimichesky 5,000 2.6 2.5 16 30 6.0 35.0 TOTAL 62,536 176.5 536 637 AVERAGE 2.8 3.0 8.8 1,876 Correlation between BS number and the variables 0.31 0.05 -0.16 0.36 0.2 -0.14 The preparatory stage is not considered as a part of the breakdown mechanism proper, although it creates distinctive morphogenetic features, namely the outlet features (dome shafts, domes and cupolas) favoring breakdown occurrence. This stage coincides with the late artesian speleogenetic stage, and is marked by the growth in the area of the Ratynsky bed exposures at the vaults of outlet domes and cupolas. The first stage of breakdown-proper is the failure of the Ratynsky bed into the cave and the formation of a breakout (gravitational) cupola in the lower part of the Kosovsky Formation (up to 1 to 2m above the gypsum top). It is, therefore, the stage of active development. A breakdown pile consisting of the limestone blocks and some clayey debris is formed beneath a cupola. The first stage itself can be short, probably one fall-in event in many cases, but it is followed by a prolonged period of relative stability (second stage). The second stage is marked by gradual upward stoping of a breakout cupola through the Kosovsky Clays. Destruction of the material at the cupola vault occurs as slab and chip breakdown, rarely as block breakdown (i.e. fallen rock masses span more than one bed; White and White, 2000). The fallen argillaceous material forms distinct breakdown taluses (cones) that can vary in volume from a few to many tens of cubic meters. This stage can span quite long periods, probably in the order of thousands to tens of thousands of years. Its duration depends upon the local properties of the Formation and its thickness in a given cave region; the latter varies between a few to 60m, depending on the local relief. Many breakdown structures at this stage still provide access from the cave to a stoping cupola, although when the structure reaches some height, it gets separated from the cave by a breakdown pile. Because of this, one can estimate further migration of the void and passage of BS to the next stage only on the basis of the composition of the talus material at the base. If it contains some admixture of sandy-gravel material, then the BS has reached the Quaternary bed and passed to the next stage. The third stage begins when a migrating void has reached the sandy-gravel bed of Quaternary alluvium. Two distinct mechanisms of the further development are revealed (Figs. 7, A and B). Which one occurs in a given locality depends on the presence of groundwater in the Quaternary sandy-gravel bed. Mechanism A predominated in the past, probably during the period commencing some tens of thousands years ago when the Miocene aquifer lost its confinement (i.e. when breakdown processes had intensified for the first time due to the loss of buoyant support). This continued until some

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.10 Fig.7. Mechanisms of the breakdown formation in Zoloushka Cave. 40 to 30 years ago, when breakdown development intensified again due to the start of quarrying and pumping, and related lowering of the water table and subsequent transformations in the cave system. Mechanism B predominates now, when the Quaternary aquifer is largely drained across most of the area, as breakdown structures and exploration boreholes provided numerous points of vertical leakage through the Kosovsky Formation. Mechanism A: When a migrating void reaches a sandy-gravel bed that contains groundwater, hydrodynamic component processes, such as liquefaction, piping and erosion become involved in the overall breakdown development and become predominant during the third and fourth stages. Breakdown of the last remaining portion of the Kosovsky Clays, along with some sandy-gravel material, causes liquidation of a void at the top of the BS because of sand liquefaction and the formation of a zone of thinning that extends laterally along the sandy horizon as a reversed wide-angle cone (the fourth stage; see Fig. 7A). The vertical breakdown structure enables leakage of the water from the aquifer, accompanied by further material removal by piping and erosion. Wetting of the clayey column causes its settlement down into the cave, further increasing the zone of thinning. All this leads to sagging of the overlying loam sequence, with the eventual appearance of surface deformation in the form of gradual subsidence. The rate and the depth of surface subsidence depend on the intensity of the leakage and piping, and on the rate of the erosion and settling of the breakdown column. It is important to stress that Mechanism A results in a gradual subsidence type of the surface deformation, not collapse. Analysis of historical data and large-scale topographical maps for the pre-quarrying period supports the view that gradual subsidence was the prevailing type of deformation in the recent past. However, Mechanism A is still operative in a few breakdown structures where Quaternary beds still host lenses of groundwater. This is indicated by active filtration along some breakdown columns. There have been occasional direct observations of drastic activation of a breakdown cone, with apparent settlement and extrusion of wet clays down into the cave and release of a considerable amount of water within few hours. On the surface the related pre-existing gentle subsidence was reactivated, with the formation of fresh concentric cracks up to 2m deep and up to 0.3m wide. Mechanism B: This occurs where the Quaternary sandy-gravel horizon is drained and does not contain water. This situation has become increasingly predominant in the cave area since the start of quarrying and related groundwater abstraction from the main Miocene aquifer. This caused reactivation of pre-existing breakdown structures and formation of new ones that, together with numerous exploration boreholes, created a closely spaced pattern of leakage points from the perched Quaternary aquifer. This eventually caused the aquifer to drain throughout most of the area. The differences between the mechanisms start from the third stage (see Fig. 7B), which begins when a stoping void reaches the sandy-gravel bed. This stage signifies a non-equilibrium state. The void does not transform into a thinning zone as in Mechanism A but

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.11 instead it grows quickly by crumbling until it reaches the overlying loam horizon that is able to support arching. The fourth stage (quasi-equilibrium state) includes further void stoping through the loam horizon. It occurs gradually by crumbling. As the void approaches the soil horizon, the destruction process is increasingly influenced by daily and seasonal changes of temperature and moisture content. The fifth stage (non-equilibrium state) occurs in most cases as a single catastrophic event, i.e. as a collapse of the remaining roof of a void, with eventual appearance of the surface feature. Depending on local conditions, it can occur either when some part of the loams still remains at the roof or when arching is supported solely by the soil horizon. The latter case is common (with a roof thickness of about 0.3 to 1.0m), as rhizomes reinforce the soil in unploughed areas. Failure can be induced by extreme wet or dry periods, seismic events (blasting in a nearby quarry), application of additional load and ploughing. It is quite common that formation of concentric cracks and shallow subsidence precedes collapsing. Final collapse events are commonly accompanied by noise and dust ejection. This indicates that the roof collapses into a void that is already separated from the main cave. The newly formed collapses have a diameter of 3 to 5m, depth of 2 to 5m and a bottle-like cross section (the diameter at the base is 10 to 40% larger than at surface level). The full development sequence is described above. However, some variations are possible toward the reduction of the number of stages due to: 1) the presence of structural or lithological discontinuities and irregularities in the overburden and, 2) incomplete thickness and composition of the overburden, such as in the lower (IId) terrace, where the Kosovsky Formation is only a few to 10m thick and the sandy-gravel and loam beds are entirely removed. Also, the last stage, that is the appearance of the subsidence or collapse at the surface, may never occur where the thickness of the overburden is large enough to cause self-liquidation of the stoping void (this point is discussed further below). Distribution of breakdown structures. The resultant map shows most of the breakdown structures existing in the cave field, regardless of whether or not they are expressed at the surface (Fig. 8A). The mapped breakdown structures were classified according to their stages of development, as described above. The survey data suggest that the overall density of breakdown structures for the whole cave field is more than 1800 per km 2 However, this parameter varies substantially between cave regions, from 254 BS/km 2 in the Central'ny region to 5271 BS/km 2 in the Kamchatka region. As the regions differ in size and morphology of passages and in the characteristics of their patterns, it is important to examine possible relationships between the number of breakdown structures and parameters of passages and their patterns in particular regions. Respective correlation coefficients are given in the last row of Table 2. As can clearly be seen all the variables characterizing passage size show no appreciable correlation with the BS number. This agrees well with observations in the caves. Whereas some of the largest cave passages (up to 20m wide and 10m high; see photos on Fig. 6 for instance), being closely spaced and separated by only small pillars, host no or few breakdown structures, other much smaller passages contain many breakdowns. It is further illustrated by some details of the dataset under examination. The Metropoliten region, which consists of large passages (specific volume 16.1 m 3 /m) has one of the lowest breakdown structure densities (306 per km 2 ), whereas the Kamchatka region, with a specific volume of 5.8m 3 /m has the highest breakdown density. The above finding is in striking contrast with established views, which suggest that breakdown formation is controlled primarily by passage size. The lack of correlation between BS number and passages size agrees well with the observation, mentioned above, that the vast majority of breakdowns in the cave initiate and develop from solution domes and cupolas that expose the bottom of the overlying Ratynsky limestone bed to the cave. Even a rather small-sized outlet cupola that exposes a few m 2 of the Ratynsky bed may give rise to the formation of breakdown structures. In contrast, large spans of master passages tend to remain stable if no outlet features occur. The causal relationship of breakdown formation and the outlet features is discussed further below, in more general speleogenetic terms. Site-specific collapse/subsidence hazard assessment. As the BS development stage signifies a certain level reached by a stoping void in a given geological cross-section, one can readily deduce the depth of a void position below the surface by superimposing the isopachyte map on the breakdown map (Figs 8, A and B). These data allow the main questions of engineering karstology, about where and at what depth voids stoping through the overburden located, to be answered with great precision. Adopting a hazard categorization based on an understanding of the breakdown mechanisms, one can produce a map of the micro-zoning of the territory according to the degree of subsidence/collapse hazard presented at the surface (Fig. 8C). On this map some arbitrary categorization of the hazard is used that reflects the depth of the stoping void below the surface and a relative probability of the collapse/subsidence deformation at the surface. Areas of low, moderate, high and very high hazard are distinguished. Depending upon the overburden thickness, the same breakdown stages can cause different degree of hazard: the lower thickness, the earlier stage can result in the surface appearance of the collapse. The blank areas within the cave field are non-hazardous, although the blank areas outside the cave field limits are non-classified because evaluation is based on the direct mapping in the cave. Hence, the areas outside cave field cannot be assessed in the same way. Another question important to hazard assessment is that of the possible size of collapse/subsidence when it appears to the surface. The answer can be inferred from the above description of the breakdown mechanisms. The main gauging factor is the diameter of the outlet domes/cupolas, initiating breakdown at the level of the cave. Most commonly it varies between 1 and 5m. The

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.12 processes involved during the first three stages result in upward stoping without appreciable increase of the void at the top of the breakdown column. Mechanism A (the subsidence mechanism) implies a possible increase of the subsidence-prone zone of 2 to 3 times, because of lateral extension of the thinning zone in the aquiferous sandy horizon. This determines the expected size of subsidence at the surface to be within a few to 15m. Development according to Mechanism B (the collapse mechanism) can cause an increase in the expected diameter of the surface collapse in only 30 to 50% of the initial breakdown column diameter. Hence, the possible collapse size is 2 to 8m, which agrees well with the actual sizes of newly formed collapses. 5.5. The Dan'kivsky collapse The Dankivtsy area is located 12km north of Zoloushka Cave, still within the subjacent karst zone. The Collapse formed suddenly, on January 11 1998, on a gentle slope of a small stream valley. According to local people the noise of the collapse was heard at a farm lying about 1km from the site. A 22m-deep shaft has formed, with an open entrance to a cave at the bottom (Fig. 9A). The shaft walls exposed the clays of the Kosovsky Formation, which graded into loams in the upper part. The walls displayed fragments of slickensided rock, which suggests that the collapse occurred along a fault. This is supported by the presence of leakage patterns at the soil/loams contact in the upper part of the walls, which indicate prolonged vertical percolation along the fault at the site of the subsequent collapse. The bottom of the shaft was almost entirely occupied by breakdown material. Its arrangement (the presence of large blocks of clay with fragments of slickensided material) suggests a single-event collapse. From the top of a breakdown pile it was possible to climb down into a chamber, a widened and domed part of the passage where 1.5m airspace occurred in the otherwise totally inundated cave (see plan and profile on Fig. 9, B and C). A water-filled passage, about 8to 9m wide and 7m high continued in a NE direction. The cave was surveyed in April, and in May it became inaccessible due to the continuing filling of the entrance shaft by loose sediments. By October, the shaft was already transformed into a bowl-shaped sinkhole 4m deep. It is evident that this surface feature will soon assume a gentler shape, quite similar in morphology to numerous subsidence features identified in the vicinity. The sinkhole mouth lies at an altitude of 173m, some 8.3m above the bottom of the small valleyand 67.3m above the floor of the Dniester Valley some 5km to the north. The bed of the surface stream in the local valley is 19.2m above the groundwater table exposed in the cave, so that the stream is perched on the clays above the vadose zone. In April, the water table was at 2.7m below the top of the gypsum. Fig. 8. The fragment of the map of breakdown structures in Zoloushka cave (A), the profile showing different stages of their development the heights of their propagation to the surface (B) and the map of micro-zoning of the territory according to subsidence hazard. A key to the Fig. 8A: 1 = cave passages; 2 = passages destroyed by the quarry; 3 = isopachytes;, 4 7 = breakdown structures with the breakout cavities positioned at various levels: 4 at the bottom of the Ratynsky bed, 5 within the Kosovsky Clays, 6 within the sandy-gravel bed, 7 within the loam bed; 8 = surface karst features; 9 = the quarry faces.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.13 Fig. 9. The Dankivsky Collapse, an example of the sudden single-event collapse: A the collapse shaft at the surface in three month after the formation, B plan, C profiles. The form of the documented part of the cave suggests that it is a fragment of an extensive maze cave system analogous to Zoloushka Cave. Morphology and passage size are quite similar. This analogy is also supported by geophysical survey results indicating a labyrinthine pattern in the vicinity of the collapse. The Dan'kivsky collapse exemplifies a rare case of a sudden single-event collapse. It occurred along a prominent fault (probably two closely related faults) at a locality where the presence of the underlying enlarged cave passage with a cupola (an outlet?) had reduced stability to a critical level. Percolation along the fault had reduced friction within the clay along the fault plane, and this further conditioned the collapse to occur through the entire thickness of the clays. Another important lesson from the Dan'kivsky case is that the shape of a surface form is not necessarily indicative of its collapse (sudden) or subsidence (gradual) origin. With the presence of soft sediments within the overburden, an original collapse shaft can be transformed into a gentle-sloped doline within few years. 6. Breakdown development in the entrenched karst zone: Mlynki, Slavka and Verteba caves The zone of entrenched gypsum karst in the Western Ukraine lies mainly to the north of the Dniester valley (Podol'sky region). The deeply incised river valleys of Dniester and its left sub-parallel tributaries separate the Miocene sequence into a number of isolated deeply drained interfluves where the gypsum and clay overburden remain largely intact. The Miocene sequence is almost entirely drained and only in the central parts of the inter-valley plateaus do the sub-gypsum units contain unconfined underground water, locally occupying also the lowermost part of the gypsum. Maze cave systems in the gypsum are presently relict. Modern dissolution is restricted to the lower part of gypsum, where the water table is present, at rare points of focused vertical percolation (where vertical dissolution pipes develop) and along short linear underground streams that are fed via swallow holes that receive periodic surface flow. Sinkholes are generally few within the high interfluves, but their density increases locally where the capping clays are removed, as within high river terraces or the floors of perched valleys. When compared to the settings of the Zoloushka area there are some distinctly different features in the lithoand hydrostratigraphy of the overburden, and these are important to breakdown development: The 2 to 5m-thick Upper Badenian unit, which immediately overlies the Ratynsky bed, is composed of marly lithothamnion carbonates (the Ternopol'sky beds). This material is capable of crumbling gradually, to support breakout cupola development. The formation lying next above is represented by massive, rather homogenous, fine marine clays (the Lower Sarmatian) up to 60m in thickness depending on local relief. This material is quite coherent when dry, and if it is thick enough it can prevent further upward migration of a void. However, where wet (along tectonic or stratigraphical discontinuities that support groundwater percolation across the otherwise almost impervious thickness), it demonstrates a kind of viscous-flow behavior, and can be extruded into the cave through breakdown structures, like toothpaste from a tube. Also, the Sarmatian clays can shift down as blocks, by sliding between two closely spaced faults if cave and breakdown development result in a decrease of support from below. Alluvial pebble/gravel sediments of the ancient upper terraces of Dniester occur overlying the clays. In most of the region these are effectively drained due to the high degree of erosional dissection. Hence, Mechanism A of the breakdown development described for the Zoloushka Cave area does not operate in the Podol'sky region. These peculiarities lead to some distinctive variations in the breakdown development in the entrenched karst zone, as compared to the development in the subjacent and deep-seated karst zones described above.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.14 6.1. Mlynki Cave The cave lies at the northern edge of the entrenched karst zone. The entrance opens into a valley slope within the gypsum outcrop. The thickness of the overburden increases to 25 to 30m towards the plateau. Only two sinkholes are recorded at the surface within the explored cave field. The cave is a maze, currently surveyed for 26km, in which passages occur on two levels. In the upper level the passages are mainly slot-like in shape, 1 to 2m wide and up to 5m high. In the lower level passage cross-sections are commonly wider, and many have a rift extension down to the base of the gypsum. Fig. 10. Breakdown structures in Mlynki Cave. The cave map is a courtesy of the Chortkiv Speleological club. Breakdown survey has been performed with an assistance of Vladimir Snigur. Complete mapping of breakdown structures has been performed for five cave regions, enabling estimation of density values. In total, 144 breakdown structures have been mapped (Fig. 10). In 57 cases breakout cupolas at the top of BS are positioned within the Ternopol'sky carbonate bed and are accessible from the main cave. In 87 cases migrating voids are separated from the cave by breakdown talus. Only in few breakdown structures is the Sarmatian clay identified in the talus, indicating unambiguously that the stoping void had entered the clay thickness above the Ternopol'sky bed. Almost all breakdown structures in Mlynki Cave developed from outlet cupolas (see photos on Fig. 4, A and B). The extrapolated density of breakdown structures varies from 700 to almost 3000 per km 2 between cave regions, with the average value for the whole set being 1609. These characteristics are quite similar to Zoloushka Cave despite the many differences in passage size, morphology and geohydrological setting. This can be explained by the similarity of the initiation conditions, by the fact that in both cases breakdown structures initiate from outlet cupolas/domepits. Hence, these morphogenetic features impose the most important control of breakdown initiation and distribution. However, in contrast to the Zoloushka area, most of the breakdown structures do not reach the surface and they remain stable and hidden in the subsurface after reaching the base of the Sarmatian clay. This reflects the fact that both the cave and the overburden are fully drained and do not demonstrate any considerable hydrogeological activity.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.15 6.2. Slavka Cave The cave lies within a spur of the interfluve plateau, bordered by a stream valley and its two small tributaries perched karst valleys. The entrance is a collapse sinkhole on the slope of one such valley. Overburden thickness increases to 20 to 25m toward the interfluve. The cave is currently explored for 9.2km and consists mainly of high (3 to 10m) slotand rift-like passages. Feeder conduits, commonly separate, form a lower level, relative to the master passages; in places they form small networks. Breakdown structures in Slavka Cave are of two types: 1) "common" breakdowns formed from outlet features, and 2) breakdowns related to "rhythmolitic" bodies (see below). The latter represent a special case of breakdown formation, occurring widely in Slavka Cave but rare in other caves. Rhythmolites is the local term for highly gypsiferous bodies of closely interbedded aleurolits, sands and coaly streaks occurring within the upper part of the gypsum. Such bodies can be of 5 to 10m across and 3 to 4m in vertical thickness. Although "rhythmolite" bodies are found in many other caves in the region, they are unusually abundant throughout the Slavka cave field. Their contact with the gypsum is irregular and generally has a coneor bowl-like shape. The nature of the "rhythmolite" bodies is not well understood, but one can assume that they are paleokarstic (syngenetic?) features. Fig. 11. Breakdown structures in Slavka Cave. The cave map is a courtesy of the Kiev Speleological club. Breakdown survey has been performed with an assistance of Natalia Yablokova.Being particularly brittle, closely stratified and fissured, slabs of rhythmolitic material fall readily into any passage whose ceiling has intersected the bodies from below, giving rise to breakout cupolas/domes (see Fig. 4C). Breakdown piles beneath these consist almost entirely of rhythmolite slabs. In total 42 breakdown structures of this type were mapped within the cave (Fig. 11), giving an extrapolated density of about 480 features per km 2 Most breakout forms are confined to the rhythmolite bodies or terminate at their contact with the overlying Ratynsky limestone or Ternopol'sky unit. Examination of such cupolas and domes suggests that breakdown structures of this type are not related to any discontinuities in the overlying formations; the latter remain largely intact above such breakdown. The scarcity of surface subsidence and collapse features above the Slavka Cave suggests that this type of breakdown structure is generally not sufficiently potent to produce a surface expression, despite the overburden being of rather small thickness. Considering that rhythmolite bodies are not common in other cave areas examined, this type of geological influence on breakdown initiation can be regarded as site-specific. "Common" breakdown structures, i.e. those formed from outlet features, are quite similar to those in Mlynki Cave. Only 13 structures of this type were found, which suggests an extrapolated density of about 150 features per km 2 Both density values given above are somewhat underestimated, because some structures have probably been overlooked in the marginal parts of the labyrinth. However, it is evident that in this particular case breakdown structures related to "rhythmolites" predominate over structures related to outlet features. 6.3. Verteba Cave Verteba Cave lies in the neck of a large meander of the Seret River, where the cover sediments are almost entirely denuded. Only isolated patches of the Ratynsky limestone and Ternopol'sky beds, from 1 to 3m in thickness, remained within the meander, but through most of the area the gypsum is covered only by soils. Many sinkholes of cone-, bowland plate-like shapes are scattered throughout the area (Fig. 12). The cave is a shallow-lying labyrinth with 7820m of closely-spaced, wide but low (primarily due to the high level of clay cave filling) passages that occur within a narrow strip. Numerous breakdown structures examined in this cave fall into three groups: 1) "Common" breakdowns formed from outlet cupolas; 2) Breakdowns formed from vertical dissolution pipes; 3) Breakdowns formed along prominent vertical cracks in the cave ceiling. Massive fall-ins of blocks are rare. Even in this shallow cave they do not cause total breakdown of the cave ceiling with sudden collapsing at the surface. Gravitational breakdown of any remaining "bridges" at the top of the cupolas and the vertical pipes is also infrequent, because of low lithostatic loads. The vast majority of breakdown structures in Verteba Cave is associated with prominent vertical cracks in the cave ceiling and involve mainly filtrational mechanisms, such as suffosion and erosion. Through such cracks, poorly consolidated fragments of the Ternopol'sky beds and remnants of the soil cover are washed into the cave, giving rise to numerous suffosion sinkholes at the surface and related breakdown piles in the caves (see Fig. 4G).

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.16 The breakdown piles are rather small in size and consist of mainly (sometimes solely) of washed-in soil. Artifacts have been found in some piles, originally dumped into the sinkholes by inhabitants of the Tripil'sky settlement (about 5000 years BC) located above the cave. This gives some idea of the rate of filtrational mass flux from the surface to the cave. Fig. 12. Distribution of sinkholes above Verteba Cave. The cave map is a courtesy of the Ternopil' Speleological club 7. Discussion and conclusions Speleological observations have allowed the identification of several different mechanisms for breakdown development in the gypsum karst of the Western Ukraine. Distribution of breakdown structures and the mechanisms of their development are influenced mainly by: 1) Speleogenetic factors (distribution, type and size of the breakdown-initiating cavities or their particular components); 2) Lithological and structural discontinuities in the gypsum encountered by caves (combined speleogenetic and geological guidance); 3) Lithostratigraphy of the overburden and the geotechnical properties of its individual units; 4) Lithological and structural discontinuities in the overburden; 5) Hydrogeological conditions at the level of the gypsum and in the overburden. 7.1. Breakdown initiation Ultimately karst breakdown development is related to the presence of karstic cavities and dissolutionally enlarged fissures. However, in contrast to established views, this study suggests that breakdown initiation in the Western Ukraine is not guided directly by the size of cavities. Some of the largest passages and chambers in the major caves remain stable and untouched by gravitational destruction. In many other cases breakdown of large gypsum blocks occur from the ceiling of passages, but the respective breakout surfaces remain stable, still within the gypsum (corresponding to particular prominent bedding planes). Even when the breakout surface occurs along the base of the Ratynsky bed, it remains stable in many cases. Only in rare situations can the massive breakdowns terminating some large passages be assumed to form as single-event collapses of the cave roof. Apparently, such cases are guided mainly by geological factors. This survey suggests strongly that the great majority of breakdowns initiate at specific speleogenetically or geologically "weakened" localities (factors 1 and 2 in the list above) that classify into few distinct types. Speleogenetic controls. In the Western Ukrainian caves, two types of speleogenetic situations that favour breakdown initiation are distinguished, both creating exposures of the Ratynsky bed in the caves: 1) outlet features (cupolas, domes and domepits of "ascending" origin) and, 2) vertical pipes formed by free downward percolation. In all the caves examined most of the breakdown structures initiate from outlet features. Such features represent places where water has discharged from a cave to the upper aquifer during a period of transverse artesian speleogenesis. By virtue of their hydrological function, the outlet features tend to form at places where the integrity of the immediately overlying Ratynsky bed and of the next higher formation are somewhat disrupted and, hence, permeability is enhanced. In other words, all of the most weakened zones at the gypsum/Ratynsky limestone contact were exploited speleogenetically, to form outlet features during transverse artesian speleogenesis. This is the single fundamental cause of breakdowns initiating predominantly from outlet features. Therefore, distribution of outlet features is the most important influence upon breakdown initiation. By way of contrast, the above reasoning is supported by the fact that the Ratynsky bed commonly forms relatively stable ceiling in large exposures created by occasional separation and breakdown of gypsum blocks into the underlying master passages, if this does not relate to outlet features. The second speleogenetic situation favouring breakdown initiation is where vertical dissolution pipes form under present unconfined settings in the entrenched and subjacent karst zones. Such pipes develop at points of a focused descending percolation to the gypsum from the overlying beds. They are 1 to 3m wide, extend downwards from the gypsum upper contact through the full thickness, and are commonly superimposed upon relict artesian passages (Fig. 13). Although in a different way, the vertical percolation pipes also expose the base of the Ratynsky limestone to the caves. Also, as in the case of the outlet features, the vertical percolation pipes commonly indicate weakened zones in the Ratynsky limestone bed and in the overlying formations. This is why breakdown structures are readily initiated from such pipes. The mechanisms for breakdown development remain largely the same as for the structures formed from outlet features, although some differences can be imposed

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.17 by continuous active percolation, hence the erosion of breakdown structures. Active percolation facilitates upward stoping through larger overburden thicknesses. The density of vertical pipes in the main caves of the region (a few to a few tens of pipes per km 2 ) is much lower than that of outlet features. However, in other regions their density is known to be much higher, for instance about 300 pipes per km 2 in the Kungursky Cave area of the Urals, Russia (Andrejchuk, 1999). There, breakdown development from vertical pipes is considered to be the main cause of collapse/subsidence development at the surface. Fig. 13. The development of breakdown structures after vertical dissolution pipes. Geological factors. Various kinds of geological discontinuities occurring within the gypsum can initiate breakdowns in the ceilings of "normal" passages. One type is exemplified by Slavka Cave, where most of the breakdowns relate to (palaeokarstic?) bodies of gypsum-rich "rhythmolites". Generally breakdown structures of this type are not sufficiently potent to produce an expression at the surface, even if the overburden has a relatively small thickness. Considering that "rhythmolite" bodies are not common in other cave areas that have been examined, this type of geological influence on breakdown initiation can be regarded as site-specific. Another type of geological influence is where breakdown is initiated at points where prominent tectonic faults disrupt both the gypsum and the overburden. If faults are pre-speleogenetic, they can cause development of larger passages that may collapse suddenly when their ceiling strength is exceeded by the load of the overburden. The presence of prominent sub-vertical discontinuity planes in the overlying clays facilitates massive breakdown. It is the presence of such guiding discontinuities that allow overburden material to collapse in a single event, even where there is a considerable thickness of overburden. It is presumed that all the deep collapses known in the region are of this type. They appear at the surface suddenly, as catastrophic collapses, forming 10 to 30m shafts, as exemplified by the Dan'kivsky collapse. Such collapses are quite rare both throughout this region and more generally, but they are the most hazardous, due to their considerable vertical magnitude (energy involved) and the difficulties inherent in their prediction. 7.2. Breakdown propagation through the overburden The propagation of a void through the overburden by stoping and the possibility that a breakdown structure will eventually manifest itself at the surface as a subsidence or a collapse depend on the thickness and lithostratigraphy of the overburden and on the particular mechanism involved. Lithostratigraphy of the overburden. The presence, layered structure and lithological composition of an overburden are among the major factors that determine stages of karst breakdown development and the component processes involved, i.e. the mechanisms of karst breakdown. Multi-stage development is governed by the stratified nature of the overburden, with varied lithological, geomechanical and hydrogeological properties for individual units. Generalizations, derived from the major publications on the problem and supported by this study, are as follows. Beds of loose, permeable sediments (i.e. sands) serve as predominantly as a setting for processes of hydrodynamic decomposition (such as suffosion, liquefaction, erosion, etc.). In contrast, low-permeability or fully-drained beds of more coherent sediment or solid rock promote arching to support void development by stoping, and serve as the setting for mainly gravitational destruction. Consequently, in the overall propagation process some non-equilibrium stages give way to quasi-equilibrium stages. The capability of some beds within the overburden to bridge a void is the main pre-requisite for collapse-style in the eventual surface deformation (as against gradual subsidence). Hydrogeological conditions The role of hydrogeological conditions in creating solution cavities that initiate breakdowns (speleogenetic factors) is not discussed here. However, these conditions play an important role in determining breakdown initiation (triggering) and development. In the Western Ukrainian gypsum karst, one of the most important effects that triggered breakdown development was the loss of buoyant support when the Miocene aquifer had been losing its confinement due to geomorphic development. The loss of buoyant support can disturb the metastable state of a cave roof at points where speleogenetic and geological factors have already brought its bridging capacity (resistance to failure) close to a critical level. This situation generally signifies the transition from deep-seated karst to subjacent karst. The effect is illustrated by many quarries in the deep-seated (confined) karst zone (Jazovsky sulfur quarry, Nikolaevsky clay quarry, etc.), where overburden removal and massive groundwater withdrawal from the Miocene aquifer resulted in an abrupt drop of potentiometric surfaces and dramatic intensification of collapse/subsidence formation

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.18 in the vicinity of the quarries (Klimchouk and Andrejchuk, 1996). In unconfined settings, hydrodynamic activity at the gypsum level promotes destabilization of breakdown columns that rest on the clay cave fill. Their basements can settle down, due to shrinkage and compaction of bulk cave fill, and be eroded by focused streams.In all cases this causes settlement of the breakdown columns. This is why lowering or fluctuation of the water table in the gypsum commonly activates breakdown development and subsidence formation at the surface, a case that is well exemplified by the Zoloushka Cave area. A complete draining of the gypsum promotes stabilization of the breakdown structures and slows down their propagation to the surface (the Mlynki Cave). Another important consideration is the presence of perched aquifers within the overburden. In the region, an aquifer hosted in the sandy-gravel alluvial terrace sediments perched on the Kosovsky/Sarmatian clays is present in many places. In the confined karst zone it contains groundwaters almost universally throughout the area, but in the subjacent and, especially, in the entrenched karst zones it is drained in part or in full by erosion valleys and subsurface breakdown structures. Where a stoping void at the top of a breakdown structure reaches the bottom of this water-bearing horizon, the set of hydrodynamic destruction processes evolves, such as suffosion, liquefaction, erosion, etc. Besides suffosion, liquefaction and thinning occurring at the sandy horizon itself, leakage along the breakdown structure, if continuous and intense enough, may cause considerable destabilization and settling of the breakdown column due to erosion and damping, hence promoting activation of the overall process. 7.3. Breakdown mechanisms The factors considered above (initiation conditions, lithostratigraphy and hydrogeological conditions) together determine the mechanisms of breakdown propagation and surface deformation. Five mechanisms identified by this study are summarized on Fig. 14. In two of them (2nd and 5th) the processes of gravitational destruction overwhelmingly predominate, but in the other three the processes of filtrational (or hydrodynamic) destruction play an important part, either in particular stages or during the entire breakdown development. Consequently these mechanisms are termed "gravitational" and "gravitational/filtrational". Note that the initiating situation only directly determines Mechanism 5, whereas other mechanisms strictly do not depend on the ways in which breakdown started. The specifics of the mechanisms are determined mainly by lithostratigraphical and hydrogeological conditions. Fig. 14. Mechanisms of the breakdown development in the gypsum karst of the Western Ukraine.

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.19 7.4. The critical thickness of the overburden The ability of a breakdown structure to reach the surface in the form of a collapse or subsidence depends on: 1) the size of the initial "breakdown window" at the gypsum/Ratynsky bed contact and the receptacle capacity of a master cavity beneath it, 2) the coefficient of loosening of the fallen material, 3) the thickness of the overburden and, 4) the involvement of the processes of hydrodynamic destruction. As most breakdowns in the region initiate where artesian outlet features or vertical percolation pipes expose the Ratynsky bed to a cave, the diameter of such exposures is the main gauging factor that determines the cross-sectional size of the breakdown column. The receptacle capacity of a master passage beneath the initiating feature is determined by passage width and height. Hence, these parameters influence the initial size of a stoping void at the top of the breakdown structure. When a breakdown talus reaches the gauging "window" and separates the migrating void from the main cave, the height of the void decreases in the course of its further upward migration, because of the loosening of the fallen material. This means that at a certain height of breakdown column void propagation may cease and the breakdown will never manifest at the surface. Hence, given some maximum parameters for the initial receptacle capacity of a cave, and a characteristic coefficient of loosening for the overburden material (it commonly varies from 1.1 to 1.3 for the region), one may speak about the critical thickness of the overburden, above which surface deformation will never occur, regardless of the degree of underground karstification. Empirical study of the relationship between sinkhole density and the thickness of the overburden performed for the three different areas (Fig. 15), seemingly supports this assumption. The critical thickness of the coverbeds is found to vary from 45 to 55m between the three different areas. The shape of the curve for area 3 (Zoloushka Cave) differs from the other two because of the intensification of collapse and subsidence formation caused by man's impact (quarrying activity). Apparently, the critical thickness of the overburden will be specific for each region. It depends on the size of the cavities and the structure and composition of the cover. This scheme is strictly valid, however, only for purely gravitational mechanisms that do not involve the processes of hydrodynamic destruction and removal, as the latter can maintain the non-karstic growth of a stoping void or restoration of the receptacle capacity of an initial karstic cavity. And, finally, it does not apply for situations where the cover is made of sediments that have a coefficient of loosening close to 1.0, such as sands. In sands, a stoping void can propagate through great thicknesses of up to 100m or more. 7.5. Final considerations Breakdown initiation at the karstified horizon (at the cave level) occurs due to various causes. Simple gravitational breakdown of blocks and slabs from the cave ceiling rarely gives rise to destruction of the overburden. The most important conclusion of general significance derived from this study is that breakdown of the overburden is caused predominantly by structures related to specific morphogenetic features in cave systems (outlet features), or to specific genetic types of conduits (vertical solution pipes), not merely to large unsupported roof spans. This is because, by virtue of their origin and hydrogeological function, such features exploit the points of lowest integrity within the main bridging unit (the Ratynsky bed in the Western Ukraine) and the entire overburden. This is also the reason why such breakdown structures are sufficiently potent to propagate through the overburden, whereas those related to occasional block breakdown of the cave ceilings are commonly not. Therefore, passage size is not an important influence upon breakdown initiation. Breakdown formation can proceed through a variety of mechanisms. In the intrastratal and covered karsts, manifestation of karstic features at the surface does not adequately reflect the degree and character of karstification at depth. The shape and size of sinkholes is not indicative of their origin. Structure and composition of the cover and the processes therein play the major role in transmission of breakdowns to the surface. This study demonstrates that speleogenetic analysis plays one of the most important roles in understanding breakdown pre-conditions and mechanisms, and in eventual subsidence hazard assessment. Direct cave observations aimed at both speleogenetic investigation and breakdown characterization at regional or site-specific levels should be employed wherever possible. Acknowledgement This study was partially supported by the ROSES (Risk of Subsidence due to Evaporite Solution) Project ENV4-CT97-0603 funded by the EC Framework IV Programme. The authors sincerely thank Dr. David Lowe and Dr. Armstrong Osborne for the correction of English. Fig. 15. The density of sinkholes versus thickness of the overburden: 1 = in the Seret-Nichlava interfluve (entrenched karst), 2 = in the Cherny Potok area (subjacent karst), 3 = in the Zoloushka cave area (subjacent karst).

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A.B.Klimchouk and V.N.Andrejchuk / Speleogenesis and Evolution of Karst Aquifers 1(1) January 2003, p.20 References Andrejchuk V.N. 1984. The regularities of karst development in the south-east of the zone of junction between the Russian platform and Carpathian foredeep. PhD Thesis, Chernovitsky University (russ.). Andrejchuk V.N. 1988. The tectonic factor and peculiarities of the sulfate karst of Bukovina: geology, geomorphology and hydrogeology of karst. Sverdlovsk, 66 p (russ.). Andrejchuk V.N. 1999. Collapses above gypsum labyrinthic caves and stability assessment of karstified terraines. Prut: Chernovtsy. 51 p (russ.). Andrejchuk V.N. and Korzhik V.P. 1984. The Zoloushka karst system. Peshchery (Caves). Types and Metods of Investigation. Perm, Perm University, 12-25 (russ). Klimchouk A.B. 1990. Artesian origin of large labyrinth caves in the Miocene gypsum of the Western Ukraine. Doklady AN Ukr.SSR, Ser B Geol chem and biol sciences 7, 28-32 (russ). Klimchouk A.B. 1992. Large gypsum caves in the Western Ukraine and their genesis. Cave Science 19 (1), 3-11. Klimchouk A.B. 1996. Gypsum karst in the Western Ukraine. In: Klimchouk A.B, Lowe D.J, Cooper A.H and Sauro U. (Eds.), Gypsum karst of the World. Int. Journal of Speleology Theme issue 25 (3-4), 263-278. Klimchouk A.B. 2000. Speleogenesis of the great gypsum mazes in the Western Ukraine. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of karst aquifers. Huntsville: Natl. Speleol. Soc., 261-273. Klimchouk A.B. and Andrejchuk V.N. 1988. Geologic and hydrogeologic conditions of the development of large gypsum caves in the Western Ukraine and their genesis. Peshchery (Caves). Gypsum and Anhydrite Caves. Perm University, Perm, 12-25 (russ). Klimchouk A.B. and Andrejchuk V.N. 1996. Breakdown development in cover beds, and landscape features induced by intrastratal gypsum karst. In: Klimchouk A.B, Lowe D.J, Cooper A.H and Sauro U. (Eds.), Gypsum karst of the World. Int. Journal of Speleology Theme issue 25 (3-4), 127-144. Klimchouk A.B. et al.. 1985. The study of geological and hydrogeological conditions of karst development of the Pridnestrovsky Podolia in the connection with the establishment of a karstological monitoring site. In: Sokolovskij I.L. and Klimchouk A.B. (Eds.), Karst of the Ukraine. Fizicheskaja Geografija i Geomorfologija, vol.32. Vyshcha Shkhola, Kiev, 47-54. (russ.). Shestopalov, V.M. (Ed.). 1989. Water exchange in hydrogeological structures of the Ukraine. Water exchange under natural conditions. Kiev, Naukova Dumka, 288 pp. (russ.). White E. and White W. 2000. Breakdown morphology. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of karst aquifers. Huntsville, Natl. Speleol. Soc., 427-429.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal www.speleogenesis.info Dynamics of cave development by allogenic water Arthur N. Palmer Department of Earth Sciences, State University of New York, Oneonta, NY 13820, USA Re-published by permission from: Acta Carsologica vol. 30, n.2, 2001, 14-32. Abstract Streams that drain from non-karstic surfaces tend to have great di scharge fluctuations and low c oncentrations of dissolved soli ds. Where these streams encounter karstic rocks they can form caves with hydraulic and ch emical dynamics quite different from those fed by autogenic recharge (e.g. through dolines). In either cas e, caves form only along those paths in which the discharge can increase with time. Only a few favorable pa ths achieve this goal, while the others st agnate with small and diminishing enlargem ent rates. Caves in carbonate rocks that are fe d by allogenic streams have a relatively short inception period, after which the mea n-annual rate of dissolutional wall retreat is typically about 0.01 cm/yr. Most of the annual growth takes place during a few major floo ds that occupy only a small fraction of the year. Local growth rates can be enhanced by abrasion from sediment. During floods, highly aggressive water is de livered rapidly to points deep within th e karst aquifer. As flood discharge increas es, cave streams become ponded by constrictions caused by detrital sediment, insoluble beds or collapse material. The head loss ac ross a constriction varies with the fifth power of the diameter ratio under pipe-full conditi ons. Head loss also increases with the sq uare of the discharge. Because the discharge duri ng a flood rises by several orders of magn itude, the head loss across constrictions ca n increase enormously, causing water to fill parts of the cave under c onsiderable pressure. This highl y aggressive water is injec ted into all available openings in the surrounding bedrock, enlarging them at a rapid and nearly uniform rate. Depending on the structur al nature of the bedrock, a dense array of b lind fissures, pockets, anastomoses, or sponge work is formed. Many such caves develop traversable mazes that serve either as by pass routes around constrictions, or as "karst annexes", which store and later release floodwaters. Many features that are sometim es attributed to slow phreatic flow or mixing corrosion are actually generated by po nded floodwaters. In caves that experience severe flooding, adjacent fi ssures or bypass routes with initial widths at least 0.01 cm can grow to traversable size within 10,000 years. Keywords: speleogenesis, allogenic recharge to karst Introduction Allogenic drainage into ponors from non-karstic rocks tends to be highly aggressive toward carbonate and also varies greatly in discharge. Caves formed by this water have a much more dynamic developmental history that those formed by autogenic recharge from a karst surface. They represent the underground aspect of “border corrosion” described by Gams (1965). Not only does allogenic recharge tend to enlarge caves more rapidly than autogenic water, but it also can determine their entire passage pattern. Solution pockets, anastomoses, blind fissures, and crude mazes are commonly superimposed on the primary cave passages. Some caves formed in this way consist entirely of network or anastomotic mazes. Severe flooding is characteristic of most of these caves. Cave enlargement by floodwaters can therefore be considered epiphreatic. However, floodwater inundations are quite different from the slow water-table fluctuations typical of other hydrologic settings. Floodwater is the most dynamic member in the broad spectrum of epiphreatic conditions. Dissolution rates and conduit competition Rates of limestone dissolution have been measured by various researchers (e.g. Rauch and White, 1979; Plummer and Wigley, 1976; Plummer et al., 1978; Sjberg, 1976; Sjberg and Rickard, 1984; Buhmann and Dreybrodt, 1985a, 1985b). At normal groundwater temperatures and CO2 partial pressures, dissolution rates remain high over a broad range of calcite concentrations but drops rapidly to very low values beyond about 70% saturation (see Fig. 1). Using rates recalculated from the experimental data of Plummer et al.

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.2 Fig. 1. Rates of limestone dissolution vs. percent calcite saturation at 10o C and 0.01 atm CO2 partial pressure (calculated from references in text). (1978), Palmer (1981, 1991) used finite-difference analysis to show how the mean enlargement rate in conduits responds to passage size, discharge, flow distance, and water chemistry (Fig. 2). Discharge (Q) and flow length (L) have identical but opposite effects on rate of wall retreat and can be combined as a ratio (Q/L) on the graph. Enlargement rates are proportional to Q/L, and the only way for a specific passage to increase its enlargement rate is to gain discharge. However, there is a limit to the solutional growth rate, determined by kinetics, beyond which the discharge has little effect. This is because in carbonate rocks the rate is limited by the reaction rate at the solid surface, rather than by mass transport within the water (see Plummer and Wigley, 1976). The maximum rate is typically about 0.01-0.1 cm/yr, depending on the local water chemistry. These figures are compatible with field measurements in cave streams (e.g. High, 1970; Coward, 1975; Lauritzen, 1990) and with the results of recent numerical and analytical modeling (e.g. Dreybrodt, 1996; Gabrovšek, 2000). Fig. 2 shows a few representative points during the early development of a karst aquifer, when small amounts of water pass through narrow openings. The solutional enlargement rate differs greatly among the various flow routes (zone 1 in Fig. 2). This holds true whether the recharge is allogenic or autogenic. It is clear that the growth rate of any given conduit can increase only if the discharge increases. As dolines and ponors develop, the paths fed by them enlarge rapidly (e.g. paths A, B, and C), while other routes stagnate with low or even diminishing enlargement rates (e.g. D and E). Only those paths that gain discharge with time are able to grow into caves. The favored ones reach the maximum enlargement rate (zone 2 in Fig. 2), after which all passages with similar water chemistry grow at roughly the same rate. Enlargement rates vary with discharge (e.g. on a seasonal basis), because the water entering the aquifer tends to be more aggressive during high flow. This process is accelerated where the karst aquifer is fed by allogenic streams. Only a few major flow routes reach the maximum growth rate, and they enlarge rather rapidly because the water feeding them tends to be more aggressive. Also, a steep hydraulic gradient is maintained by the large flow of water from ponors. Maximum growth rate in mature conduits is not limited by dissolution, since mechanical erosion aided by coarse sediment load can approach or ev en exceed the effect of dissolution during floods (Newson, 1971; Smith and Newson, 1974). The flow distance (L) in Fig. 2 needs some explanation. This is the distance from where aggressive water first enters a particular conduit or other opening. In Fig. 2 an initial saturation ratio of zero is assumed (i.e. no dissolved carbonate). However, the saturation ratio is usually considerably higher, even in streams that drain relatively insoluble rock s. The shape of the diagram is still valid for water of any aggressiveness, but the maximum dissolution rate is smaller if the saturation ratio is greater than zero (see Palmer, 1991). Fig. 2. Mean rates of solutional wall retreat in limestone conduits (from Palmer, 1991). Q/L = ratio of discharge to flow distance. Early stages of karst aquifer development are represented by zone 1, which includes a great variety of dissolution rates. Q increases with time in only a few conduits, allowing them to grow at a faster rate (A, B, C), while others stagnate at low and usually diminishing enlargement rates. Only those that reach the maximum rate (zone 2) grow into traversable caves. Graphs are similar for conduits with non-circular cross sections, but the lines have gentler slopes.

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.3 Fig. 2 applies mainly to relatively pure limestones. In dolomites the disparity in enlargement rates is even greater, and although the maximum rate in zone 2 is still valid, the rates in zone 1 are lower. The figure does not apply to evaporites, because their dissolution rates are higher and are proportional to flow velocity. Floodwater dynamics In the early stages of conduit development by allogenic streams, subsurface flow is limited because floodwaters simply overflow onto the surface. The impact of allogenic streams on cave patterns is greatest when the conduits have enlarged sufficiently to carry all (or most) floodwater. By this time the caves are partly airfilled, because the low-flow discharge is unable to keep them filled with water. Initially most such caves consist of only a few passages, owing to the limited number of inputs. During floods, the caves carry highly aggressive water deep into the karst aquifer (Fig. 3). Floodwaters tend to pond behind constrictions in the stream passages. This is typical where there has been collapse, accumulation of sediment, or a narrowing caused by relatively insoluble rock. In turbulent flow, head loss is proportional to the fifth power of the passage diameter. Therefore, if one water-filled passage has only half the diameter of another, the smaller one will require a hydraulic gradient about 32 times greater to transmit the same discharge. The head loss also increases with the square of the discharge. During low flow there is generally very little ponding at constrictions. However, the discharge rises by several orders of magnitude during a typical flood, and so the head loss across constrictions increases enormously, causing water to back up and to fill parts of the cave under great pressure (Fig. 4). Rises in water level of more than 100 m are common in some caves fed by allogenic streams. This water has been ca rried in from the surface very rapidly and is highly aggressive. In areas of ponding it is injected into all available openings in the surrounding bedrock. Because of the steep gradients and short flow distances, these openings are enlarged at a rapid and nearly uniform rate at the top of the growth-rate graph (Fig. 5). Cave enlargement by allogenic floodwater is considerably faster than nearly all other speleogenetic processes. In caves that experience severe flooding by aggressive water, fractures with initial widths at least 0.01 cm can grow to traversable size within 10,000 years (Palmer, 1991). Fig. 3. Head loss across a 10-meter-long constriction in a water-filled conduit, at va rious discharges (Q) and diameter ratios (d1/d2). During floods, water can pond to great depths in the upstream parts of a cave because of such constrictions. Because the floodwater retains most of its aggressiveness as it passes through the cave, the factor L in Fig. 5 applies essentially to the distance into the surrounding fissures that receive the injected water. The maximum enlargement rate will be slightly smaller than that shown in Fig. 2, because the saturation ratio is not zero. Effects on cave morphology Cave enlargement by floodwater can be recognized by a variety of features, even in relict cave fragments where the original hydrologic context is unclear: !" Diversion routes around constrictions. These typically have irregular profiles and ungraded intersections (i.e., the passage floors and ceilings of intersecting passages are at different levels). They may have a maze pattern, but often consist only of a few alternate routes. !" Blind fissures and network mazes. These are typical of highly fractured bedrock, especially at shallow depth below the surface. Such passages may serve as “karst annexes” (Mangin, 1975), which receive floodwaters during rising flow and release it more slowly as the flood subsides.

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.4 Fig. 4. Distance of penetration of highly aggressive flow into water-filled conduits. In this example, incoming allogenic water is assumed to be already 50% saturated (concentration / saturation concentration = 0.5). The graphs show the distance required for the water to reach only 70% saturation, beyond which the dissolution rate begins to decrease rapidly. Variations with discharge (Q) and passage diameter (d) are shown. A typical cave stream can penetrate a great distance while retaining most of its initial aggressiveness. Fig. 5. Same as Fig. 2, but under floodwater conditions, when aggressive water is injected into many openings for short distances with steep hydraulic gradients (i.e. large Q/L ratios). All openings larger than a certain minimum size are enlarged at similar ra tes near the top of the diagram, producing maze patterns. !" Anastomoses and anastomotic cave patterns. Anastomoses are common where bedding-plane partings are the dominant openings, and where there have been frequent fluctuations in water level. Not all anastomoses are formed by floodwater, however: some are remnants of the initial flow paths that eventually developed into conduits (see Ewers, 1966). Anastomoses of this second type are common only in the upstream ends of flow paths, where the water is aggressive and the flow distance is short. In this way their origin is chemically identical (see Fig. 5). Anastomotic passages form either two-dimensional or threedimensional arrays of tubes, in which many or all can be simultaneously water-filled and enlarged during floods. For this reason these systems are sometimes called “underground deltas” (Maire, 1990). !" Solution pockets and spongework. Spongework forms where no significant fractures or partings are present. Once formed, ceiling pockets can be enlarged by turbulent eddies (Slabe, 1995) or by elevated CO2 partial pressure resulting from air compression by rising water (Lismonde, 2000). !" Large contrasts in the grain size and distribution of sediments. Floodwater conditions usually involve great variations in flow velocity, with the result that large cobbles and boulders occupy the main stream routes, and fine sediment fills regions of static flooding. In local areas, highvelocity floodwater can prevent sediment from depositing at all. Many of the dissolution features described here are often attributed to slow phreatic flow or mixing corrosion. Maze patterns, solution pockets, blind fissures, and anastomoses have all been used by various authors as indicators of phreatic cave development. In a sense they are correct, but not in the way they envisioned. These forms are best explained as the result of intermittent flooding above the low-flow water table. For example, in the dynamic hydrologic conditions of alpine karst, many tubular conduits that are often assumed to be phreatic have been show n by Choppy (1991) and

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.5 Audra (1994) to be actually the product of fluctuations in water level within the epiphreatic zone. Most floodwater enlargement of existing passages is upward and lateral if the floor is armored by sediment. Although periodically refreshed interstitial water within sediment can dissolve the underlying bedrock, the rapid dissolution that takes place during high flow is apparently exerted most vigorously on the bare bedrock surfaces. Some deep fissures can form below the general level of the main cave streams, but these are uncommon. Although mazes are a typical result of cave development by allogenic floodwaters, most large maze caves do not share this origin (e.g. the large network caves of Ukraine and the Black Hills of South Dakota). But the simultaneous enlargement of many alternate flow paths shown in Fig. 5 is valid for these caves as well. The large Q/L ratio necessary for maze origin is the result of small flow distances (L), so all significant openings are enlarged simultaneously. These conditions apply to leakage from overlying or underlying insoluble rocks, and also to mixing or H2S oxidation where aggressiveness is produced right within the aquifer, resulting in cave development at small L values. Examples Several examples from the USA are described here to illustrate the variety of patterns produced by allogenic recharge in contact-karst settings (see locations in Fig. 6). Blue Spring Cave Indiana, is fed mostly by autogenic recharge, but it can be used as a standard for comparison with the extreme floodwater examples that follow. It is almost entirely a jointinfluenced branchwork cave, but it contains two prominent maze sections where blockage of its large main stream has caused local ponding and diversion of water (Fig. 7). The catchment area for the cave is at least 35 km2, which provides enough flow that the hydraulic conditions in these two constricted areas are, in a modest way, similar to those in caves fed by allogenic recharge. Blue Spring Cave is located in the Carboniferous Salem Limestone in an extensive low-relief sinkhole plain. The Salem dips only about 0.3 degree to the west. In the main stream passage, at a junction with a large tributary, extensive collapse has triggered local flooding and development of a network diversion maze around the breakdown (Palmer, 1991). With less than 25 m of overlying rock, the collapse has opened subsidiary fractures in the bedrock that have or ientations quite different from those that guided the main passages (see inset a in Fig. 7). Passages in the network are highly scalloped, indicating high-velocity flow. This maze is located in the typical massive facies of the Salem Limestone. Farther upstream, in a bedded facies, chert beds have caused local passage constrictions, resulting in an anastomotic maze (inset b in Fig. 7). These passages are also prominently scalloped by high-velocity water. Floodwater features are uncommon elsewhere in the cave. The relation between flooding and maze development is clear. Fig. 6. Locations of caves described in this paper. (1) Blue Spring Cave, Indiana; (2) Onesquethaw Cave, New York; (3) = Skull Cave, New York; (4) Mystery Cave, Minnesota; (5) Big Brush Creek Cave, Utah. Fig. 7. Network and anastomotic mazes superimposed on the branchwork pattern of Blue Spring Cave, Indiana (from Palmer, 1991).

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.6 Passages in the network maze (Fig. 7a) are narrow fissures that are c oncentrated at the same level as nearby cave streams, and extending an average of 2 m both above and below the present streams. Many contain deep pools. Since the streams in nearby passag es are now entrenched about 7 m below the initial passage ceilings, the concentration of maze passages at the stream level indicates that the maze developed rather recently within the zone of floodwater fluctuation. A few maze passages extend up to 5 m above the present water table, and as much as 10 m below (as shown by depth measurements in pools that contain outflowing water). Floodwater dissolution can apparently extend not only above the low-flow water table, but below it as well. Onesquethaw Cave in New York State, was used by Palmer (1972) as an ideal example of cave development by allogenic runoff (Fig. 8). It is fed by a single sinking stream with a steeply sloping 3.5 km2 catchment area of shale and sandstone. The cave extends through the Devonian Onondaga Limestone, which is locally folded with considerable folding and faulting, and with local dips up to 25 degrees. The cave reaches no more than 20 m below the overlying land surface. Chert beds up to 20 cm thick have created many abrupt constrictions, which cause the cave to fill completely with water during high flow (approximately once every 2-5 years). Fig. 8 shows the contrast between low-flow and high-flow conditions, with the pressure head and velocity head indicated for the high flow. Many blind fissures extend laterally and upward from the main passages, and long sections of sub-parallel anastomotic tubes have developed around the main stream passage (Fig. 9). Most passages are strongly scalloped, showing high-velo city flow, but scallops are absent in the narro west passages because of abrasion by suspended sediment. During a single flood in 1969, more than 3 m of gravel and sand accumulated in the cave, filling some of the lowlevel fissures and blocking some of the flow routes. Fig. 8. Profile of Onesquethaw Cave, New York, showing contrast between low-flow and high-flow conditions. Note large variation in passage size anirregular hydraulic gradient. Head-loss calculations by M. Palmer (1975); profile from A. Palmer (1972).

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.7 Fig. 9. Anastomotic passages in Onesquethaw Cave, which developed simultaneously by periodic floodwaters. Skull Cave in the Devonian Helderberg limestones of New York, is fed by two small sinking streams that drain shaly limestone capped by glacial till (Fig. 10). The stratal dip is SSW at about 2 degrees. The cave extends no more than 60 m below the overlying surface. Its spring is very inefficient because of blockage by glacial deposits, as well as trapping beneath poorly soluble shaly and dolomitic beds. Except for the upstream part of the entrance passage the en tire cave floods to the ceiling nearly every year. Conspicuous fissure networks have formed in the two areas of most intense flooding (Fig. 11). Blind fissures also extend as much as 20 m upward from the main passages (Fig. 12). Armoring of the floors by sediment encourages upward dissolution during floods. In contrast with Onesquethaw Cave, anastomotic passages are not present, because jointing is much more prominent in the Helderberg limestones. Flooding is also rather static, with high pressure but little velocity. The networks contain no scallops, except in the sections near the main stream. In some areas their walls are coated with clay up to two centimeters thick. Delicate differential solution of th e bedrock fabric behind the clay shows aggressiveness by periodically renewed interstitial floodwater. The fissure networks receive water during every flood and release it from storage as the flood subsides. They are good examples of “karst annexes” as envisioned by Mangin (1975). Fig. 10. Map of Skull Cave, New York showing network mazes formed by static flooding (from Palmer, 1975). Fig. 11. Typical floodwater passage in Skull Cave. Note irregular floor profile.

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.8 Fig. 12. Profiles of the main stream passage and the main east-west passage of Skull Cave, showing floodwater fissures in ceiling. Fig. 14. Calcite saturation index (SI calcite) of allogenic water entering Mystery Cave vs. river stage on West Fork of Root River. Equilibria calculated by Palmer and Palmer (1993) from chemical data by Grow (1986). Approximate conversion from C/Cs in Fig. 4 is SI ~ log [(C/Cs)2.86]. Fig. 13. Mystery Cave, Minnesota, a subterranean meander cutoff of the West Fork of the Root River Cave map by Minnesota Speleological Survey, reproduced with permission. Mystery Cave in the highly fractured Ordovician Stewartville and Dubuque Formations of Minnesota, is a subterranean meander cutoff of the Root River (Fig. 13). The entire cave consists of a fissure network totaling 21 km of mapped passages at a depth of 20-40 m below the surface. The local dip is 0.3-0.8 degrees to the west, and the groundwater flow, which is to the east, is discordant to the bedding along fractures. Although it is not strictly an example of contact karst, as in the previous two examples, the hydraulic and chemical setting, as well as the resulting influence on the cave pattern, are similar to those of true contact karst fed by allogenic streams. Fig. 14 shows the variation in saturation index with stage in the river. Much of the riverbed is on carbonate rock, and only the floodwaters are solutionally aggressive. Big Brush Creek Cave is located at an elevation of 2500 m in the Uinta Mountains of Utah in the Carboniferous Deseret and Humbug Formations (Fig. 15). Brush Creek drains about 65 km2 of high-altitude metamorphic rock, mainly quartzites, and today has a peak discharge of about 30 m3/sec. Much of the water is now tapped for irrigation, but there is evidence for larger flows in the past. Quartzite boulders up to 2 m in diameter have been carried into the cave (Fig. 16). There is still enough water today to fill the entire cave during high flow. The cave is fairly deep by American standards, extending to more than 250 m below the land surface. Its spring lies 650 m below the entrance. To reach th e spring, the cave water must rise along fractures in an overlying sandstone formation, which presents a considerable impediment to floodwaters. Also, many constrictions in the cave have formed by sediment and rafted logs that have accumulated in large piles. The Humbug Formation is an intraclastic carbonate breccia with fragments up to half a meter in diameter. Fractures and partings are very sparse, and most openings are narrow interstices between

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.9 Fig. 15. Partial map of Big Brush Creek Cave, Utah (from Palmer, 1975). Fig. 16. Entrance of Big Brush Creek Cave. clasts. The earliest cave development followed the contact with the underlying dolomitic Deseret Formation, which dips about 15 degrees to the south, but most floodwater dissolution has been upward into the Humbug. Extensive spongework and 3-dimensional mazes have been dissolved in the Humbug (Fig. 17). Spongework and maze development is most prominent in the upstream part of the cave. A long, deep section not shown on the map contains unpleasant concentrations of carbon dioxide from organic decomposition. It is mainly a unitary passage with few of the floodwater features described above. Fig. 17. Spongework in carbonate breccia, Big Brush Creek Cave. Note large quartzite cobbles on floor.

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Arthur N. Palmer / Speleogenesis and Evoluti on of Karst Aquifers, 1 (1) January 2003, p.10 Conclusions With the exception of Brush Creek Cave, all of these examples are located in thin, shallow aquifers, which are most susceptible to the development of floodwater mazes. In Brush Creek Cave the mazes and spongework are most conspicuous in the upstream parts, where the water is most aggressive and overburden pressures are least. The effect of allogenic water on cave patterns is less conspicuous in well-drained, massive limestones. This is especially true for caves that lie at considerable depth below the land surface. Many of the important karst areas of the world fall in this category. Karst aquifers with few fractures and constrictions generally respond only with great enlargement of active cave passages and perhaps simple diversion routes around blockages. Some of the best examples of allogenic recharge are located in the classic Karst of Slovenia, including the wellknown Škocjanske jama (Mihevc, 2001) and Postojnska jama (Šebela, 1998). In these caves the effects of flooding are expressed mainly as irregular patterns with solution pockets and blind fissures. Except for a few local fissure networks there are no extensive mazes like those in the shallower examples described above. How can floodwaters create maze caves? As shown in Fig. 5, it is possible for all openings with sufficient Q/L to enlarge simultaneously at about the same rate. However, the graph also shows that openings below a certain size will not achieve this same rate. The exact threshold size depends on the hydraulic gradient forcing the water into the openings; but it is unlikely that openings less than 0.005 cm in initial width or radius will be competitive. This means that mazes are most likely in shallow aquifers that have experienced considerable erosional unloading, or perhaps also stress from the weight of continental glaciers. The New York caves described above, and possibly also Mystery Cave, have probably benefited from the latter. The network maze in Blue Spring Cave was clearly enhanced by the opening of new fractures around the collapse zone. But why do mazes form, rather than simple parallel diversion routes? Mazes are so common because as the major flow paths enlarge, secondary fissures and tubes are formed by floodwaters injected into the walls of these new routes. The pattern of secondary passages therefore grows outward in a ramifying way until the various blind spurs intersect, forming a maze. In addition, water in the various diversion routes has highly irregular hydraulic gradients during flooding (see Fig. 8). As a result, adjacent flow routes tend to have disparate head values, resulting in the development of crossover passages between the main routes (see Fig. 7b). Many of the smaller features described in this paper are also found to a smaller extent in caves fed by autogenic recharge. These include anastomoses, solution pockets, and blind fissures. Knowing how these features are formed in caves that undergo severe flooding, we can better understand the dynamics of how they are formed in caves fed by allogenic water as well. References Audra P. 1994. Karsts alpines; gense de grands rseaux souterrains. Karstologia Memoires 5, 279 p. Buhmann D. and Dreybrodt W. 1985a. The kinetics of calcite solution and precipitation in geologically relevant situations of karst areas. 1: Open system. Chemical Geology 48, 189-211. Buhmann D. and Dreybrodt W. 1985b. The kinetics of calcite solution and precipitation in geologically relevant situations of karst areas. 2: Closed system. Chemical Geology 53, 109-124. Choppy J. 1991. Notions lmentaires sur le creusement des grottes. Splo-club de Paris, Premire rencontre d’Octobre, 18-23. Coward J.M.H. 1975. Paleohydrology and streamflow simulation of three karst basins in southeastern West Virginia. Ph.D. thesis, McMaster University, 394 p. Dreybrodt W. 1996. Principles of early development of karst c onduits under natural and man-made conditions revealed by mathematical analysis of numerical models. Water Resources Research 32, 2923-2935. Ewers R.O. 1966. Bedding-plane anastomoses and their relation to cavern passages. National Speleological Society Bulletin 28, 133-140. Gabrovšek F. 2000. Evolution of early karst aquifers: from simple principles to complex models. Postjna, ZRC SAZU, 150 p. Gams I. 1965. Types of accelerated corrosion. Problems of the speleological research, International Speleological Congress, Brno, Czechoslovakia, 133-139. Grow S.R. 1986. Water quality in the Forestville Creek karst basin of southeastern Minnesota. M.S. thesis, University of Minnesota, Minneapolis, 229 p. High C.J. 1970. Aspects of the solutional erosion of limestone, with special consideration of

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Arthur N. Palmer / Speleogenesis and Evoluti on of Karst Aquifers, 1 (1) January 2003, p.11 lithological factors. Ph.D. thesis, University of Bristol, U.K., 228 p. Lauritzen S.-E. 1990. Tertiary caves in Norway: a matter of relief and size. British Cave Research Association Transactions 17/1, 31-37. Lismonde B. 2000. Corrosion des cupoles de plafond par les fluctuations de pression de l’air emprison. Karstologia 35, 39-46. Maire R. 1990. La haute montagne calcaire. Karstologia-Mmoires no. 3, 731 p. Mangin A. 1975. Contribution l'tude hydrodynamique des aquifres karstiques. Annales de Splologie 29, 283-332, 495-601 & 30, 21-124. Mihevc A. 2001. Speleogeneza Divaškega krasa. Postojna, Inštitut za razusjivanje krasa, ZRC SAZU, 180 p. Newson M.D. 1971. The role of abrasion in cavern development. Cave Research Group of Great Britain, Transactions 13, 101-107. Palmer A.N. 1972. Dynamics of a sinking stream system, Onesquethaw Cave, New York. National Speleological Society Bulletin 34, 89110. Palmer A.N. 1975. The origin of maze caves. National Speleological Society Bulletin 37, 5676. Palmer A.N. 1981. Hydrochemical controls in the origin of limestone caves. Proceedings of 8th International Speleological Congress, Bowling Green, Kentucky, 120-122. Palmer A.N. 1991. Origin and morphology of limestone caves. Geological Society of America Bulletin 103, 1-21. Palmer A.N. and Palmer M.V. 1993. Mystery Cave, Forestville State Park, Fillmore County, Minnesota. Interpretive report for Minnesota Department of Natural Resources, 97 p. + 20 folio maps. Palmer M.V. 1976. Ground-water flow patterns in limestone solution conduits. M.A. thesis, State University of New York, Oneonta, 150 p. Plummer L.N. and Wigley T.M.L. 1976. The dissolution of calcite in CO2-saturated solutions at 25o C and 1 atmosphere total pressure. Geochimica et Cosmochimica Acta 40, 191-202. Plummer L.N,, Wigley T.M.L. and Parkhurst D.L. 1978. The kinetics of calcite dissolution in CO2water systems at 5o to 60o C and 0.0 to 1.0 atm CO2. American Journal of Science 278, 179216. Experimental data in National Auxiliary Publication Service Document 03209. Rauch H.W. and White W.B. 1970. Lithologic controls on the development of solution porosity in carbonate aquifers. Water Resources Research 6, 1175-1192. Šebela S. 1998. Tectonic structure of Postojnska jama Cave System. Postjna, ZRC SAZU, 112 p. Slabe T. 1995. Cave rocky relief and its speleological significance. Ljubljana, ZRC SAZU, 128 p. Sjberg E.L. 1976. A fundamental equation for calcite dissolution kinetics. Geochimica et Cosmochimica Acta 40, 441-447. Sjberg E.L. and Rickard D.T. 1984. Temperature dependence of calcite dissolution kinetics between 1 and 62o C at pH 2.7 to 8.4 in aqueous solutions. Geochimica et Cosmochimica Acta 48, 485-493. Smith D.I. and Newson M.D. 1974. The dynamics of solutional and mechanical erosion in limestone catchments on the Mendip Hills, Somerset. In K.J. Gregory and D.E. Walling (eds.), Fluvial processes in instrumented watersheds. Institute of British Geographers, Special Publication 6, 155-167.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal www.speleogenesis.info Speleogenesis in carbonate rocks Arthur N. Palmer Department of Earth Sciences, State University of New York, Oneonta, NY 13820, USA Re-published by permission from: Gabrovšek, F. (Ed.), Evolution of karst: from prekarst to cessation. Postojna-Ljubljana, Zalozba ZRC, 43-60. Abstract This paper outlines the current views on cav e origin in carbonate rocks, combining id eas from a variety of sources. A typical dissolution cave develops in several stages that grade smoothly from one to the next: (1) Initial openings are slowly enlarged by water that is nearly at solutional equilibri um with the local bedrock. (2) As the early routes enlarge, t hose with the greatest amount of flow grow fastest. (3) These favoured routes eventually become wide enough that groundwater is able to retain most of its solut ional aggressiveness throughout the entire distance to the spring outle ts. This breakthrough time usually requires times on the order of 104 to 105 years and ends the inception phase of spel eogenesis. (4) Discharge along these selected routes increases rapidly, allowing the m to enlarge into cave passages rather uniformly over their entir e length. Maximum enlargement rates are roughly 0.001-0.1 cm/yr, depending on the local water chemistry and lithology. (5) The cave acquires a distinct passage pa ttern that depends on the natu re of groundwater recharge, the geologic setting, and the erosional hist ory of the region. Branchwork patterns dominate in most carbo nate aquifers. Maze caves are produced by any of the following: steep hydraulic gradients (e.g. during floods), short flow paths, un iform recharge to many openings, and mixing of wate rs that contrast in chemistry. (6) Enlargement rate usually decreases as passages become air-filled, owing to loss of aggressiveness as carbon di oxide escapes through openings to the surface. (7) The cave typi cally evolves by diversion of water to new and lower routes as the fluvial base leve l drops. (8) The cave is eventually destroyed by roof collapse and by intersection of passages by surface erosion. At an y given time, different parts of the same cave may be experie ncing different stages in this sequence. Keywords: cave origin in carbonate rocks Introduction Caves are present in all rather pure carbonate rocks that are in geologic settings and climates that allow abundant groundwater recharge. For this reason, it is clear that cave origin requires no special chemical mechanism beyond the circulation of meteoric groundwater. Dissolution caves can form by other processes, but this is the common speleogenetic mode in mo st carbonate aquifers and is the main topic of this paper. Most of the concepts presented here are not new, but, where possible, alternate viewpoints are given in the hope of encouraging further discussion. Cave inception Speleogenesis requires one basic thing: Groundwater must dissolve the bedrock rapidly enough to produce caves before the rock is removed by surface erosion. This requires the through-flow of large amounts of solutionally aggressive water along stable flow paths. The earliest stages At great depth beneath the surface there is very little groundwater flow because openings in the rock are narrow and few, and hydraulic gradients are feeble. But as uplift and erosion expose these rocks near the surface, increasing amounts of groundwater pass through them. Along any given flow path, the solutional enlargement rate is controlled by a simple mass balance. The mass removed from the walls of the growing conduits is equal to that which is carried away in solution. The increase in volume thus depends on how much water passes through the conduit, and how rapidly the water dissolves the rock In other words, the two major controls are discharge and chemical kinetics. Early in the development of a carbonate aquifer, all groundwater becomes nearly saturated with dissolved calcite and/or dolomite before it emerges at the surface. The total amount of rock removed along any flow path is nearly independent of chemical kinetics, because the water has enough time to equilibrate with the rock, regardless of

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.2 dissolution rates. The saturation concentration depends on the minerals present, CO2 concentration, type of system (open vs. closed), temperature, and interactions with other dissolved components. All these show considerable variation, both spatially and temporally, but it is unlikely that there will be major differences between neighboring flow paths within a given aquifer. In contrast, there are great variations in discharge from one flow path to another – and this is the main control over which the early paths are able to grow into caves. Most dissolution takes place at the upstream ends of the flow paths, where aggressive water first enters the carbonate rock. (“Upstream” and “downstream” in the following discussion refer to the up-gradient and down-gradient ends of the system, even where the flow is only laminar seepage.) With time and distance, there is an increase in saturation ratio of the dissolved minerals (actual concentration divided by saturation concentration, C/Cs ). At first the dissolution rate decreases in a roughly linear manner. But as the saturation ratio rises above approximately 60-90% (the exact value depends on temperature and CO2 content), the dissolution rate decreases much more rapidly. The result is that the final approach toward saturation is very slow (see Berner and Morse, 1974; Plummer and Wigley, 1976; Plummer et al., 1978; Dreybrodt, 1990; Palmer, 1991). Dissolution is so rapid in the upstream sections that if the remainder of the dissolution followed the same trend, the water would lose virtually all its aggressiveness after only a short distance of flow. Dissolution would be so slow in the rest of the aquifer that cave development would be almost impossible (Palmer, 1984). Except in the most ideal situations (wide, short fractures with steep hydraulic gradients, e.g. along escarpments), enlargement of the initial openings to cave size would require many millions of years, during which the carbonate rock is likely to be entirely removed by erosion. Interestingly, it would be unlikely for caves to form with either the rapid dissolution at low saturation ratios or the slow dissolution at high saturation ratios. Early slow dissolution along the entire flow path is essential for preparing the way for the rapid dissolution that follows. But the slow dissolution alone cannot enlarge the routes rapidly enough to form caves within a geologically feasible time. Rapid dissolution at low saturation ratios is necessary to achieve this. But, as shown above, the rapid dissolution by itself cannot form caves in most situations. Geological aspects of cave inception The initial width of fissures (e.g. fractures and partings) is perhaps the most uncertain of all field conditions that influence cave inception. By the time a cave is large enough for humans to enter, the evidence has long disappeared. Initial fissure width is a slippery concept, because the widths increase with time even without being dissolved, simply by release of stress as the overlying rocks are eroded away. Field evidence suggests that a minimum initial fissure width of about 0.01 mm is required for cave development (Bcker, 1969). However, this value depends on hy draulic gradient, flow distance, water chemistry, and length of time available, and so the threshold for initial fissure width is not a fixed value, but instead depends on the local setting. To clarify how wide the initial fissures in limestone might be, it is helpful to gather data from relatively insoluble rocks that are approximately as brittle as limestone. Intrusive igneous rocks such as granite should give a close approximation. Water wells in these rocks have fairly small yields, especially at depths of more than 50 m below the surface (Freeze and Cherry, 1979, p. 158). But even with conservative estimates for hydraulic gradient and fissure frequency, the observed well yields require fissures that are roughly 0.1-0.5 mm wide. Surely only a few of the many fissures are this large, but they are important ones, which in soluble rock would grow into caves. Inception horizons were originally defined by Lowe (1992) as beds or stratal interfaces that provide a chemical environment that favours cave development. The presence of pyrite along a geologic contact was cited as a typical example, whereby oxidation of the sulphide to sulphuric acid might give a substantial boost to cave development. One difficulty with this particular example is the deficiency of oxygen in most deep groundwater. Structural and hydraulic factors are also crucial in determining which initial openings are able to develop into caves. The presence of interbedded sulphates within carbonate rocks provides a suitable environment for cave inception. Dissolution of the sulphates can boost porosity, although this process forces calcite to precipitate by the common-ion effect. Because of differences in molar volume, the precipitated calcite usually does not occupy all the porosity generated by dissolution of gypsum or anhydrite. This process is even more potent when limestone, dolomite, and gypsum interact within the same system. As calcite is forced to precipitate, the solubility of gypsum rises to almost 1.5 times more than that of gypsum alone, and the solubility of dolomite rises to several

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.3 times its normal value. Because dolomite dissolves so slowly, the process is drawn out over long distances and times, potentially resulting in long, continuous paths of increased porosity that may pave the way for later cave development. The geochemical process has been validated by field measurements (e.g. Bischoff et al., 1994), but its impact on cave development is still unclear. Breakthrough Eventually the entire length of an incipient cave becomes large enough to allow water to pass all the way through while still retaining most of its aggressiveness. At this time there is a fairly sudden transition (“breakthrough”) to rapid dissolution throughout the entire flow path. From then on, the entire route enlarges rapidly at a roughly uniform rate of about 0.001-0.01 cm/yr, depending on the water chemistry. This rate varies with the amount of turbulence, but only at low saturation ratios (Plummer and Wigley, 1976; White, 1984). At the high saturation ratios of most karst water the effect is minor. In mature caves, abrasion by coarse sediment load can incr ease local rates of cave development (Smith and Newson, 1974). These factors are insignificant compared to the truly great differences in growth rate that distinguish true cave passages with low saturation ratios from narrow flow paths whose water is nearly saturated with dissolved carbonates. Fig. 1 shows the mean enlargement rate in an ideal fissure as a function of discharge and flow length. The steep parts of the curves represent the slow dissolution rates governed by the mass balance, and the horizontal segments at the top represent the rapid dissolution controlled mainly by kinetics (Palmer, 1991). Because the enlargement rates are not uniform throughout the fissure, the rates shown in Fig. 1 cannot be translated directly into the time required for an incipient cave to reach breakthrough. To do this, finite-difference modelling is necessary. The results are shown in Fig. 2. The time required for chemical breakthrough can be considered the “gestation time” through which an incipient cave must pass in order to allow it to grow into a true cave. It is difficult to specify exactly when this time begins. In some ways, it involves the entire age of the carbonate aquifer, if one includes all the effects of early diagenesis, burial, and uplift in order to reach its present state (Klimchouk and Ford, 2000). But before cave growth can truly begin, there must be a substantial hydraulic gradient. Thus it is customary to start the clock when the carbonate rock is first exposed above base level, at the time when both recharge zones and discharge zones are well defined. Computer models can track the growth of idealized fissures of specified initial width, length, hydraulic gradient, and chemical attributes. These show that the breakthrough time is approximately proportional to w-3 ( i / L )-1.4 P-1, where w = initial fissure width, i = mean hydraulic gradient, L = flow distance, and P = initial PCO2 (Palmer, 1988, 1991). Dreybrodt (1996) provided an analytical derivation arriving at nearly the same functional relationships. Fig. 1. Mean enlargement rate of a fissure in limestone, as a function of discharge (Q) and flow length (L). Q = discharge per metre of fissure height (long dimension of fissure cross section). Assumptions include closed conditions, T = 10o C, initial PCO2 = 0.01 atm. (See Palmer, 1991.) Fig. 2. Approximate breakthrough times for cave inception along fissures in limestone. The main part of the graph shows closed conditions at T = 10o C and initial PCO2 = 1%. Variation of breakthrough time with initial fissure width, temperature, and initial PCO2 are shown. The combined variable i/L represents the hydraulic gradient ( h/L) divided by flow distance (L). Modified from Palmer (1991). See also Dreybrodt (1996).

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.4 Laminar discharge through the fissure is proportional to w3 i which is essentially the inverse of two of the most important variables that determine breakthrough time. Thus the paths that develop most rapidly into caves are those with high discharge and short flow distance. High PCO2 is also favourable, as long as CO2 is not lost by degassing. Temperature plays a complex role in determining how long it takes for breakthrough to occur. Higher temperatures speed the chemical reactions, but in long flow systems this can increase the breakthrough time by depleting most of the water’s solutional capacity in the upstream parts, leaving less for the downstream parts. High temperature increases the flow velocity by reducing the viscosity of the water, but it also decreases the amount of limestone or dolomite that can be dissolved. The net result is an increase in breakthrough time with rising temperature. However, another complication is that in warmer climates the CO2 production in the soil is greater, which shortens breakthrough times. As shown in Fig. 2, breakthrough time decreases as much as 5 times if the CO2 consumed by carbonate dissolution is quickly replaced, for example when the water is in close contact with a CO2 source such as soil. This is rare. In general, the early phase of growth takes place in an approximately closed system, where CO2 is used up as dissolution proceeds. In caves with open atmospheres, CO2 is likely to be lost by air exchange with the surface, which more than offsets the apparent advantage of the open system. Fig. 2 shows that initial fissures 0.01-0.1 cm wide would require no more than a few thousand or tens of thousands of years to reach the maximum enlargement rates, from the time aggressive groundwater first begins to flow through the limestone. For example, in a fissure 1 kilometre long, with an initial width of 0.02 cm, hydraulic gradient of 0.02 (20 m/km), PCO2 of 0.05 atm, temperature of 10o C, and closed to further uptake of CO2, the maximum rate of enlargement is reached in about 30,000 years. These conditions are typical, perhaps even conservative. Lab work and computer modelling by Dreybrodt (1990, 1996) suggest even shorter breakthrough times that are probably more valid. Acids can also be generated within passages by oxidation of organic compounds in the water or iron sulphide in the bedrock, diminishing the breakthrough times. Water chemistry and flow vary with the seasons, but their effects average out over the years. Time required for a cave to reach traversable size Beyond the breakthrough time, growth rate of a cave depends chiefly on the nature of its water input. In dense, rather pure limestone, the rate of wall retreat ( S ) can be estimated with the equation S = 11.7 k (1 – C/Cs )n cm/yr where C/Cs is the saturation ratio, k is a reaction coefficient, and n is the reaction order (see Palmer, 1991 for units and further details). Values for k and n vary with PCO2, and k also varies with temperature. For quick applications, some representative averages can be given. At C/Cs < ~0.7, k and n are approximately 0.015 and 1.7 respectively. At C/Cs > ~0.7, k and n are roughly 0.24 and 4 respectively. Because (1C/Cs ) is less than 1, the larger exponent gives a smaller value of S For example, water that collects on insoluble rock and then flows as a sinking stream directly into a limestone cave usually has a PCO2 of about 0.001-0.005 atm. This value is higher than that of the outside atmosphere (0.00036 atm) because even though the stream is open to the atmosphere, it acquires CO2 from seepage that enters the stream through the soil. At ponors, most sinking streams have saturation ratios of about 0.1-0.5. Under these conditions, limestone surfaces in the cave will dissolve as fast as 0.15 cm/yr. Ideally, a water-filled cave can increase its diameter up to 2-3 m in 1000 years. (The diameter increases at twice the rate of wall retreat, S .) Measurements with dial micrometers, repeated over several years, have verified these rates in caves fed by sinking streams (High, 1970; Coward, 1975). In contrast, many caves are fed by water that infiltrates through soil and reaches the caves only after having traveled for a considerable distance along the soil-limestone contact and through narrow fissures in the epik arst. This water has a high PCO2 (about 0.01-0.05 atm) but has a high saturation value, usually about 0.75-0.95 by the time it reaches the caves. The diameter of a waterfilled passage grows no more than about 20 cm per 1000 years under those conditions. Organization of conduits It has been shown that caves in a typical karst aquifer are able to form only along flow paths that increase their discharge with time. This can be achieved in either of two ways:

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.5 !" By increasing the flow efficiency in a system with a fixed head difference. An example is leakage of water from a stream or other body of water that drains to a lower outlet. As the initial fissures widen, the discharge rises dramatically. The upstream head begins to decrease only when the conduit becomes large enough to transmit the entire stream flow. By that time breakthrough has already taken place. !" By increasing the catchment area that feeds an incipient cave passage. At first, water drains into the growing caves as widely dispersed seepage. Dolines form by subsidence into the rapidly growing voids at the soil/bedrock interface. As dolines increase their catchme nt area, their meanannual discharge increases to the caves that they feed. Discharge to th e caves increases in an irregular manner, much less rapidly than in routes fed by leaking streambeds, and hydraulic gradients decrease with time, even during the earliest periods of growth. The difference between these two systems is important. Because the routes fed by surface streams can increase their flow much more rapidly, they are usually the first parts of a cave to form. Passages fed by depressions of limited catchment area require longer times to form, and they usually join the earlier passages as tributaries of a branchwork system. The fi rst passages to form in a cave are usually short and direct, except where short paths are prohibited by the geologic setting. With time, these early p assages serve as low-head targets for passages having more remote recharge sources (Ford and Ewers, 1978; Ford et al., 2000). Less time is required for a cave to grow in small steps (i.e. where new, relatively short upstream passages link to earlier downstream ones) than for a single long passage to form. This is partly due to the non-linear relation between breakthrough time and flow distance. Although the growth of any single passage propagates in the downstream direction, the overall system grows in the upstream direction, away from the springs, by addition of new passages (Ewers, 1982; Ford et al., 2000). A typical sequence is shown in Fig. 3. Assume, for simplicity, that passage segments B-A and C-B have identical lengths and gradients. The breakthrough time for a single passage from C to A would be ( LC-A / LB-A)1.4 longer than the breakthrough time for either of the two segments – i.e. about 2.6 times longer. This is 30% longer than it would take for segments B-A and C-B to reach breakthrough separately, one after the other. Just as importantly, the gradient of C-B would normally be less than that of B-A until the head dropped in B-A (Ford et al, 2000). The tributary from doline (D) has a smaller catchment area and is slower to reach cave dimensions. Fig. 3. Evolution of a typical branchwork cave by successive piracy of sinking streams and development of recharge sources through dolines. Segment B-A forms first because of the short pa th length and steep gradient. Segment C-B links up later, aided by steepening of the gradient as segment B-A develops. (C-B does not necessarily join B-A at point B.) The passage from doline D is last to form b ecause of its limited catchment area. See Ford and Ewers (1978) and Ford et al. (2000) for descriptions of hardware models that illustrate this concept. Since the flow of water through carbonate aquifers is controlled partly by the history of river entrenchment, the vertical arrangement of cave passages also reflects this control. The traditional view is that the largest passages are formed when base level is relatively static (Sweeting, 1950; Davies, 1960). At such times, rivers develop floodplains, and springs are held at fairly constant elevations for lengthy periods of time. Meanwhile the passages that feed the sp rings are able to grow large. In contrast, passages that form during rapid river entrenchment are usually small. The major passages form different levels, which in most cases decrease in age downward. Fluvial aggradation may cause some or all neighboring cave passages to fill with sediment over the vertical range of base-level rise. This conceptual model h as been well validated in Mammoth Cave, Kentucky (Palmer, 1989; Granger et al., 2001). However, in many caves there are several complications that disrupt this simple interpretation. Vadose passages may be perched on insoluble strata and grow to large size above base level. Most phreatic passages contain vertical loops

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.6 that descend far below the local base level. Some phreatic caves follow favourable stratigraphic units such as zones of former su lphates. Even the ideal cave levels controlled by pauses in fluvial entrenchment are not perfectly “level”. For this reason, many people prefer to call them storeys or tiers and either of these terms is preferred in general applications. However, the term cave level is still appropriate where there is a clear relation to fluvial base level. But the critical elevation is not the average elevation of a phreatic passage, but instead where there is a clear transition from vadose to phreatic morphology (for example, a transition from canyon to tube). This transition is not clear in some passages. Fig. 4 is an idealized profile through a multistoreyed cave, as described by Ford (1971). Three main stages of cave development are shown, with decreasing loop amplitudes from the highest storey to the lowest. This is not a characteristic of all multi-storeyed caves, but it is a conceptual ideal. Ford (1971) ascribed the decrease in amplitude to increasing fissure frequency in the host rock with time. Fissures are sparse at first, and passages are constrained to only a few deeply descending loops. As erosional unloading and cave development persist, fissures become more numerous until eventually the passages are able to form more or less along the water table, with minimal phreatic looping. In some caves the greater amplitude of loops in upper passages is instead caused by floodwaters, which superpose ungraded, looping bypass routes around low-flow routes that have more uniform gradients (Palmer, 1972). In the same vein, on the basis of studies in the Alps, Audra (1994) and Huselmann et al. (2001) ascribe an epiphreatic origin to looping passages. Fig. 4. Vertical layout of a typical cave, showing decreasing amplitude of phreatic loops with depth. This trend has been interpreted by Ford (1971) and Ford and Ewers (1978) to be the result of increasing fissure frequency with time. Successive positions of the water table are shown as gray lines. Some researchers consider these lines to represent the upper extent of epiphreatic flow (see text). The earliest passages in a cave system (usually fed by sinking streams) may not show a clear distinction between vadose and phreatic development, because their discharge fluctuates a great deal, and because the initial potentiometric surface is relatively high. As a result, most of these passages are subjected to a variety of flow conditions – phreatic at first, and then vadose and epiphreatic at later times. Still, many of them show a fairly sharp transition from vadose canyons (with continuous downward trends) to phreatic tubes (with low gradients and usually irregular looping profiles). This transition is more sharply defined in secondary passages fed by karst depressions of limited catchment area, because the flow is more uniform with time and the water sources are usually well above the potentiometric surface. Because of their gravitational flow, many vadose passages have a strong down-dip component, especially those in well-bedded rocks. Phreatic passages show no consistent relation to the dip, except where that is the only path to potential outlets, or where prominent fractures also extend in that direction. In well-bedded rocks, the intersection between the dipping beds and lowgradient water table encourage many phreatic passages to develop nearly along the strike of the beds. These relationships tend to be obscure where the geologic structure is complex. Origin of branching systems Branching cave patterns are by far the most common for several reasons: !" As passages enlarge, the local hydraulic head within them decreases. Groundwater flows from surrounding smaller openings, where the potentiometric surface is higher, toward the low heads of the early conduits. !" Vadose passages have no inherent tendency to converge, because they are hydraulically independent. However, the structures that they follow often intersect, forcing independent streams to join as tributaries. Examples include intersecting fractures, and synclinal structures in bedding-plane partings. !" Water from broad recharge areas converges toward outlets of limited extent, generally stream valleys, which causes a natural tendency for conduits to converge simply by competition for space. After two streams have converged, there is little opportunity for them to diverge farther downstream. The exception is in the vicinity of the spring outlet, where local distributary systems may develop because of collapse, backflooding, and widening of fissures by erosional stress release.

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.7 Maze development Besides branchworks, most other caves are mazes in which all the passages form more or less simultaneously. A maze cave can form only if the growth rate is similar along many alternate flow paths. This can happen if all passages evolve simultaneously at the maximum rate shown in Fig. 1. If the ratio of discharge to flow distance ( Q / L ) is large in many alternate flow routes, they will enlarge at roughly the same rate (Palmer, 1991). Specifically, this condition is achieved if Q / rL > 0.001 (cgs units), where r = mean conduit radius. In fissures, this condition is reached if Q / bL > 0.001, where b = long dimension of the fissure cross section, perpendicular to the narrow dimension w Specific settings where this condition is met include: A. High-discharge or high-gradient flow during floods. Water is forced into all fissures in adjacent carbonate rocks under steep gradients, causing them to enlarge at approximately the maximum possible rate (Palmer, 2001). This process is most active in the vicinity of constrictions in the main stream passages, which result fro m collapse, sediment chokes, or poorly soluble strata. B. Short flow paths from where the water first enters the soluble rock. Because of the short flow distances, all fissures ex cept for the narrowest enlarge simultaneously at similar rates. The epikarst is an example. Network mazes are also formed by recharge through a permeable but insoluble material such as quartz sandstone (Palmer, 1975, 2000). C. Uniform recharge to all fissures, regardless of their width. This can be achieved by seepage through porous, insoluble materials, as in B above. D. Sustained high gradients, for example beneath dams. E. Mixing zones, where the groundwater aggressiveness is locally boosted by mixing of waters of contrasting CO2 content or salinity, or by oxidation of sulphide-rich water. Over short flow distances, many alternate routes are enlarged. Mixing of waters of varied CO2 content can decrease breakthrough times, but large differences in CO2 concentration are necessary (Gabrovšek, 2000). The differences in maze types depend partly on geologic structure. Network mazes consist of intersecting fissures, with a pattern resembling city streets. They require many intersecting fractures (joints or faults), which are typical of massive or thick-bedded rock. Most are formed by processes B, C, or E above. Anastomotic mazes have a braided pattern of intersecting tubes, usually arranged two-dimensionally along a single parting or fault. They are nearly all formed by process A above. Spongework mazes form where primary (matrix) porosity is dominant. In pattern they resemble the intersecting holes in a sponge. Most of them form by process E, and less commonly by process A. A two-dimensional variety can form along bedding-plane partings. Ramiform mazes consist of rooms with offshoots extending outward from them at various elevations. They usually include areas of network or spongework maze development and are formed mainly by process E. Many network and anastomotic mazes, and a few spongework mazes, are merely superimposed on a basic branchwork pattern and represent only part of the entire cave development. Fig. 5 provides a summary of typical cave patterns, showing their relation to source of aggressive water and to dominant structural characteristics. Supporting evidence from computer models Finite-difference computer models support and clarify some of these relationships. Conspicuously absent from the list of ways to form maze caves is slow groundwater flow through artesian aquifers. This origin seems logical, and many maze caves are indeed located in aquifers that are partly artesian. However, artesian conditions by themselves do not produce maze caves. Modelling by Palmer (1991) showed that different-sized branches of a loop are least likely to enlarge at the same rate in slowmoving water near saturation. Dreybrodt and Siemers (2000) supported this idea by showing that as breakthrough time increases, passages tend to become unitary and exhibit less complexity. Modelling by Clemens et al. (1997) verified the development of network mazes by uniform seepage through an insoluble caprock, as described in B above. The insoluble cap encourages maze development because it is permeable, rather than a confining unit. Conduit growth and modification At the breakthrough time, when an incipient cave reaches its maximum growth rate, several other changes take place more or less simultaneously (White, 1977). The cave water changes from laminar to turbulent, which increases the solution rate slightly (see earlier discussion). The flow also becomes competent enough to transport detrital sediment. For example, it is able to carry away the soil that subsides into caves through karst depressions, allowing the depressions to grow more rapidly. The sediment load can also help to enlarge

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.8 Fig. 5. Common patterns of solutional caves. Dot sizes show the relative abundance of cave types in each of the listed categories. Single-passage caves are rudimentary or fragmentary versions of those shown here.caves by mechanical abrasion, but, in places, sediment accumulates in thick beds that retard dissolution and erosion. Where sediment accumulates, upward dissolution by paragenesis is a possible consequence, especi ally in caves enlarged by periodic floodwaters. However, water within the sediment is often undersaturated and can still dissolve the underlying rock (Vaughan et al., 1998). When a cave is able to transmit the entire flow from its recharge area, the average flow can increase no further. Instead the head within the passage decreases as the cro ss section continues to enlarge. Much of the upstream part of the cave becomes vadose, and streams may entrench canyons in the passage floors. As caves acquire entrances that allow air exchange with the surface, many free-surface cave streams lose part of their aggressiveness. Inflowing water is fairly rich in soil-derived CO2, and may acquire even more by oxidation of organic materials as it flows through the caves (Bray, 1972). Loss of CO2 through entrances and other openings can drive the stream water to supersaturation with dissolved calcite or dolomite, so that many vadose cave streams are aggressive only during high flow. Some vadose stream channels even acquire a thin coating of calcite in sections of supercritical flow during dry seasons. These deposits are usually removed during the following wet season, but with only a small net amount of solutional entrenchment each year. Measurements in caves of New York State show that the overall entrenchment rate of active stream canyons of this type can be as slow as 10-20 mm per thousand years (Palmer, 1996), despite the continuous flow of water. During six months of continuous monitoring in the largest stream in Mammoth Cave, Meiman and Groves (1997) found that 70% of the passage enlargement took place during the highest 7% of flow. Dating of cave sediments by 26Al/10Be isotope ratios in quartz-rich cave sediment is a powerful tool for interpreting rates of cave development. Usually this sediment is deposited by the most recent active stream flow and gives a minimum age for the passage. At Mammoth Cave, 26Al/10Be dating suggests that the development of each passage level required at least 105 years (Granger et al., 2001). This value is compatible with the range of probable times required for breakthrough (Fig. 2) and for later enlargement to the present diameters of about 5-10 m in the major passages. Headward erosion of resistant beds by cave streams can require a surprisingly long time. For example, sediment on ledges above an entrenched canyon in Mammoth Cave were dated at 1.13 million years, validated by samples at similar elevations elsewhere in the cave (Granger et al., 2001). The passage is floored by a metre-thick sequence of shaly and cherty limestone, which has been breached by a deep canyon that post-dates the sediment. Headward entr enchment has progressed

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Arthur N. Palmer / Speleogenesis and Evolution of Karst Aquifers, 1 (1) January 2003, p.9 only 360 m along the passage, and only about half of that has occurred upstream from the sampling site. The entrenching stream is still active today and is quite capable of transporting gravel. The rate of headward entrenchment app ears to be less than half a metre per thousand years. But under favourable conditions, diversion of passages from one level to another can take place rather rapidly. Post-glacial diversion of water in New York State caves h as formed traversable passages up to a metre in diameter and 200 m long since the last glacial retreat about 13,000 years ago (Mylroie, 1977). In many vadose canyons throughout the world, examples can be seen where loops or cutoffs have developed along prominent bedding-plane partings exposed in the canyon floor (Fig. 6). As a result, the floor of the upper level coincides with the ceiling of the lower level. The new passage must develop before the parting is bypassed by deepening of the original canyon. This implies that the breakthrough time for the diversion route is virtually nil, allowing the new narrow path to enlarge competitively with the old wellestablished one. Most such diversions are short. As the land surface becomes dissected by erosion, patterns of groundwater recharge change. The few large initial water sources may be divided into many smaller ones. Vadose water must travel increasingly greater distances to reach the water table, and extensive complexes of vadose canyons and shafts can form. The resulting pattern of active cave streams is much denser than that of the original surface drainage. Growing dolines eventually form a continuous karst surface. Eventually the only surface st reams that retain their flow are the main entrenched rivers and the ephemeral upstream ends of sinking streams. Fig. 6. Stream diversion in an entrenching vadose canyon. The lower loop illustrates nearly zero breakthrough time along the guiding bedding-plane parting, as shown by the minimal entrenchment of segment A below the lower parting. This is a common occurrence, especially in well-bedded carbonates, but it is not a general rule. The final stage As the land erodes, the surface intersects underlying cave passages, segmenting them and eventually destroying them entirely. Evidence for the cave may persist for a while as a canyon-like feature or a rubbly zone of collapsed blocks. This final episode in the life of a cave passage usually occupies tens of thousands or even hundreds of thousands of years. However, newer passages continue to develop where the soluble rock extends to lower elevations. In dipping carbonate rocks, new areas of rock are uncovered by erosion at about the same rate as they are eroded away in the up-dip areas. This process ends when the entire soluble rock in the cave region is eroded away. References Audra P. 1994. Karsts alpins; gense des grands rseaux souterrains. Karstologia Mmoirs 5, 279 p. Berner R.A. and Morse J.W. 1974. Dissolution kinetics of calcium carbonate in sea water; IV: Theory of calcite dissolution. American Journal of Science 274, 108-134. Bischoff J.L., Julia R., Shanks W.C. and Rosenbauer R.J. 1994. Karstification without carbonic acid; bedrock dissolution by gypsumdriven dedolomitization. Geology 22/11, 995998. Bcker T. 1969. Karstic water research in Hungary. International Association of Scientific Hydrology Bulletin 14, 4-12. Bray L.G. 1972. Preliminary oxidation studies on some cave waters from south Wales. Cave Research Group of Great Britain, Transactions 14, 59-66. Clemens T., Hckinghaus D., Sauter M., Liedl R. and Teutsch G. 1997. Simulation of the evolution of maze caves. 12th International Congress of Speleology, and 6th Conference on Limestone Hydrology and Fissured Media, La Chaux-de-Fonds, Switzerland, 2, 65-68. Coward J.M.H. 1975. Paleohydrology and streamflow simulation of three karst basins in southeastern West Virginia. Ph.D. dissertation, McMaster University, Hamilton, Ontario, Canada, 394 p. Davies W.E. 1960. Origin of caves in folded limestone. National Speleological Society Bulletin 22/1, 5-18. Dreybrodt W. 1990. The role of dissolution kinetics in the development of karst aquifers in

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Arthur N. Palmer / Speleogenesis and Evoluti on of Karst Aquifers, 1 (1) January 2003, p.10 limestone: a model simulation of karst evolution. Journal of Geology 98/5, 639-655. Dreybrodt W. 1996. Principles of early development of karst c onduits under natural and man-made conditions revealed by mathematical analysis of numerical models. Water Resources Research 32, 2923-2935. Dreybrodt W. and J. Siemers 2000. Cave evolution on two-dimensional networks of primary fractures in limestone. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of Karst Aquifers. Huntsville, Ala.: National Speleological Society, 201-211. Ewers R.O. 1982. Cavern development in the dimensions of length and breadth. Ph.D. dissertation, McMaster University, Hamilton, Ontario, 398 p. Ford D.C. 1971. Geologic structure and a new explanation of limestone cavern genesis. Transactions of the Cave Research Group of Great Britain 13/2, 81-94. Ford D.C. and Ewers R.O. 1978. The development of limestone cave systems in the dimensions of length and depth. Canadi an Journal of Earth Sciences 15, 1783-1798. Ford D.C., Lauritzen S.-E. and Ewers R.O. 2000. Modeling of initiation and propagation of single conduits and networks. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of Karst Aquifers. Huntsville, Ala., National Speleological Society, 175-183. Freeze, R.A., and J.A. Cherry. 1979: Groundwater. Englewood Cliffs, N.J., Prentice-Hall, 604 p. Gabrovšek F. 2000. Evolution of early karst aquifers: from simple principles to complex models. Postojna, Slovenia, Inštitut za razusjivanje krasa ZRC SAZU, 150 p. Granger D.E., Fabel D. and Palmer A.N. 2001. Pliocene-Pleistocene incision of the Green River, Kentucky, determined from radioactive decay of 26Al and 10Be in Mammoth Cave sediments. Geological Society of America Bulletin 113/7, 825-836. Huselmann P., Jeannin P.-Y. and Monbaron M. 2001. Relation between al pine paleogeography and cave genesis: the case of the cave system of Sieben Hengste (Berne, Switzerland). In Huselmann P. and Monbaron M. (Eds.), Cave genesis in the Alpine belt. Proceedings of 1st Workshop for Alpine Speleogenesis, University of Fribourg, Fribourg, Switzerland, 115-123. High C.J. 1970. Aspects of the solutional erosion of limestone, with special consideration of lithological factors. Ph.D. dissertation, University of Bristol, Bristol, U.K., 228 p. Klimchouk A. and Ford D.C. 2000. Types of karst and evolution of hydrogeologic setting. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of Karst Aquifers. Huntsville, Ala., National Speleological Society, 45-53. Lowe D.J. 1992. The origin of limestone caverns: an inception horizon hypothesis. Ph.D. dissertation, Manchester Polytechnic University, U.K., 512 p. Meiman J. and Groves C. 1997. Magnitude/ frequency analysis of cave passage development in the Central Kentucky Karst. Proceedings of 6th Annual Mammoth Cave National Park Science Conference, 11-13. Mylroie J.E. 1977. Speleogenesis and karst geomorphology of the Helderberg Plateau, Schoharie County, New York. Ph.D. dissertation, Rensselaer Polytechnic Institute, Troy, N.Y., 336 p. Palmer A.N. 1972. Dynamics of a sinking stream system, Onesquethaw Cave, New York. National Speleological Society Bulletin 34/3, 89-110. Palmer A.N. 1975. The origin of maze caves. National Speleological Society Bulletin 37, 5676. Palmer A.N. 1984. Recent trends in karst geomorphology. Journal of Geological Education 32, 247-253. Palmer A.N. 1988. Solutional enlargement of openings in the vicinity of hydraulic structures in karst regions. Dublin, Ohio, Proceedings of 2nd Conference on Environmental Problems in Karst Terranes, Association of Ground Water Scientists and Engineers, 3-13. Palmer A.N. 1989. Geomorphic history of the Mammoth Cave System. In W.B. White and E.L. White (eds.), Karst Hydrology: concepts from the Mammoth Cave area. New York, Van Nostrand Reinhold, p. 317-363. Palmer A.N. 1991. Origin and morphology of limestone caves. Geological Society of America Bulletin 103, 1-21. Palmer A.N. 1996. Rates of limestone dissolution and calcite precipitation in cave streams of eastcentral New York State [abstract]: Northeast Section, Geological Society of America 28/3, 89.

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Arthur N. Palmer / Speleogenesis and Evoluti on of Karst Aquifers, 1 (1) January 2003, p.11 Palmer A.N. 2000. Maze origin by diffuse recharge through overlying formations. In: Klimchouk A., Ford D., Palmer A. and Dreybrodt W. (Eds.), Speleogenesis: Evolution of Karst Aquifers. Huntsville, Ala., National Speleological Society, 387-390. Palmer A.N. 2001. Dynamics of cave development by allogenic water. Acta Carsologica 30/2, 1332. Plummer L.N. and Wigley T.M.L. 1976. The dissolution of calcite in CO2-saturated solutions at 25o C and 1 atmosphere total pressure. Geochimica et Cosmochimica Acta 40, 191-202. Plummer L.N., Wigley T.M.L. and Parkhurst D.L. 1978. The kinetics of calcite dissolution in CO2water systems at 5o to 60o C and 0.0 to 1.0 atm CO2. American Journal of Science 278, 179216. Smith D.I. and Newson M.D. 1974. The dynamics of solutional and mechanical erosion in limestone catchments on the Mendip Hills, Somerset. In: Gregory K.J. and Walling D.E. (Eds.), Fluvial processes in instrumented watersheds. Institute of British Geographers, Special Publication 6, 155-167. Sweeting M.M. 1950. Erosion cycles and limestone caverns in the Ingleborough District of Yorkshire. Geographical Journal 124, 63-78. Vaughan K., Groves C. and Meiman J. 1998. Carbonate chemistry of interstitial fluids within cave stream sediments [abstract]. Conference on Carbon Cycling in Karst, International Geological Correlation Program, Western Kentucky University, Bowling Green, Kentucky, 33-34. White W.B. 1977. Role of solution kinetics in the development of karst aquifers. In: Tolson J.S. and Doyle F.L. (Eds.), Karst hydrogeology. International Association of Hydrogeologists, 12th Memoirs, 503-517. White W.B. 1984. Rate processes: chemical kinetics and kast landform development. In: LaFleur R.G. (Ed.), Groundwater as a geomorphic agent. Boston, Allen and Unwin, 227-248.



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Speleogenesis and Evolution of Karst Aquifers The Virtual Scientific Journal www.speleogenesis.info Patterns of dissolution porosity in carbonate rocks Arthur N. Palmer Department of Earth Sciences, State University of New York Oneonta. NY 13820-4015, USA. E-mail: palmeran@snyoneva.cc.oneonta.edu Re-published by permission from: Palmer, A.N., Palmer, M.V., and Sasowsky, I.D. (e ds.), Karst Modeling: Special Publication 5, The Karst Waters Institute, Charles Town, West Virginia (USA), 71-78. Abstract This paper reviews the hydrochemical proce sses that determine the patterns of caves and other solutional features within carbon ate rocks. The model presented relies on the f unctional relationships expressed by chemi cal mass balances, flow equations, and kine tic expressions for dissolution rate. Although it shares many aspects of purely conceptual models and is backed by field evidence, its quantitative basis places it into the realm of analytical models. The conclusions merely summarize earlier wo rk (mainly Palmer, 1981, 1991). Solutional enlargement of caves and other karst features is highly selective in water that is close to equilibrium with dissolved carbonate minerals, enlarging only the most f avorable openings – i.e. those that transmit the great est discharge. This is char acteristic of long flow paths within a typical karst aq uifer. In contrast, solutional enlargement will be rather uniform along many competing flow paths where there is (1) high discharge, (2) sustained steep hydraulic gradients, (3) short flow paths, or (4) local renewal of aggressive ness by mixing, oxidation of sulfides, etc. These conditions produce maze caves and epikarstic networks. In general, this condition prev ails if Q/rL > 0.001 cm/sec (tubes), or / bL > 0.001 cm/sec (fissures), where Q = discharge, r = tube radius, b = long dimension of fissure cro ss section, and L = distance of flow from where the initial aggressive solution come s in contact with the carbonate rock. Keywords: solution porosity in carbonate rocks, speleogenesis, modeli ng of karst aquifers Introduction Unlike most geologic processes, the origin of dissolution porosity lends itself readily to analytical solutions. Four salient "laws" govern the process: two mass balances (water balance and chemical mass balance) and two kinetic equations (which describe the dissolution rate and the flow rate of water), and in combination they provide a theoretical basis for quantifying the solutional history of karst aquifers. The greatest difficulty is in applying these clean-cut analytical tools to the complex and rather disordered world of geology. It is impossible to model a karst aquifer in all its details, because most of the details are unknown. However, a great deal can be learned about the origin and distribution of dissolution porosity by using the analytical approach to obtain a battery of governing concepts that can be applied to all karst aquifers. This paper summarizes the evolution of a conceptual model whose deta ils were first developed on the basis of field observation and hydraulics, and only later substantiated by chemical kinetics. It applies specifically to carbonate rocks, although the general approach can be modified to fit any geologic setting by substituting the appropriate expressions for kinetics and fluid flow. General distribution of karst porosity The appropriate first step relies on concepts that are well known to everyone. Water, where if first enters a soluble rock, is undersaturated with respect to the minerals in that rock. Its saturation ratio ( C/Cs) with respect to these minerals is at, or close to, zero. ( C = concentration of dissolved mineral, e.g. calcite; Cs = saturation concentration, which is greatly enhanced by acids. Cs for calcite in typical groundwater is about 2-3 mmol/L.) Dissolution is most rapid where C/Cs is lowest, and its rate diminishes as the dissolv ed load increases. The

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.2 2 greatest dissolution occurs where the water first enters the soluble rock and diminishes with flow distance. Thus most of the water's solutional aggressiveness is usually squandered at and near the bedrock surface, simply lowering the surface and widening the upstream ends of fissures that penetrate the rock (Fig. 1). Underground water circulating through the rock has greatly diminished aggressiveness, but unless Cs decreases (e.g. by CO2 degassing, rising temperature, or common-ion effects) the water will never quite reach saturation. Since the flow of water is continuous from recharge source to outlet, and the relatively high C/Cs allows dissolution to proceed only at a slow rate, the dissolution porosity is drawn out along the entire length of the flow paths. Aggressiveness can be maintained by a gradual rise in Cs for example by oxidation of organic compounds. Aggressiveness can also undergo a local burst because of mixing of different waters or by oxi dation of hydrogen sulfide. However, in a typical karst aquifer the solutional porosity tends to form c ontinuous conduits rather than isolated voids (Fig. 1). The diminution of dissolution porosity with dept h in karst aquifers is well documented by borehole data (Fig. 2). Introductory geology textbooks typically portray karst porosity in cross sections that resemble blocks of Tilsit cheese, with the holes elongated along beds and fractures as a gratuitous nod to the influence of geology. This portrayal would be amusing, were it not for the fact that many scientists and engineers, perhaps subconsciously, use the same conceptual view of karst when attempting to solve environmental problems. The real distribution of karst porosity is more complex, but also more predictable. Fig. 1. Generalized cross section of a karst aquifer formed by meteoric water, showing relative distribution of dissolution voids. The apparent porosity is exaggerated in the diagram for clarity. Below the epikarst the total dissolution porosity rarely exceeds a few percent and is commonly far less (see paper by Worthington in this volume). Fig. 2. An example of the variation in karst porosity with depth, as observed in deep boreholes in Herzegovina (Modified from Milanovic, 1981). Within karst aquifers, most of the dissolution porosity consists of condu its, usually arranged in dendritic patterns in which tributaries join each other to produce fewer but larger conduits in the downstream direction. When we visit a cave of this kind, our conception of dissolution porosity can easily be-skewed. The impression is that the overall dissolution porosity must be enormous. However, it is concentrated in only a relatively few conduits, which, if compared to the overall rock volume, yield only a tiny porosity, usually less than a percent (see Worthington, in this volume, for representative examples). In the most complex part of Mammoth Cave, Kentucky, Palmer (1995) calculated the total dissolution porosity to be only 4%, even though the figure includes many more inactive relict passages than active ones. Mo st accessible caves are surrounded by rock in which the vast majority of openings have hardly enlarged at all. The conduits are not surrounded by porous zones, with walls like a sponge, where progressively smaller openings extend indefinitely into the cave wall. The conduits are quite discrete. As shown by the morphology of cave passages, the flow of aggressive water through karst aquifers takes place mainly along fractures and bedding-plane partings, and much less so along primary pores. It is

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.3 3 important to note that fractures and partings diminish in width and number with depth (see Ford and Ewers, 1978, and Ford in this volume). Provided the openings are too small to admit turbulent flow, the flow is described by the Hagen-Poiseuille equation: # $ % & L h b w Q d d 12 !3 (1) where Q = discharge, w = fissure width, b = fissure breadth (long dimension of fissure cross section), != specific weight of water, = dynamic viscosity of water, and dh/dL = hydraulic gradient. Although groundwater can pass through all openings in varying amounts, note the strong dependence of flow on fissure width. Since w diminishes downward, one can assume that the water will preferentially follow shallow paths, far more so than in aquifers whose pore size does not decrease so significantly with depth (for example, gravel). This applies to fractured bedrock aquifers of all types, but especially in soluble rocks where the initial openings are enlarged by the flow. As these openings grow, some will enlarge more rapidly than others. They are the ones that grow to cave size and which eventually dominate the flow pattern within the aquifer. The onset of turbulent flow is often used to distinguish the birth of a karst conduit. This discussion applies mainly to phreatic water. Gravitational vadose water is easily perched for some distance on resistant or relatively insoluble beds, providing a strong lateral component to the flow, interrupted by shafts where the perching beds are breached by fractures. The tendency for conduits not to penetrate far below the potentiometric surface is disrupted to some extent by faulting and folding. In tectonically disturbed areas it is possible for certain flow paths to extend to considerable depths, especially along faults. For example, in a catalog of solutional voids encountered beneath river beds during drilling by the Tennessee Valley Authority (Moneymaker, 1941), the deepest voids are encountered in the folded and faulted Appalachians and in the fault zone of western Kentucky. Ford (1971) emphasized low fissure frequency as the main criterion for why certain caves extend deep beneath the potentiometric surface, whereas Palmer (1969) emphasized fissure width. The two contrasting views are nevertheless compatible. Despite this disruption, the shallowest paths are still the most favorable, even in tectonically disturbed areas, and deep conduits are relatively rare. Pervasive deep flow is likely only in confined settings, or where favorable stratigraphic boundaries allow deep dissolution (for example, along sulfatecarbonate interfaces). Worthington (1991) cites many examples of sulfate-rich springs fed by groundwater that follows deep basin-wide paths. Rates of conduit growth In a fissured aquifer with myriad flow paths, which ones are most likely to enlarge into solution conduits? This is not a trivial question, because the presence and distribution of conduits is one of the most important variables in the assessment of a karst aquifer. Phreatic conduits form potentiometric lows and are the main paths of groundwater discharge, as well as the major paths of contaminant transport. The configuration of vadose channels determines how sources at the surface relate to the points of recharge at the underlying water table. In most carbonate aquifers only a small percentage of flow routes enlarge into turbulent-flow conduits. Early in the flow history of the aquifer, fissures are narrow and the flow is dispersed among many different routes, each with its own overall hydraulic gradient, mean fissure width, total flow length, and mean discharge. Groundwater discharge and velocity are so small that the wate r becomes nearly saturated with dissolved bedrock long before it emerges at the surface. This can be verified by measuring the chemistry of inflowing seepage through narrow openings into accessible caves. Even some substantial flows of several cm3/sec arrive essentially at C/Cs =1.0 after less than 50 m of flow. Therefore, in any single flow route, the rate of dissolutional enlargement depends simply on the chemical mass balance. The mass of material removed from the walls of the opening is equal to the mass removed in solution by the groundwater flow. This rule is independent of the shape of the opening (tube, fissure, etc.). The chemical mass balance can be stated as follows: Mass removed from the walls of the opening = mass carried away in solution. Change in mass with time = # V/ # t = Q # C where = rock density, # V/ # t = change in volume with time, and # C = change in dissolved load over the length of the conduit. A conduit of circular cross section is assumed for convenience; the actual conduit shape is not important, since the functional relationships are the same. Thus, within any single conduit segment of rather uniform dimensions, the rate of wall retreat (S) can be stated () L r C C Q S* +2 56 310, [cm/yr] (2)

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.4 4 Fig. 3. Mean rate of wall retreat (cm/yr) as a function of discharge (Q), flow distance ( L ,), hydraulic gradient ( i ), and tube radius ( r ) in tubular conduits at 10 C and PCO2 = 0.01 atm ( CS = 212 mg/L CaCO3). Values of i/L are valid only for laminar flow. A = region of varied growth rates (Q/rL < 0.001 cm/sec); B = maximum possible solutional growth rate, limited by chemical kinetics. (From Palmer, 1981). where C0 = initial concentration of dissolved bedrock (mg/L), r = conduit radius (cm), and L = conduit length (cm). All other terms are in cgs units. The numerical coefficient converts the result to cm/yr. A uniform radius is not required, since we are concerned only with mean values of S, Q, and r, and again the functional relationships are not affected. Early in the evolution of a karst aquifer, when the water emerges from each opening essentially at saturation, if we assume that C0 = 0, then # C = CS, which is constant at a given temperature and CO2 partial pressure. Under these conditions, the mean rate of wall retreat within each conduit is therefore a linear function of Q/rL. The mean S values plot as the family of lines shown in group A of Fig. 3 (10 C, PCO2 = 0.01 atm). At any value of r, the larger the Q/L ratio, the greater the rate of enlargement. Large Q and/or small L favor rapid enlargement. Dashed lines show the corresponding values of i/L, where i = hydraulic gradient, are also shown on Fig. 3, valid only for laminar flow (derived from the tubular version of eq. 1, where w3b/12 is replaced by + r4/8). The numerical values for S in Fig. 3 are misleading, because (1) they assume that the water enters each conduit at zero saturation; (2) most of the dissolution is concentrated in the upstream end of the conduit, so that the rate throughout the majority of the conduit is less than the mean; and (3) they do not consider mixing or branching between different conduits. However, it is the relative values the general relationships among the terms that concern us here, not the specific numerical values. Actual growth rates are best determined with finitedifference calculations or analytical methods (e.g. Palmer, 1984, 1991;Dreybrodt, 1990,1996; Groves and Howard, 1994a and b; Clemens et al, 1996; Hanna and Rajaram, 1998). The earliest conduit growth usually begins somewhere in (or beyond) the lower left portion of Fig. 3. Growth rates are far too low to allow turbulent-flow conduits to develop in a geologically feasible time. Within any given conduit, the growth rate can increase only if Q increases, since L does not change. If Q does not increase, the enlargement rate will remain static or actually decrease, as shown by the negative slope of the lines in group A as the conduit radius increases. (Conduits of non-circular cross section would have more gently sloping lines in group A.) But the growth rate reaches an upper limit beyond which it cannot rise. This is typically about 0.0010.01 cm/yr, depending on the chemical conditions. So far the analysis h as been focused on narrow openings in which the water emerges near saturation. Now consider a conduit with such a large Q/L ratio that water is able to pass through the conduit while retaining nearly all its aggressiveness. The entire conduit enlarges at a nearly uniform rate (shown as B in Fig. 3), which is a function of the dissolution kinetics, rather than of the mass balance. Rates of wall retreat are almost uniform and independent of Q/L. Experiments by Berner and Morse (1974), Plummer and Wigley (1976), Plummer et al. (1978) show that in turbulent flow the dissolution rate for calcite is governed mainly by the chemical reactions at the wall, rather than by mass transfer within the fluid, and that turbulence and flow velocity have little effect on. Mass transfer does have an effect at low flow rates in limestone conduits (Curl, 1968; Buhmann and Dreybrodt, 1985a and b, Dreybrodt, 1988), and in evaporites at any flow rate, but the pattern of lines in group A in Fig. 3 is not affected. For carbonate rocks, the dissolution rate is expressed by ()n SC C V k A t C / 1 d d [mg/L-sec] (3) where A' = surface area of rock in contact with water, V = water volume, k = reaction coefficient, and n = reaction order. It is more common for the parenthetical term to be expressed as (CS C), but the form shown in eq. 3 avoids the problem of having to adjust the units of k whenever n changes. Values of k and n depend on the acid content of the

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.5 5 water (for example, PCO2), and k also varies with temperature and lithology (Palmer, 1991, derived from the original lab measurements of Plummer et al., 1978, Plummer and Wigley, 1976, Rauch and White, 1977, and other sources). The reaction order (n) increases rather abruptly from 1-2 to 4 or more at C/CS values that range from about 0.6 to 0.9, depending on temperature and PCO2 (Palmer, 1991; Dreybrodt et al. in this volume). In typical groundwater, n = 2 (or 1, in narrow conduits) and k = 0.01 at C/CS <=0.65. At greater C/CS n=4 (occasionally more; see Dreybrodt et al., 1999) and k = 0.1. Because C/CS < 1, an increase in reaction order represents a decrease in dissolution rate. Ironically, this decrease appears to be essential to the origin of nearly all solution conduits (Palmer, 1984). If the low-order (rapid) kinetics were to prevail throughout the initial opening, virtually all the aggressiveness would be consumed within a few meters of flow, except in usually wide fissures, and the growth rate would be so slow in the rest of the conduit that it would never achieve turbulent flow within a geological feasible time. On the other hand, the highorder (slow) kinetics alone are too slow to enlarge the conduits to the size of traversable caves. White (1977) called the change from slow to rapid dissolution a "kinetic trigger" that represents the beginning of true cave development. Thus cave origin enjoys the best of both worlds: slow growth when the aggressiveness must persist for long distances through narrow fissures, and later rapid growth to achieve large size. Combining eqs. 2 and 3, and substituting Q = V/t and A' = 2+rL, gives the following general equation for dissolutional wall retreat in carbonate rocks: () *n SC C k S / 1 56 31 [cm/yr] (4) which is valid for all types of flow and conduit geometries (Palmer, 1991). The maximum rate of wall retreat (aside from occasional mechanical erosion during floods in large conduits) can be determined by this equation, where C/CS = saturation ratio where the water enters the conduit. In Fig. 3, the maximum rate at B is for C/CS = 0. Higher saturation ratios, even as high as 0.9, still provide enlargement rates that are rapid by geological standards, and the overall shape of the graph in Fig. 3 remains valid. The composite graph in Fig. 3 shows both the early laminar flow (S dependent on Q/L) and the late-stage flow (S independent of Q/L). The curvature of the lines where Zone A meets Zone B was determined by finite-difference analysis. It is not realistic to assume that water can pass through an entire aquifer without changing its saturation ratio. The water acquires most of its solute load at the upstream end, especially where it accumulates in small openings that feed the main conduits. The maximum enlargement rate in a conduit is thus limited in part by the value of C0 at its upstream end, which is rarely less than 0.5. HighQ flow can pass through a cave-size conduit for great distances while gaining only a few mg/L of dissolved load. Because of the large Q, this still represents a substantial rate of mass flux. Competition between enlarging flow paths: unitary conduits and branchworks The specific configuration of conduits within the aquifer is determined by the relative growth rates of competing flow paths. The most common situation, where meteoric groundwater passes through a carbonate unit from an upland recharge surface to outlets at lower elevation, can be described as follows: 1. Early in the flow history of the aquifer, the many alternate flow routes have low Q/L ratios and plot in or beyond the lower left comer of Fig. 3. Their Q/L values and rates of conduit growth span a wide range of many orders of magnitude. All growth rates are small, but some will be much greater than others. The dots on Fig. 4 show some representative flow paths early in the aquifer development. These are idealizations, because no single flow path behaves entirely independently. 2. Growth rate can increase only if the Q/L ratio increases. For any given path, this can be achieved only by an increase in discharge. As each opening grows, its discharge tends to increase because more water is able to pass through. But there is a maximum amount of available infiltration, and eventually a conduit can acquire additional Q only at the expense of its neighbors. This is achieved in two ways: (a) As a conduit grows, its hydraulic head decreases (despite increasing Q ) because of the reduced amount of head loss required to transmit the flow. Water is drawn fro m neighboring openings in which the head is greater, which increases the discharge in the major conduits, (b) Sinkholes develop as the major openings and their tributaries enlarge, especially at the upstream ends, funneling water into the largest conduits but bypassing lesser openings. As a result, the openings with the largest initial Q/L or i/L are those that are most likely to acquire increasing flow and to increase their

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.6 6 Fig. 4. Varied growth histories of competing flow paths in the early stages of a karst aquifer. Paths A, B, and C accelerate in growth and reach the maxium rate by increasing their discharge. These are the routes that become major dissolution conduits. In contrast, other paths (e.g. D and E) stagnate at low and usually diminishing enlargement rates. enlargement rate. As shown by the arrows in Fig. 4, some openings are favored over their neighbors and increase rapidly in both Q and S. Others languish with negligible and generally decreasing Q and S. 3. Enhanced discharge is able to increase the enlargement rate only up to a certain point, beyond which the rate becomes insensitive to further increases in Q. The enlargement rate is now limited mainly by the dissolution kinetics. Those relatively few conduits that reach this state grow at approximately the same rapid rate, with only minor differences caused by local variations in chemistry, flow, and passage configuration. Traversable caves are formed by water that has achieved this state. Single-passage stream caves and branchwork caves are the normal result. Branching caves are by far the most common, because of the tendency for convergence of flow toward the relatively low head of the major conduits, and because of fortuitous intersections between p assages. The typical passage pattern is similar to that in Fig. 5a. This example is located in prominently bedded, low-dip strata. Greater structural complexity leads to comparable passage complexity, but the overall cave pattern is usually a branchwork. Uniform dissolution among many competing flow routes: labyrinthine porosity Under certain conditions, nearly every competing flow route enlarges at comparable rates, and a labyrinth of interconnected openings is formed. According to Fig. 3, the only way this can happen is to expose many openings simultaneously to high values of Q/L or i/L. Beyond a certain threshold (Q/rL > 0.001 cm/sec in tubes, Q/bL>0.001 cm/sec in fissures), they will all enlarge at rather similar rates, regardless of size, discharge, gradient, or flow length. The result is a labyrinth of interconnected passages consisting of openings that have grown simultaneously to cave size. Caves formed under these conditions have maze patterns (Fig. 5 b, c, and d). Favorable conditions include the following: Where aggressive water first enters the soluble rock (small values of L). Distances of flow from the entry points are short, and all openings have large Q/L values regardless of opening size or discharge. The most common example is the epikarst, the zone of extensive dissolution in the upper few meters or tens of meters of the soluble rock, located either beneath a soil cover or exposed directly at the surface (Fig. 1). The same approximate result is achieved where water passes through a porous, nonsoluble rock before entering the carbonate rock, forming a maze cave. Network caves formed along intersecting fractures are the most common type (Figure 5b). The flow of water can be downward from the overlying surface or upward as artesian flow from an underlying formation (provided the water has not encountered substantial amounts of carbonate rock beforehand). Fig. 5. Examples of soluti onal cave patterns: (a) Crevice Cave, Missouri, (b) part of Crossroads Cave, Virginia, (c) part of Holoch, Switzerland, (d) main rooms of Carlsbad Cavern, New Mexico. Maps courtesy of Paul Houck, H.H. Douglas, Alfred Bgli, and Cave Research Foundation, respectively. E = entrance.

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.7 7 It is also possible to enlarge many alternate flow routes simultaneously where steep hydraulic gradients are imposed by flooding. This is most noticeable where soluble rock is exposed adjacent to entrenched rivers that flood severely. Water enters the ground as bank storage having high i/L ratios because of the steep gradients and short flow distances. The same effect is achieved in local areas within preexisting caves where flow constrictions (breakdown, sediment fill, interference by relatively insoluble beds) cause floodwater to pond upstream from them. Seasonal or storm-related flooding periodically injects aggressive water into every available opening, enlarging them rapidly and simultaneously. Irregular network caves (western part of Fig. 5b) are formed where vertical or steeply inclined fractures dominate. Anastomotic caves are formed where bedding-plane partings or low-angle faults are the main cave-forming units. Where matrix porosity provides the major flow paths, for example in diagenetically young limestones, breccias, or reef rock, a spongework pattern will form (as illustrated in parts of Fig. 5d). Fig. 6. Many alternate routes are able to enlarge simultaneously at roughly the same rate, regardless of size, if they are all able to sustain hi gh discharges (e.g. because of steep hydraulic gradients) or short flow distances. Maze caves, epikarst, and pervasive labyrinthine porosity are produced in this way. Very small openings, or those with low Q/L or i/L will not be competitive. Where steep hydraulic gradients are sustained (e.g. beneath dams), growth rates rise steeply in all conduits as r increases, as shown by the dashed lines in Fig. 3, until they all reach the maximum rates of wall retreat at or near the top of the graph. A network of similarly enlarged conduits is expected (Palmer, 1988; Dreybrodt, 1996; Bauer et al, 1999). Solutional aggressiveness can be renewed in zones of mixing between chemically contrasting waters, for example between infiltrating high-CO2, freshwater and low-CO2 seawater. Local network and spongework patterns are produced, not only because the flow distances from the source of aggressiveness are short, but because of the diffuse nature of most water flow under these circumstances. Carbonate aquifers in coastal and island settings are noted for this kind of porosity (Back and others, 1984; Mylroie and Carew, 1990; Mylroie and Vacher, 1999). Mixing between rising (often thermal) waters and meteoric water is capable of producing considerable aggressiveness. Caves formed in this way tend to have network or irregular ramiform patterns (Fig. 5 b and d). Oxidation of rising hydrogen sulfide rising from depth into oxygen-rich zones at or near the water table produces a burst of localized dissolution. This process usually results in network and ramiform caves, consisting of large irregular rooms with sequential branches exiting to the surface (Fig. 5d). The outflow usually coalesces into discrete conduits as aggressiveness is lost and flow length increases. Summary This simplified view of the distribution of karst porosity leads to several conclusions, which are summarized below: Karst porosity is greatest near the land surface in areas of groundwater recharge (for example, in the epikarst). It diminishes in the downflow direction but coalesces into relatively few major conduits that are continuous through the entire aquifer. Except in mixing or redox zones, karst porosity rarely occurs as isolated voids. Branching conduit patterns are the most common. Conduits form only where the setting is favorable for certain flow paths to gain discharge at the expense of their neighbors, e.g. by development of sinkholes. Dissolution labyrinths, in which every accessible opening is enlarged to comparable amounts, form in several settings: (a) within short distances of flow from where aggressive water first enters a soluble rock, or where mixing or redox reactions produce local zones of undersaturated water within a karst aquifer; (b) in areas of steep hydraulic gradient, where Q/L and i/L are large. Dissolution porosity diminishes greatly at depths beneath the local base level because of the very strong influence of fissure widths on resistance to flow. Use of these concepts can aid in the prediction of porosity distribution and geometry. However, it must be recognized that relict karst porosity can also occur where conditions favorable to its origin are no longer present. Also, details of geologic structure must be

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.8 8 considered. These ideas are developed further in the literature. The concepts described here have not changed significantly since their first brief publication in 1981, although they have been explored at greater length since by more advanced geochemical and digital models (see Palmer, 1991 and 1995). The model described in Fig. 3 was generalized from the behavior of individual conduits. However, subsequent digital modeling of multiple-conduit networks has validated these concepts. References Back W., Hanshaw B. and Van Driel J.N. 1984. Role of groundwater in shaping the eastern coastline of the Yucatan Peninsula, Mexico. In: LaFleur, R. G. (Ed.): Groundwater as a geomorphic agent, Boston, Mass..Alien and Unwin, Inc., 281-293. Berner R.A. and Morse J.W. 1974. Dissolution kinetics of calcium carbonate in sea water, IV: Theory of calcite dissolution. American Journal of Science 274, 108-134. Buhmann D. and Dreybrodt W. 1985a. The kinetics of calcite solution and precipitation in geologically relevant situations of karst areas. 1: Open system. Chemical Geology 48, 189-211. Buhmann D. and Dreybrodt W. 1985b, The kinetics of calcite solution and precipitation in geologically relevant situations of karst areas. 2: Closed system. Chemical Geology 53, 109-124. Clemens T, Hckinhaus D., Sauter M., Liedl R. and Teutsch G. 1996. A combined continuum and discrete network reactive transport model for the simulation of karst development. In: Calibration and reliability in groundwater modeling. Proceedings of the ModelCARE 96 Conference held at Golden, Co., Sept. 1996: UAGS Publ. No. 237. Curl R.L. 1968. Solution kinetics of calcite. Proceedings of 4th International Congress of Speleology, Ljubljana, Slovenia, 61-66. Dreybrodt W. 1988. Processes in karst systems: physics, chemistry and geology: Berlin, Germany, Springer-Verlag, 288 p. Dreybrodt W. 1990. The role of dissolution kinetics in the development of karst aquifers in limestone: a model simulation of karst evolution: Journal of Geology 98 (5), 639-655. Dreybrodt W. 1996. Principles of early development of karst conduits under natural and man-made conditions revealed by mathematical analysis of numerical models: Water Resources Research 32, 2923-2935. Ford D.C. 1971. Geologic structure and a new explanation of limestone cavern genesis: Transactions of the Cave Research Group of Great Britain 13 (2), 81-94. Ford D.C., and Ewers R.O. 1978. The development of limestone cave systems in the dimensions of length and depth: Canadi an Journal of Earth Sciences 15, 1783-1798. Groves C.G., and Howard A.D. 1994a. Minimum hydrochemical conditions allowing limestone cave development. Water Resources Research 30 (3), 607-616. Groves C.G., and Howard A.D. 1994b. Early development of karst systems, 1. Preferential flow path enlargement under laminar flow. Water Resources Research 30 (10), 2837-2846. Hanna R.B., and Rajaram H. 1998. Influence of aperture variability on dissolutional growth of fissures in karst formations. Water Resources Research 34 (11), 2843-2853. Milanovic P.T. 1981. Karst hydrogeology. Littleton, Colorado, Water Resources Publications, 434 p. Moneymaker B.G. 1941. Subriver solution cavities in the Tennessee Valley. Journal of Geology 49, 74-86. Mylroie J. E., and Carew J.L. 1990. The flank margin model for dissolution cave development in carbonate platforms: Earth Surface Processes and Landforms 15, 413-424. Palmer A.N. 1969. A hydrologic study of the Indiana karst: Ph.D. thesis, Indiana Univ., Bloomington, Ind., 181 p. Palmer A.N. 1981. Hydrochemical controls in the origin of limestone caves. 8th International Speleological Congress, Proceedings, Bowling Green, Kentucky, p. 120-122. Palmer A.N. 1984. Recent trends in karst geomorphology. Journal of Geological Education 32, p. 247-253. Palmer A.N. 1988. Solutional enlargement of openings in the vicinity of hydraulic structures in karst regions. Association of Ground Water Scientists and Engineers, 2nd Conference on Environmental Problems in Karst Terranes and their Solutions, Proceedings, p. 3-15. Palmer A.N. 1991. Origin and morphology of limestone caves. Geological Society of America Bulletin, 103, 1-21. Palmer A.N. 1995. Geochemical models for the origin of macroscopic solution porosity in

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A.N.Palmer / Speleogenesis and Evolution of Karst Aquifers 1, January 2003, p.9 9 carbonate rocks. In: D.A. Budd, Harris, P.M. and Sailer, A. (Eds.), Unconformities in carbonate strata: their recognition and the significance of associated porosity. American Association of Petroleum Geologists, Memoir 63, p. 77-101. Plummer L.N., and Wigley T.M.L. 1976. The dissolution of calcite in CO2-saturated solutions at 25 C and 1 atmosphere total pressure. Geochimica et Cosmochimica Acta 40, p. 191202. Plummer L.N., Wigley T.M.L. and Parkhurst D.L. 1978. The kinetics of calcite dissolution in CO2water systems at 5 to 60 C and 0.0 to 1.0 atm CO2. American Journal of Science 278, p. 179216. Rauch H.W, and White W.B. 1977. Dissolution kinetics of carbonate rocks. 1. Effects of lithology on dissolution rate. Water Resources Research 13, p. 381-394. White W.B. 1977. Role of solution kinetics in the development of karst aquife rs. In: Tolson J.S. and Doyle F.L. (Eds.), Karst hydrogeology. International Association of Hydrogeologists, 12th Memoirs, p. 503-517. Worthington S.R.H., 1991. Karst hydrogeology of the Canadian Rocky Mountains. Ph.D. dissertation, McMaster University, Hamilton, Ont, 227 p.