Hydrogeochemistry of Dripwaters in Selected Caves of Moravian Karst Ph D Dissertation

MASARYK UNIVERSITY FACULTY OF SCIENCE DEPARTMENT OF GEOLOGICAL SCIENCES

Pavel. . Pracný

SUPERVISOR: DOC. ING. JIŘÍ FAIMON, DR. BRNO 2017 BIBLIOGRAPHIC ENTRY

Author Mgr. Pavel Pracný Faculty of Science, Masaryk University Department of Geological Sciences

Title of Dissertation Hydrogeochemistry of Dripwaters in Selected Caves of Moravian Karst

Degree Programme Geology

Field of Study Geological Sciences

Supervisor doc. Ing. Jiří Faimon, Dr. Faculty of Science, Masaryk University, Department of Geological Sciences Faculty of Science, Palacký University, Department of Geology Academic Year 2016/2017

Number of Pages 53+65

Keywords Cave dripwater; Moravian Karst; Anomalous drip; Car- bon dioxide; Mg/Ca ratio; Kinetic modeling; Limestone dissolution; Degassing; Mg-calcite BIBLIOGRAFICKÝ ZÁZNAM

Autor Mgr. Pavel Pracný Přírodovědecká fakulta, Masarykova univerzita Ústav geologických věd

Název práce Hydrogeochemie skapových vod ve vybraných jeskyních Moravského krasu

Studijní program Geologie

Studijní obor Geologické vědy

Školitel doc. Ing. Jiří Faimon, Dr. Přírodovědecká fakulta, Masarykova univerzita, Ústav geologických věd Přírodovědecká fakulta, Univerzita Palackého, Katedra geologie Akademický rok 2016/2017

Počet stran 53+65

Klíčová slova Skapové vody; Moravský kras; Anomální skap; Oxid uhličitý; Poměr Mg/Ca; Kinetické modelování; Rozpouštění vápenců; Odplyňování; Mg-kalcit ABSTRACT

Karst dripwaters are an important factor of speleothem formation. These cave precipitates provide various proxy data (e.g. stable isotopes, minor and trace elements or grow laminae) about paleoenvironment. To better understand the interrelationship between proxies and environment, an investigation of recent karst processes is important. A dripwater hydrogeochemistry and cave PCO2 were studied in the dry part of Punkva Caves (Moravian Karst). The sampling was conducted twice per month from February 2012 to March 2013. Additional dripwater samples for stable isotopes analyses were collected in April and November 2014. An anomalous drip was identified showing hydrogeochemical properties significantly different from other regular drips in the cave system as well as other caves in Moravian Karst. The anomalous drip showed a low SIcalcite ~ 0.14±0.11 (standard deviation), low specific conductivity 297±22.2 μS cm−1 and enhanced values of δ13C (−7.85 to −8.35‰ VPDB), Mg/Ca × 1000 ratio (45.7±3.3) and Sr/Ca × 1000 ratio (0.65±0.06). In contrast, the regular drips showed saturation SIcalcite in range from 0.83 to 1.07, high specific conductivity (604±32 μS cm−1) and lower Mg/Ca × 1000 (17.0±1.4) and Sr/Ca × 1000 (0.31±0.02) ratios as well as lower δ13C values (−10.34 to −10.94‰). The data analysis supports conclusion that the anomalous drip properties were a consequence of a prior calcite precipitation or/and water mixing. This idea is supported by the position of the drip on a crevice edge.

The partial pressure of CO2 measured in cave air, PCO2(air), was in range from 10−3.31 to 10−2.49 (0.06–0.32 vol%). These values were compared to CO2 partial pressures calculated from dripwater hydrogeochemistry as a partial pressure of CO2 corresponding to aqueous carbonates, PCO2(W) (10−2.91 to 10−2.35 or

0.12–0.45 vol%), and hypothetical CO2 partial pressure participating on the initial dripwater formation, PCO2(H) (10−1.77 to 10−1.49 or 1.7–3.2 vol%). Both the

PCO2(air) and PCO2(W) showed clear seasonal variations with maxima in summer and minima in winter. It seems that the cave air CO2 had been controlled by cave ventilation modes: the higher PCO2(air) were a result of a downward airflow mode during the period of active ventilation with increased influx of CO2 from epikarst and vadose zone. In contrast, the PCO2(H) was very stable without any significant seasonality indicating independence on seasonally changing surface conditions. It could mean that the source of CO2 is deployed deeper in karst profile under the soil. The anomalous drip represented an exception with lower and varying PCO2(W) and PCO2(H) close to PCO2(air) indicating prior CO2 degassing and calcite precipitation. A geochemical model of CO2 degassing shows that the regular dripwaters data are plotted along a degassing line with slope ~ −1 pointing to a unique value of PCO2(H) regardless of season. In addition, it shows that the anomalous drip data are much more scattered and estimated values of PCO2(H) are most probably incorrect due to previous calcite precipitation changing dripwater hydrochemical properties. The possibility of dripwater conversion into solution aggressive with respect to calcite due to anthropogenic CO2 influx into cave were studied in Výpustek Cave. The model showed that it is possible to reach sufficient cave

CO2 concentrations during longer events with enhanced attendance (500 people). Ordinary guided tours (50 people, ca. 0.5 h) seem to be of inconsequential effect. A dynamic model of the Mg/Ca ratio theoretical evolution during limestone dissolution in epikarst (T = 10 °C, logPCO2 = –1.5) was designed. The limestone was modeled as a dolomite and Mg-calcite mix with various content ratios. Two distinct stages of the dissolution were observed: (a) an initial stage with stoichiometric release of Ca and Mg (congruent dissolution) and (b) an advanced stage beginning when the solution reached calcite saturation, characterized by a continuous release of Mg and concurrent Ca decrease due to calcite precipitation (incongruent dissolution). The overall Mg/Ca ratio evolution, represented by shape of dissolution reaction paths, is determined by the Mg- calcite composition and the ratio of Mg-calcite and dolomite (D/C). The dynamics of Mg-calcite dissolution dominates for all ratios under D/C = 1, when the reaction paths divert from pure Mg-calcite paths. A minor factor influencing the reaction path of Mg/Ca evolution was identified in the ratio of limestone surface to water volume ({L}/V). However, the {L}/V ratio controls overall interaction dynamics. In epikarst, the dissolution dynamics is given by conditions of a system open to gaseous CO2, leading to enhanced epikarst dissolution. The ratio is probably higher deeper in vadose zone, but the dissolution is limited because the system is closed to CO2. The actual Mg/Ca ratio in dripwater depends on water residence time (i.e., water-rock interaction time) which controls how far the dissolution proceeds along the reaction path. Dripwaters from Punkva Caves and other sites over the world were compared with the model results. Most of the waters showed Mg/Ca ratios similar to reaction paths for Mg-calcite and low-dolomite limestone, whereas dripwaters from dolostone were similar to evolution during dissolution of pure dolomite. Both the difference in Mg/Ca ratios of the anomalous drip compared to regular drips and the Mg/Ca evolution model show the importance of water flow paths in dripwater formation. However, the water flow paths could change both temporally and spatially with the evolution of karst system, independently on climate conditions. In addition, the Mg-calcite composition and dissolution dynamics seem to play substantial role in dripwater Mg/Ca ratio evolution. Therefore, it is important to take these factors into consideration in paleoenvironmental studies of karst proxies. ABSTRAKT

Krasové vody představují důležitý činitel při vzniku speleotém. Tyto jes- kynní sedimenty jsou významným zdrojem proxy dat (např. stabilní isotopy, stopové prvky, přírůstkové linie) o environmentálních podmínkách v době svého vzniku. Pro lepší pochopení vazeb mezi proxy daty a klimatem je důležité studium recentních krasových procesů. Hydrogeochemické vlastnosti skapo- vých vod a PCO2 v Punkevních jeskyních (Moravský kras) byly studovány dva- krát měsíčně od února 2012 do března 2013. Dodatečné vzorky pro izotopické analýzy byly odebrány v dubnu a listopadu 2014. V Punkevních jeskyních byl identifikován anomální skap s výrazně odliš- nými vlastnostmi ve srovnání s běžnými skapy v této i jiných jeskyních Morav- ského krasu. Anomální skap vykazuje nízké přesycení ke kalcitu SIcalcite ~ 0,14±0,11 (směrodatná odchylka) a nízkou specifickou vodivost 297±22,2 μS cm−1 a zvýšené hodnoty δ13C (−7,85 až −8,35 ‰ VPDB), poměru Mg/Ca × 1000 (45,7±3,3) a poměru Sr/Ca × 1000 (0,65±0,06). Běžné skapy vykazují hodnoty indexu nasycení SIcalcite rozmezí od 0,83 do 1,07, vysokou specifickou vodivost (604±32 μS cm−1) a nižší hodnoty δ13C (−10,34 to −10,94 ‰) a poměrů Mg/Ca × 1000 (17,0±1,4) a Sr/Ca × 1000 (0,31±0,02). Data naznačují, že vlastnosti ano- málního skapu jsou důsledkem předběžného srážení kalcitu a/nebo mixování vod v nadloží, což podporuje také pozice skapu na hraně komína.

Praciální tlak CO2 v jeskynním vzduchu, PCO2(air), se pohyboval v rozmezí od 10−3,31 do 10−2,49 (0,06–0,32 obj. %). Tyto hodnoty byly porovnány s parciál- ními tlaky vypočítanými na základě hydrogeochemických vlastností skapových vod. Jsou to parciální tlak odpovídající obsahu karbonátů v roztoku, PCO2(W) (10−2,91 až 10−2,35, tj. 0,12–0,45 obj. %), a hypotetický parciální tlak podílející se na formování skapové vody, PCO2(H) (10−1,77 až 10−1,49, tj. 1,7–3,2 obj. %). Jak

PCO2(air) tak i PCO2(W) vykazují jasnou sezónnost s maximy v létě a minimy v zimě. Toto chování je patrně dáno jeskynní ventilací: vyšší PCO2(air) je důsled- kem režimu sestupného proudění v období aktivní jeskynní ventilace se zvýše- ným přínosem CO2 z epikrasu a vadózní zóny. Naopak hodnoty PCO2(H) byly velmi stálé bez výraznější sezónnosti, což naznačuje významnou nezávislost na povrchových podmínkách a původ CO2 formujícího skapové vody ne v půdě, ale hlouběji v krasovém profilu. Anomální skap představuje významnou výjimku s výrazně nižším a variabilnějším PCO2(W) a hodnotami PCO2(H) blízko PCO2(air) což indikuje předběžné odplyňování CO2 a srážení kalcitu. Vývoj skapových vod při odplyňování CO2 a následném srážení kalcitu je ilustrován na modelu. Ten ukazuje, že data odpovídající běžným skapům leží podél linie odplyňování se směrnicí ~ −1 směřující k unikátní hodnotě PCO2(H) nezávisle na ročním období. Dále je vidět, že data z anomálního skapu jsou mnohem rozptýlenější a odhad hodnoty PCO2(H) z linie odplyňování bude nesprávně ukazovat hodnotu danou předchozím srážením kalcitu, které změnilo hydrogeochemické vlastnosti vody. Pří větších kulturních akcích v jeskyni Výpůstek byl posuzován také vliv antropogenního CO2 na krasové vody a možnost jejich konverze na vody agre- sivní ke kalcitu. Modelování ukazuje, že při delších akcích se zvýšenou návštěv- ností (500 lidí) je možné dosáhnout dostatečných koncentrací CO2. Běžné pro- hlídky (50 lidí, cca 0,5 h) se zdají být bez významnějšího účinku. Byl sestaven model teoretického vývoje poměru Mg/Ca v průběhu roz- pouštění vápence v epikrasu (T = 10 °C, logPCO2 = –1,5). Vápenec byl simulován jako mix různých poměrů dolomitu a hořečnatého kalcitu. Byly pozorovány dvě zřetelné fáze rozpouštění: (a) počáteční období stechiometrického uvolňování Ca a Mg (kongruentní rozpouštění) a (b) pokročilá fáze začínající po dosažení nasycení kalcitem, odpovídající nekongruentnímu rozpouštění s nárůstem Mg a souběžným poklesem Ca díky srážení kalcitu. Celkový vývoj poměru Mg/Ca reprezentovaný tvarem reakční cesty rozpouštění je dán složením Mg-kalcitu a poměrem Mg-kalcitu a dolomitu (D/C). Při poměrech D/C menších než 1 domi- nuje dynamika rozpouštění Mg-kalcitu a reakční cesty pro vápenec se překrý- vají s cestami pro čistý Mg-kalcit. Další faktor v omezené míře určující tvar reakční cesty je poměr povrchu vápence ku objemu vody ({L}/V). Poměr {L}/V nicméně určuje celkovou dynamiku interakcí. V epikrasu probíhá rozpouštění za podmínek systému otevřeného vůči CO2, což vede k intenzivnějšímu roz- pouštění. Navzdory tomu, že v puklinách hlouběji ve vadózní zóně budou hodnoty {L}/V vyšší, rozpouštění je limitováno uzavřením systému vůči CO2. Oka- mžitý poměr Mg/Ca ve vodě je dán tím, jak daleko podél reakční cesty postou- pilo rozpouštění, což je určeno dobou zadržení vody (tj. doba interakce voda- hornina). Modelové výsledky byly srovnány se skapovými vodami z Punkevních jeskyní i dalších lokalit ve světě. Většina skapových vod vykazovala poměry Mg/Ca odpovídající reakčním cestám Mg-kalcitu s malou příměsí dolomitu, za- tímco vody z dolomitických hornin se podobaly vývoji rozpouštění čistého dolomitu. Jak rozdíl v poměrech Mg/Ca mezi anomálním a běžnými skapy, tak model vývoje poměrů Mg/Ca ukazují na důležitost cest proudění vody v krasu pro vznik skapových vod. Nicméně cesty proudění se mohou měnit v čase i prostoru tak, jak se vyvíjí krasový systém, a to nezávisle na klimatických podmínkách. Navíc se zdá, že složení Mg-kalcitu a dynamika rozpouštění hrají podstatnou roli ve vývoji Mg/Ca poměrů ve vápencích. A proto je při paleoenvironmentál- ních studiích krasových proxy dat důležité brát tyto vlivy v potaz. © Pavel Pracný, Masaryk University, 2017 © John Wiley & Sons, Ltd., 2015 © Springer-Verlag Berlin Heidelberg, 2015 © Springer Science+Business Media Dordrecht, 2017 © 2016 Elsevier GmbH., 2017 DECLARATION

I declare that this PhD dissertation is an original report of my reseach and was composed by myself. I confirm that the work submitted is my own, except where work which has formed part of jointly-authored publications has been included. I confirm that appropriate credit has been given within this dissertation where reference has been made to the work of others. I agree that my dissertation may be available in the library of Masaryk Uni- versity.

Brno, 12. 5. 2017 Pavel Pracný TABLE OF CONTENTS

1 Foreword ...... 13 2 Review of dripwater formation and processes ...... 17 2.1 Karst Hydrogeology ...... 17

2.1.1 Dripwater hydrology ...... 19

2.1.2 Residence times ...... 21

2.2 Hydrogeochemistry of dripwaters ...... 21

2.2.1 Carbonate system ...... 21

2.2.2 Dissolution kinetics of carbonate minerals ...... 24

2.2.3 Karst water properties and evolution ...... 25

2.3 Paleoenvironmental reconstructions ...... 28

3 Results and discussion ...... 29 3.1 Hydrogeochemistry of dripwaters in Moravian Karst ...... 29

3.2 Cave and epikarstic PCO2 ...... 32

3.3 Modeling limestone dissolution and Mg/Ca evolution in epikarst ...... 35

3.4 Anthropogenic CO2 influence on dripwaters and speleothem corrosion 38

4 Conclusions ...... 41 5 References ...... 44 Appendix 1 ...... 54 Appendix 2 ...... 70 Appendix 3 ...... 83 Appendix 4 ...... 105

12 1 FOREWORD

Limestone karst areas form large portion of continental crust surface. Be- sides being an important reservoir in biogeochemical cycle of carbon, they provide various benefits for humankind. Many islands are formed of limestone rocks, karst aquifers constitute major fresh water source and karst regions provide plentiful recreational opportunities – from marvelous beaches of southeast Asia, complex cave system of Mammoth Cave in North America to picturesque landscapes of Mediterranean, to name a few. With increasing awareness of climate changes over the last decades, the importance of various paleoclimatic data archives rapidly increased. Karst systems offer important high-resolution terrestrial archives via both the surface (tufas) and the underground (speleothems) secondary karst sediments. Espe- cially speleothems, conveniently sheltered in mostly inaccessible caves, preserve detailed record of environmental conditions at the time of their formation. In order to recover this information, we have to understand mechanisms of speleothem growth as well as the processes, which formed karst waters precipitating speleothems. Although seemingly elementary in its composition, the carbonate karst environment exhibits intriguing levels of complexity. Phenomena considered in holistic approach to karst include biological activity (esp. CO2 production in soil), processes on water-atmosphere boundary (e.g., CO2 degassing in cave), water-rock interaction during limestone dissolution, petrological and miner- alogical characteristics of bedrock, hydrogeological properties of vadose zone among many other aspects. Such intricacy allows for many climatic dependent variables to be reflected in speleothems, while, on the other hand, a careful separation of paleoenvironmental signals is required. Arguably, the most utilized paleoenvironmental proxies are stable isotopes, albeit other parameters, esp. trace elements or growth laminae, provide valuable support for reconstructions. Despite utilization of trace elements in reconstructions, the exact mechanisms of their transfer into the original karst water are not very well understood.

13 Additionally, speleothem destruction repeatedly emerges as a lively discussed topic among environmentalists of all kinds (Baker & Genty 1998; Auler & Smart 2004; Dreybrodt et al. 2005; Martín-Peréz et al. 2012). Although previous studies found solid evidence of structural damage causing the destruction in Moravian Karst (Faimon et al. 2004), the possibility of corrosive effects of dripwater cannot still be ruled out. The systematic long-term research in the Punkva Caves might produce decisive insights. This thesis describes results of a research focused on hydrogeochemistry of dripwaters in Moravian Karst. It was particularly aimed to (1) improve our understanding of processes determining the dripwater parameters, possibly applicable as paleoenvironmental proxies (especially minor (Mg2+) and trace (Sr2+) cations) and to (2) find and characterize possibly corrosive dripwaters. The basis for all considerations is a dataset collected in the Punkva Caves in Moravian Karst from February 2012 to March 2013. Although previous works described dripwater properties in some Moravian Karst caves (Faimon et al. 2004; Faimon et al. 2012), our dataset represents an unprecedentedly detailed and thorough study of one site. From hydrogeochemical properties, the relatively low-saturated anomalous drip in the Punkva Caves was identified. Based on differences between the anomalous and the regular drips, we could infer possible implications for paleoenvironmental reconstructions regarding Mg/Ca ratios. Moreover, the detailed dataset allowed a comparison of seasonal variations of cave and dripwater CO2 in addition to reconstruction of a hypothetical epikarst PCO2 showing low seasonality. Finally, we devised a kinetic model of limestone dissolution in epikarst, which provided valuable insights into the Mg/Ca evolution during incongruent dissolution of carbonate minerals. This PhD dissertation presents (a) an introductory review of karst hydrogeology and hydrogeochemistry with regards to cave dripwater and (b) an overview of the journals’ papers where I have participated as the main author. Other publications related to this topic with my minor contribution or conference presentations are not included in the thesis, but are referenced in the review part.

14 The initial impulse which started the whole research endeavor occurred in early 2011, when Ing. Luděk Kabelka, an analytical chemist whose curiosity was fortuitously triggered by a newspaper article about speleothem destruction, contacted Dr. Jiří Faimon, an associate professor at Department of Geo- logical Sciences, Masaryk University, to discuss some insights into the problem. The result was a serendipitous scientific disagreement, which was later drafted into a dissertation topic and offered to me. Dr. Jiří Faimon became my PhD advisor and I am deeply thankful for this opportunity – he taught me many principles of scientific thought, work and writing, spent numerous hours discussing (and on his part mostly explaining) various geochemical topics and helped me to tame my somewhat chaotic and spontaneous nature in order to become more organized and precise – to name a few things I learned from him. In addition, without assistance from Ing. Luděk Kabelka from GEOtest s.r.o., who participated in large portion of the field campaign, provided essential analyses in hydrogeochemical laboratories of GEOtest s.r.o. and negotiated generous donation, the planned intensity of sampling would be impossible to achieve and process. I would like to thank my colleagues and students from Department of Ge- ological Sciences, who either helped me directly with my research or with whom I had the opportunity to cooperate on other research projects, advancing my skills. Namely to Dr. Marek Lang, Ms. Miluška Hradská, Mr. Tomáš Praj, Ms. Klára Blažková, Mrs. Radka Bodláková, Mr. Erik Rzepiel, Dr. Dalibor Všianský and Dr. Josef Zeman, even though there are many more involved. I am also very thankful to Mr. Pavel Kadlec who helped me with the laboratory work and to prof. Ondra Sracek from Palacký University, who participated in isotopic research of dripwaters. The bulk of field work was conducted in the Punkva Caves and would not be possible without kind cooperation of the Cave Administration of the Czech Republic, namely Mr. Hynek Pavelka and Mr. Jiří Hebelka, who also provided

15 climatic data from administration’s meteo-station, and additional cooperation from the Administration of the Protected Landscape Area Moravský kras. At last but not least, I am wholeheartedly thankful to my whole family for lasting support through my study years and their confidence in happy ending. And finally, I would like to state with profound gratitude that I deeply admire my wife Simona for her immense kindness, patience and encouragement in face of my PhD adventures. Thank you!

16 2 REVIEW OF DRIPWATER FORMATION AND PROCESSES

2.1 KARST HYDROGEOLOGY The most important karst rock is limestone, a rock composed of carbonate minerals, mostly calcite (CaCO3) and dolomite (CaMg(CO3)2) and to a lesser extent of aragonite (CaCO3). A significant component might be Mg-calcite which contains up to 20% of magnesium, although it is less stable than calcite and tends to re-crystallize during diagenesis (Mackenzie et al. 1983; Ford & Williams 2007). Geomorphologically, limestone karst in temperate climate regions is composed of wide plateaus intersected by deep valleys with various surface features formed by water (sinkholes, polje, limestone pavement etc.). One of the primary sources of water flowing through karst profile is infiltrated precipitation. The infiltrating water moves (a) by a slow matrix flow through interstitial spaces between soil grains and/or (b) via preferential flow along the macropores, e.g., animal dens/holes, root residue, or mud cracks (Kogovšek & Šebela 2004). Although soil thickness significantly increases water storage capacity, water percolation into caves is commonly observed even in regions without any soil coverage (Klimchouk 2004). Another substantial source of karst water is infiltrating water flowing from non-karstic terrains into carbonate rocks (Ford & Williams 2007). Due to the high solubility of limestone, any surface water in karst areas rapidly disappears under surface shaping characteristic underground cave phenomena. Therefore, the conditions in karstic aquifers are largely determined by the geological and lithological properties of given karst region. Per some authors (e.g., Ternan 1972; Trček 2003; Smart & Worthington 2004) the underground environment can be divided into two hydrogeological structures: A. A network of interconnected karst conduits with high overall permeability, fracture porosity and quick preferential flow along these conduits. The water residence time is short. Therefore, the structure serves as a drainage of less permeable rock. B. Rock massive with low permeability, i.e., limestone bedrock with primary porosity and secondary tectonic permeability. In spite of the low

17 permeability, the flow velocity is significantly lower and the residence time is usually very long. Another division of karst underground (Palmer 2006; Ford & Williams 2007) is a zonation based on the hydrogeological properties and describes it in terms of three major zones: 1. The uppermost unsaturated zone of strongly weathered rock underneath soil – epikarst. 2. Lower unsaturated zone (vadose zone) with large underground caverns. The lowest part of vadose zone can be periodically saturated (e.g., during floods) forming an epiphreatic (sub)zone. 3. Saturated zone (phreatic zone), the zone of permanent water saturation under the water table where water flows primarily horizontally via conduit systems. As the vertical permeability in epikarst zone rapidly diminishes on the boundary with vadose zone, excessive water is stored in pores, fractures and joints forming perched aquifers. From the aquifers water flows along fissures downward into the vadose zone and is eventually released into the cave as dripwater (Perrin et al. 2003a; Ford & Williams 2007; Williams 2008; Jones 2013). Therefore, epikarst is of major hydrologic importance. Its overall water capacity is given by i. epikarst zone thickness, ii. porosity, iii. water inflow and outflow balance. Whereas the thickness and porosity determine the sum of space available for the water, the ability to retain water is given by the water inflow/outflow balance. The outflow is in addition to hydrostatic pressure determined by the vertical hydraulic conductivity of the vadose zone, which is dependent on fracture porosity and can be very variable. Therefore, the epikarst water capacity might significantly vary even on small scale.

18 2.1.1 Dripwater hydrology Rudimental dripwater hydrology studies are based on discharge response to precipitation (Smart & Friederich 1987; Baker et al. 1997; Genty & Deflan- dre 1998; McDonald & Drysdale 2007), but the research focus is gradually shifted towards analyses of relation between precipitation and geochemical properties of dripwater (Perrin et al. 2003a; Musgrove & Banner 2004; Cruz et al. 2005; Schwarz et al. 2009; Riechelmann et al. 2011; Faimon et al. 2012; Kamas et al. 2015). A few modes of water flow were distinguished in karst unsaturated zone and are consistent with the division described in previous chapter: water flows (A) via a system of interconnected fissures, joints and conduits (conduit flow) and/or (B) seeps through the bedrock matrix (matrix flow). These flow types may differ in residence times and hydrogeochemical properties of resulting dripwaters. A classification of karst underground water flow was based on the discharge and discharge variability (Friederich and Smart 1982), which was latter improved by Baker et al. (1997) and is commonly utilized (e.g., Spötl et al. 2005; Baldini et al. 2006a; McDonald & Drysdale 2007; Hartland et al. 2012; Fairchild & Baker 2012; Pracný et al. 2016). Dripwaters show various responses to atmospheric precipitation leading to linear, non-linear, or even no direct correlation (e.g., McDonald et al. 2007; Pronk et al. 2009; Riechelmann et al. 2011; Faimon et al. 2016). The drips with stable discharge during dry periods without atmospheric precipitation might indicate a water source in slowly percolating infiltration or very large perched aquifer. On the contrary, the drips with rapid response to precipitation may have faster connection with the surface. Based on a long-term monitoring Tooth and Fairchild (2003) described four types of precipitation responses: 1. Rapid response without time-lag, after which the discharge slowly decreases. 2. Rapid response with associated time-lag, after which the discharge slowly decreases. 3. Intermittent response, when a particular threshold of water input must be exceeded, before the drip discharge responses, occasionally even by decreasing.

19 4. No response to particular precipitation; the changes in discharge intensity are not correlated with precipitation. The non-linear responses to precipitation are explained by an aquifer with an overflow – it maintains a constant water head and in case of intensive influx the excessive water is drained via otherwise dry outlet (and could feed a seasonal drip). Tracer studies (Kogovšek & Šebela 2004; Goldscheider et al. 2008; Pronk et al. 2009; Kogovsek & Petric 2014) showed that the tracer concentration in dripwater diminished exponentially suggesting a simple linear flow. Neverthe- less, in case of some drips, the concentration increased again after another rainfall event indicating existence of periodically drained primary collectors. Be- sides, significant lateral dispersion during the wet periods was observed in some dripwater tracer experiments (Bottrell & Atkinson 1992). Another process considerably affecting tracer concentrations is a mixing along the water flow path. It seems that hydraulically the epikarst is composed of numerous flow paths that can be in limited contact and allow partial water mixing (Perrin et al. 2003a; Perrin et al. 2007). The observed lateral dispersion proves horizontal spread of water during wet periods, indicating overflow from the primary aquifer into adjacent free spaces enabling increased mixing and feeding of seasonal drips. Conversely, the epikarst aquifer becomes fractionated into separate reservoirs feeding individual drips (or groups of drips) during dry periods. An insightful study by Genty and Deflandre (1998) showed that the volume of drop in dripwater was ~0.15 mL in 94% cases. Nevertheless, the volume varies under extreme discharge. At higher discharges (over ~50 drops/min), the drop volume decreases, possibly due to limitations of fluid physical properties (density, surface tension, water pressure connected to the flow) or development of side-drips. On the contrary, the volume increases at lower discharges (under 1 drop/min). In addition, Fairchild et al. (2006b) report that drop volume for drips with various discharges was in range from 0.146 to 0.154 mL.

20 2.1.2 Residence times An important mechanism participating in dripwater hydrogeochemistry could be the piston flow effect. Piston flow is a water transport mechanism occurring when the infiltrating water pushes the water retained in flow paths since previous precipitation event towards the outlets. In this case, water discharge in cave increases almost instantly after major precipitation events. In addition, the water pushed by the piston flow is often flushed out from parts of the reservoir with longer residence times as indicated by the increased trace element content and isotope ratios (e.g., Aquilina et al. 2006, Tooth & Fairchild 2003, Emblanch et al. 2003). The overall residence times of water in vadose zone are quite variable and ranging from days (Bottrell & Atkinson 1992; Genty & Deflandre 1998; Perrin et al. 2003b; Kogovšek & Šebela 2004; Kamas et al. 2015; Faimon et al. 2016) to months and even years (Spötl et al. 2005; Aquilina et al. 2006; Kluge et al. 2010; Kogovšek & Petric 2014) as a result of very complicated water flow paths or extremely slow matrix flow. During the wet periods, a continuous flow through all types of hydraulically connected fissures is enabled. Therefore, a relatively fast transport of solution from surface into cave occurs. In contrast, during the dry periods, only a portion of flow paths is active and the major part of infiltrated precipitation is stored in less permeable parts of vadose zone. This water might be flushed out after intense precipitation events (Kogovšek & Šebela 2004; Kogovšek & Petric 2014).

2.2 HYDROGEOCHEMISTRY OF DRIPWATERS The hydrogeochemical properties of dripwaters are given by processes occurring along the water flow path from karst surface into a cave and are ulti- mately embedded into the speleothem structure and composition.

2.2.1 Carbonate system Biogeochemically the most important compound of carbon is carbon dioxide CO2. The solubility of gaseous CO2 in water decreases with rising temperature. The equilibrium between CO2 in atmosphere above the solution and CO2 in the solution is expressed by Henry’s constant (Sander 2015):

21 c Hcp = CO2(푎푞) (1) PCO2

Where cCO2(aq) is the concentration of CO2 in the aqueous phase and PCO2 is the partial pressure of CO2 in the atmosphere. Oftentimes, it is defined as a dimensionless parameter given by the ratio between the aqueous phase concentration (cCO2(aq)) of a specie and its gas-phase concentration (cCO2(g)) c Hcc = CO2(푎푞) (2) cCO2(푔) The values for standard conditions are presented in Table 1. Dissolved carbon dioxide reacts with water to form carbonic acid (Stumm & Morgan 2012):

CO2(푎푞) + H2O = H2CO3 (3) with the equilibrium constant

푎퐻2퐶푂3(푎푞) K0 = (4) 푎퐶푂2(푎푞)

Although unhydrated CO2(aq) is much more abundant than H2CO3, a con- vention was adopted to include the hydration in the overall dissolution in order to facilitate calculations as the hydration step is effectively instantaneous. The two carbonate species are summarily expressed as H2CO3* and the overall reaction of CO2 dissolution becomes ∗ CO2(푔) + H2O = H2CO3 (5) Carbonic acid further dissociates into two species

* - + H2CO3 = HCO3 + H (6) and − 2− + HCO3 = CO3 + H (7)

with equilibrium constants K1 and K2 which are presented in Table 1 (and

K1 is in fact composite constant including the CO2(aq) hydration). The distribution of carbonate species is determined by pH value. Under acidic conditions of pH < 4 the system contains effectively only carbonic acid. As the pH increases, the acid dissociates into bicarbonate reaching maximum at pH = 8.3. With further pH increase, the bicarbonate ion dissociates into the carbonate ion, which effectively dominates the carbonate system when pH is above 12. The calcium carbonate (calcite or aragonite) dissolution can be described by simple dissolution equation:

22 2− 2+ CaCO3(s) = CO3 + Ca (8) with solubility products for calcite (Kc) and aragonite (Ka)

K = 푎 2+ 푎 2− (9) 푐/푎 Ca CO3 In karst environment studies, the pinnacle of interest lies in interactions of carbonate minerals with water and carbon dioxide. The dissolution under open system conditions can be expressed by equation − 2+ CaCO3(푠) + CO2(푔) + H2O = 2HCO3 + Ca (10) and total equilibrium constant

2 푎 − 푎 2+ HCO3 Ca KT = (11) PCO2 The value of calcite solubility product is temperature dependent and is also affected by variations in Ca activity caused by presence of aqueous complexes such as CaCO30 and CaHCO3+ (Jacobson & Langmuir 1974). Dissolution of dolomite in pure water is analogically described as 2− 2+ 2+ CaMg(CO3)2(푠) = 2CO3 + Ca + Mg (12) with solubility product 2 K푑 = 푎 2+푎 2+푎 2− (13) Ca Mg CO3

The equation for open system in contact with CO2 and dolomite is − 2+ 2+ CaMg(CO3)2(푠) + 2CO2(푔) + 2H2O = 4HCO3 + Ca + Mg (14) Dissolution of Mg-calcite with content of Ca = x and Mg = y where x + y = 1 is given by equation: 2− 2+ 2+ Ca푥Mg푦CO3(s) = CO3 + 푥Ca + 푦Mg (15) with solubility product 푥 푦 K = 푎 2+ 푎 2+ 푎 2− (16) 푀푔−푐푎푙푐푖푡푒 Ca Mg CO3 Due to various sources, genesis and composition of Mg-calcite, the published solubility products vary in a wide range from ca. 10−8.5 to 10−7.4 (Plummer & Mackenzie 1974; Mackenzie et al. 1983; Morse & Mackenzie 1990).

Solubility product Ksp of a mineral is defined by the equilibrium activities of minerals’ constituents. If the system is not in equilibrium the immediate species’ activities constitute ion activity product IAP. By comparing Ksp to IAP the reaction direction can be assumed – if IAP/Ksp > 1 the reaction will proceed towards the reactants and vice versa. This relation is useful to specify solution

23 saturation with respect to a mineral. The most common expression is the saturation index SI = log IAP/Ksp. Supersaturated solutions show positive SI values, whereas not saturated solutions show negative values.

Table 1 Equilibrium constants for the carbonate system − log T [°C] 0 5 10 15 20 25

K0 1.11 1.19 1.27 1.34 1.41 1.47

K1 6.58 6.52 6.46 6.42 6.38 6.35

K2 10.63 10.55 10.49 10.43 10.38 10.33

Kc 8.38 8.39 8.41 8.34 8.45 8.48

Ka 8.22 8.24 8.26 8.28 8.31 8.34

Kd - - - - - 17.2±2 All values from Plummer & Busenberg (1982) except dolomite (Sherman & Barak 2000)

2.2.2 Dissolution kinetics of carbonate minerals Dissolution and precipitation rates of carbonate minerals are determined by various factors and were studied under wide range of conditions with significant variations in resulting rate values (see Morse & Arvidson 2002 for review; Arvidson et al. 2003; Kaufmann & Dreybrodt 2007; Morse et al. 2007; Cubillas et al. 2005; Pokrovsky et al. 2005; Pokrovsky et al. 2009). The basic mechanisms of dissolution on mineral surface were described by Plummer et al. (1978) and expanded by Chou et al. (1989) as follows for calcite:

kc1  HCaCO  Ca 2  HCO  (17) 3  3 kc1

kc 2  HCaCO CO * Ca 2  2HCO  (18) 323  3 kc 2

kc3 CaCO 2  COCa 2 (19) 3  3 kc3 and for dolomite:

kd1  H2)CaMg(CO  Ca 2 Mg 2  2H CO  (20) 23  3 kd1

k d 2  H2)CaMg(CO CO * Ca 2 Mg 2  4H CO  (21) 3223  3 kd 2

24 k d 3 )CaMg(CO Ca 2 Mg 2  2CO 2 (22) 23  3 kd 3 The total forward and backward rates for calcite are than expressed as:

+ ∗ (23) 푅푓(푐푎푙푐푖푡푒) = 푘푐1푎H + 푘푐2푎H2CO3 + 푘푐3

푅 = 푘 푎 2+푎 − + 푘 푎 2+푎 − + 푘 푎 2+푎 2− (24) 푏(푐푎푙푐푖푡푒) −푐1 Ca HCO3 −푐2 Ca HCO3 −푐3 Ca CO3 The forward rate equations for calcite and dolomite differ only in reaction order. Whereas for calcite the reaction order n = 1, experiments show that the reaction order of dolomite dissolution with respect to aH+ is a fractional number. Busenberg and Plummer (1982) found that n = 0.5 (for 25 °C) and the value increases with increasing temperature. Definition of the dolomite backward rate is more complicated. Busenberg & Plummer (1982) determined that the

CO32− is not responsible for the observed backward reaction and it is independent on the activities of Ca2+ and Mg2+ below pH 6 and far from equilibrium, resulting in

− (25) 푅푏(푑표푙표푚푖푡푒) = 푅푓(푑표푙표푚푖푡푒) − 푅푡표푡푎푙 = 푘4푎HCO3

Where k4 is the rate constant of backward reaction. Generally speaking, the dissolution mechanisms of calcite are transferra- ble to other mono-cation carbonate minerals (e.g., magnesite and aragonite) whereas dissolution of dolomite is an example of composite carbonate mineral. The expression for dissolution rate of Mg-calcite is expanded to include magnesium activity and stoichiometry of Ca2+ and Mg2+.

2.2.3 Karst water properties and evolution Hydrochemical properties of karst water are defined by the content of substances in solution, which are incorporated especially via dissolution of minerals in contact with infiltrating water and gaseous CO2. The initial composition of infiltrating water is result of complex interactions in atmosphere. The chemical constituents dissolved in rainwater are usually connected to source oceanic water and particle/gases in atmosphere. Local sources of pollution in urban areas may participate. They are utilized as an indicator of the extent of anthropogenic pollution. Among the most important pollutants emitted by humans are both nitrogen and sulfur oxides, which are converted into the acids that are

25 a major cause of water acidity. In turn, acidic rainwater outwashes heavy metals present in atmosphere mainly from industrial sources and road transporta- tion emissions (e.g. Eriksson 1952; Carrol 1962; Gatz 1991; Paternoster et al. 2014; Vet et al. 2014). The seawater contribution to precipitation composition is ordinarily determined by comparing ion ratios to the same ratios in marine water (e.g. D'Alessandro et al. 2013). The dominant dissolved ions in carbonate karst waters are calcium and carbonates released from calcite dissolution. Other cations released by dissolution of natural calcite/limestone are Mg2+, Fe2+, Mn2+ and Sr2+. One of the determining constituents of karst water is dissolved carbon dioxide and its’ species (Atkinson 1977). The additional anions (e.g. Cl−; SO42− or NO3−) are either originating from minor minerals present in limestone and soil, from anthropogenic pollution, or are initially present in precipitated water (e.g., Perrin et al. 2003b). Formation of dripwater begins immediately after infiltration of atmospheric precipitation. Infiltrating water dissolves CO2 present in soil/epikarst atmosphere that participates on carbonate mineral dissolution (see equation

3). The source of CO2 is linked to biological activity – e.g., plant roots respiration or microbial decomposition of organic matter (Kuzyakov 2006). CO2 concentrations in soils are climatically and seasonally dependent (esp. in temperate climate); they are enhanced in warm and wet regions and seasons (Sanchez- Canete et al. 2011; Plestenjak et al. 2012). Albeit soil is widely accepted as a main source of CO2, enhanced concentrations (2–6 vol%) were measured in vadose zone of karst in Mediterranean area (Benavente et al. 2010). In karst soils of Moravian Karst, the concentrations usually reach up to 1 vol% with strong seasonal fluctuations leading to highest values in summer and lowest in winter (Faimon & Ličbinská 2010; Blecha & Faimon 2014). The acidic water in epikarstic perched aquifer dissolves limestone under conditions (esp. PCO2) which can be retrospectively estimated from dripwater composition (e.g. Faimon et al. 2012; Peyraube et al. 2013; Milanolo & Gabrovšek 2015; Pracný et al. 2016b). Eventually, the water is drained into

26 joints or fissures and flows downwards into the cave system. The water residence times of perennial drips are usually long enough for the water to be fully equilibrated with respect to calcite (Spötl et al. 2005; Kogovšek & Petric 2014). Then, the residence time usually does not affect the concentrations of major constituents of dripwater. Nevertheless, the limestone bedrock might contain numerous minerals with slower dissolution kinetics (e.g. dolomite or clay minerals) that release trace elements, which are important in paleoenvironmental analysis of speleothems. One of processes affecting hydrogeochemical properties of dripwater is prior calcite precipitation (PCP, Fairchild et al. 2000). It occurs whenever the aqueous CO2 is able to degas into ventilated spaces within the vadose zone between epikarst aquifer and drip site. During calcite precipitation, aqueous Ca concentration decreases and, thus, the trace element to calcium ratios in water increase, possibly archived in forming speleothems. PCP is believed to be promoted by drier climatic conditions (Fairchild et al. 2000, McMillan et al. 2005, Fairchild et al. 2006b, Tremaine & Froehlich 2013). Another process influencing composition of dripwater is mixing along the water flow paths. Waters formed in contact with different PCO2 and distinct bedrock composition, or under various surface conditions (forestation, anthropogenic pollution etc.) are mixed in karst profile above the cave (Perrin et al. 2003a; Perrin et al. 2007; Moore et al. 2009; Schwarz et al. 2009; Gabrovšek & Dreybrodt 2010). After reaching the cave, a drip is formed and the dripwater degasses the excess of dissolved CO2. The driving force of the degassing is the difference between CO2 present in water and CO2 in cave air. Because aqueous concentration of CO2 corresponds to a specific value of PCO2 in contact with the solution (see equation 3), the partial pressure is commonly used to express the content of CO2 in water and referenced to as a CO2 partial pressure in water (PCO2(W)). Therefore, the driving force of dripwater degassing can be conveniently described as the difference between PCO2(W) and PCO2(atmosphere). As the CO2 is released from the water, the driving force diminishes until the partial pressures equalize. Dripwater degassing carbon dioxide is an important source of CO2 in

27 cave air (Bourges et al. 2001; Baldini et al. 2008). As the dripwater degasses, the carbonate equilibria shift and the water becomes supersaturated with respect to calcite, which is followed by calcite precipitation and speleothem formation.

2.3 PALEOENVIRONMENTAL RECONSTRUCTIONS Cave dripwaters and their precipitates (speleothems) are studied worldwide as a source of information about environment in geological past (see McDermott 2004 or Fairchild et al. 2006a for review). As many of processes participating on water composition are climatically controlled, speleothem composition can be used to reconstruct paleoenvironmental conditions during its formation (Li et al. 2005; Griffiths et al. 2010; Borsato et al. 2016; Paar et al. 2016). Most common technique is stable isotope analysis (McDermott 2004; Drysdale et al. 2005; Verheyden et al. 2008; Lachniet 2009), nevertheless also trace elements can provide useful data (e.g., Verheyden et al. 2000; Huang et al. 2001; McMillan et al. 2005; Fairchild et al. 2006b; Cruz et al. 2007; Wong et al. 2011; Frisia et al. 2012; Jochum et al. 2012; Sinclair et al. 2012; Meyer et al. 2014; Tan et al. 2014; Orland et al. 2014; Casteel & Banner 2015; Bernal et al. 2016). The enhanced trace element ratio is interpreted as an effect of PCP or longer residence time in arid climate, and therefore is frequently tested as a paleoenvironmental proxy (Verheyden et al. 2000; Fairchild et al. 2000; Fairchild & McMillan 2007; Fairchild & Treble 2009; Tremaine & Froehlich 2013). In fact, Mg/Ca ratios may sometimes paradoxically show positive correlation with rainfall (e.g., Baldini et al. 2012). In the last few years, more thorough studies indicate that dripwater properties are climate sensitive to lesser extent than previously anticipated (e.g., Baker et al. 2016). Furthermore, the variations in dripwater hydrogeochemistry are not necessarily linked with changes in drip discharge (Musgrave & Banner 2004; Faimon et al. 2016) or the response can be non-linear (Karmann et al. 2007). Additionally, it seems that changes in calcite fabrics and crystal habits might be also a record of saturation and water discharge variability (e.g. Genty & Quinif 1996; Frisia et al. 2000; Niggemann et al. 2003 or Riechelmann et al. 2014).

28 3 RESULTS AND DISCUSSION

3.1 HYDROGEOCHEMISTRY OF DRIPWATERS IN MORAVIAN KARST Hydrogeochemical properties of dripwaters directly determine speleothem growth and composition. In order to study differences between dripwaters in a cave system, a long term monitoring of dripwater properties and cave environment was realized in Punkva Caves in Moravian Karst. Sampling sites were located in a corridor behind Přední Dome (drips PC), in Zadní Dome (drip ZD) and in Tunnel Corridor (drips TC). Collected data were furthermore compared with older dripwater research on other sites in Moravian Karst to assess possible similarities/differences between particular caves in one karst region. The study was realized from February 2012 to March 2013 and 126 samples were collected during 26 sampling events (twice per month) and compared with 45 analyses from the archive dataset. Meteorological data were provided by the Cave Administration of the Czech Republic from a station situated above the cave and operated by the Czech Meteorological Institute. The sampled dripwaters were partially analyzed directly on site (volumetric determination of alkalinity and Ca) and in a laboratory (ICP-OES analysis of Mg and Sr). Drip- waters were also twice (April and November) sampled for stable isotope analysis (LAS; δ18O, δ2H and δ13C). For detailed description of the site and used methods, see Appendix 1 and Appendix 2. Studied drips showed different flow regimes regardless of drip location – drip CP2 had very stable discharge (variation coefficient 17.6%), whereas drips CP3 and TC1 showed higher variation (v.c. ~50%) and drips CP1, TC2 and ZD showed very wide variations. The drip TC2 eventually even stopped dripping. With exception of the stable drip CP2, the decreasing discharges indicate, that the epikarstic perched aquifers feeding the drips were gradually emptied from May to December. Interestingly, no drips showed response to summer storm events, which might be caused by very wide measurement step. Another con- tributing factor might be evapotranspiration as shown by a model (see Appen- dix 1: Fig. 6). The discharges started to rise again in mid-December and rose

29 until the end of measurements in March. This development indicates a substantial replenishment of epikarst aquifers from snowmelt water. Neverthe- less, drip CP3 did not increase discharge and it seems that the perched aquifer is replenished under specific conditions. In addition, a delay of 3–4 months between modeled infiltration and drip discharges suggests that the water is seasonally drained through some preferential pathways. In such case, it would not significantly contribute to epikarst aquifers feeding the studied drips. The studied dripwaters were divided into two groups according to their hydrogeochemical properties. On one side, there is the anomalous drip TC1, on the other side the rest of the drips referred to as the regular drips. These drips showed higher values of EC, significant supersaturation with respect to calcite and low trace elements ratios. Isotopically, the drips were very similar to each other with no signs of additional processes. They demonstrated slight enrichment in δ18O and δ2H in November samples and δ13C values corresponding to calcite dissolution under closed system conditions. In contrast, the anomalous drip showed peculiar properties – systematically lower EC, saturation close to equilibrium with respect to calcite and enhanced Mg/Ca and Sr/Ca ratios compared to regular drips. These ratios were caused by significantly lower Ca concentration. The isotopic composition showed strong enrichment in δ13C indicating water degassing, but the δ18O and δ2H values were similar to values in regular drips. The δ18O and δ2H are close to Global Meteoric Water Line (GMWL) and indicate very limited effect of evaporation and fast precipitation infiltration. What is more, the summer/fall difference in δ18O values indicates incomplete mixing in the epikarst aquifer as the isotopically lighter summer rainwater remains at least partially differentiable. A few possible explanations of anomalous drip properties were proposed and discussed. Firstly (a) an enhanced water dynamic in karst profile – a situation when infiltrating dripwater reaches the cave before it could attain equilibrium with soil/epikarstic CO2 and calcite. It is therefore under-saturated with respect to calcite, shows low mineralization and Mg/Ca ratio and very variable discharge with strong reaction to rainfall. However, these properties do not fit the anomalous drip. Another possible explanation is (b) water mixing

30 along the water flow path somewhere in vadose zone. Although water mixing can explain low saturation or under-saturation with respect to calcite and lower mineralization, it does not seem possible – considering Moravian Karst limestones composition – that it could provide enhanced trace element ratios in Punkva Caves. In conclusion, an explanation of anomalous properties via (c) prior calcite precipitation (PCP) seems as the most plausible. Effects of PCP on dripwater include lower mineralization and saturation with respect to calcite, enhanced trace element ratios and increase in δ13C, while not necessarily influencing the dripwater hydrology. Even the speleological situation on site supports PCP hypothesis, as the water flows through ca. 20 long crevice with many speleothems before being sampled (Glozar 1984). Considering the possible speleothem corrosion, the anomalous dripwater does not seem to pose threat to cave environment. Its properties and supposed origin via PCP suggest water in equilibrium or slight supersaturation with respect to calcite and cave CO2. An unrealistically high PCO2 in cave air would be necessary for the water to become aggressive to calcite as demonstrated in Výpustek – another Moravian Karst cave – where influence of ventilation and increased PCO2 during large cultural events was studied (see chapter 3.4 and Appendix 4). The anomalous dripwater geochemical properties would be imprinted into a speleothem growing from the water, raising numerous questions regarding speleothem utilization as climatic proxies. The trace element incorporation in calcite is interpreted as indicator of temperature changes (Roberts et al. 1999; Huang & Fairchild 2001), dry periods (Fairchild and McMillan 2007) or changing elemental sources (Ayalon 1999). In contrast to general understanding (e.g. Fairchild et al. 2000), the correlation of Mg/Ca to the drip discharge in Punkva Caves does not show negative correlation. Although a several years long dataset would be necessary to assess the effect of long term arid periods, no intra- seasonal changes were observed. Another problematic utilization of a speleothem precipitated from anomalous dripwater is evaluation of calcite precipitation dynamic from crystal habits (e.g., Frisia et al. 2000; Niggemann et al. 2003

31 or Riechelmann et al. 2014). The final problematic interpretation is the utilization of δ13C isotopic data. The δ13C is supposed to be determined by soil CO2 composition (Dreybrodt & Scholz 2011), but the anomalous drip shows enrichment caused by the PCP. An anomalous dripwater speleothem would therefore incorrectly indicate lower biological activity in soil. The properties of the anomalous drip are of permanent character and sup- posedly controlled spatially, implying that the spatial conditions might be an important factor that should be considered with the temporal conditions (e.g. dry/wet seasons). Moreover, these spatial effects might change over long periods as karst/cave conditions develop (e.g. by weathering and karst evolution). It could be concluded that it would be appropriate to use more speleothems from one cave system in any paleoclimatic reconstruction based on speleothem proxies to eliminate possible distortion by anomalous speleothems.

3.2 CAVE AND EPIKARSTIC PCO2 The amount of limestone dissolved in dripwater is reflected by Ca2+ and

CO32− content that determines solution’s undersaturation or supersaturation with respect to calcite. The total quantity of dissolved calcite is limited by the conditions under which the dissolution occurred, especially PCO2 in soil/epikarst aquifer. When the percolating water enters a cave, it immediately starts to degas CO2, the pH rises as well as CO32− content as the carbonate balance shifts and the water becomes supersaturated to calcite. This subsequently leads to calcite precipitation and speleothem growth. The intensity of degassing is dependent on the difference between PCO2 equivalent present in the water and

PCO2 in the cave’s air: the bigger the difference, the higher supersaturation can be reached. Based on the hydrogeochemical properties of dripwater, the estimate of hypothetical CO2 partial pressure under which the water was formed can be made (Faimon et al. 2012; Peyraube et al. 2012; Milanolo & Gabrovšek 2015; Pracný et al. 2016b). A model based on reconstruction from calcium and carbonate species concentrations and pH was composed to study hypothetical

PCO2 participating on formation of dripwaters in Punkva Caves in Moravian

Karst (Czech Republic) and compare them with PCO2 in the cave’s air and in

32 sampled dripwaters to better understand some of the processes determining speleothem growth. Punkva Cave system is developed in central part of Moravian Karst in Devonian limestones of the Macocha formation. Studied dripwaters are situated in upper dry level of the cave under ca. 100 m thick layer of limestone.

Concentrations of CO2 in Moravian Karst reach up to 1% in soils (Faimon & Ličbinská 2010; Blecha & Faimon 2014) and 1–11% in caves (Otava 1995; Fai- mon et al. 2012). Dripwater and speleoclimatic data were collected during 26 measuring campaigns from February 2012 to March 2013. Meteorological data were obtained from measuring station situated on the surface above the cave. All data were statistically analyzed for presence of significant correlations and presence of cycles. For detailed description of the site and used methods, see Appendix 1 and Appendix 2.

The partial pressure of CO2 in the cave’s air (PCO2(air)) showed significant seasonality with calculated period of about 304 days. The minima were measured in winter (10−3.31, i.e., 0.06 vol%) and the maxima in summer (10−2.49, i.e.,

0.32 vol%). Both the partial pressure of the CO2 corresponding to aqueous carbonates, PCO2(W), and the partial pressure of CO2 participating on the initial water formation, PCO2(H), were calculated from dripwater hydrogeochemistry.

Whereas the PCO2(W) showed clear seasonal variations with minima in winter (10−2.91, i.e., 0.12 vol%) and maxima in summer (10−2.35, i.e., 0.45 vol%) the

PCO2(H) did not show any significant seasonality. The partial pressure was calculated in a narrow range from 10−1.77 to 10−1.49 (1.7–3.2 vol%) indicating only slight dependence on surface conditions.

The cave air CO2 seasonality might be driven by the difference between exterior and interior temperature. Whereas the daily temperature maxima below mean annual temperature (MAT) indicate upward airflow ventilation mode (UAF mode, totally 141 days), the daily temperature minima above MAT indicate daily downward airflow mode (DAF mode, totally 105 days) in the cave and both modes represent the periods of active cave ventilation (Faimon et al.

2012). In UAF mode, the PCO2(air) is systematically lower, as the air flows out from the cave through upper openings, which contrasts with DAF mode, when

33 the PCO2(air) rises probably by influx from cracks and joints leading through epikarst and soil (Lang et al. 2017). Nevertheless, the correlations of external temperature with stationary PCO2(air) data are not significant, which might be caused by time shifts in monitoring steps. In addition, neither the internal temperatures correlate with PCO2(air).

The periodicity in PCO2(W) is less pronounced than periodicity of PCO2(air) and with different periods. The raw data are correlated statistically significantly, which indicates interconnection via degassing. No significant correlation was found in the stationary data, which might be caused by additional effect, e.g., drip rate variations or variations in sampling. If the drip rate is slow, the residence time of an individual hanging drop on the speleothem increases and it can degas much more CO2 than under higher discharge. The intensity of degassing is also affected by the difference between the initial

PCO2(W) of water entering cave and PCO2(air) (Faimon et al. 2016).

In contrast to PCO2(air) and PCO2(W), the hypothetical PCO2(H) shows very stable values for all drips except for the previously identified anomalous drip. The stable PCO2(H) values in all other drips (the regular drips) indicate independence on surface conditions and suggest that the CO2 source might be situated in deeper parts of epikarst or vadose zone. Nevertheless, the precise source and presumed biogeochemical conditions are unknown. Moreover, the estimated values of PCO2(H) are higher than concentrations measured in soils (Faimon & Ličbinská 2010; Sanchez-Cañete et al. 2011; Plestenjak et al. 2012; Blecha & Faimon 2014).

On the contrary, the PCO2(H) reconstructed for the anomalous drip shows seasonal variations similar to PCO2(air) and PCO2(W). Statistical correlation as well as the calculated periodicity demonstrate a strong connection. The close relation of the partial pressures indicates existence of a mechanism disrupting the original PCO2(H). A model of geochemical evolution (based on Peyraube et al. 2012 and Milanolo & Gabrovšek 2015) illustrates the difference between the regular drips and the anomalous drip (see Appendix 2). The model shows that regular dripwaters have undergone uniform degassing evolution shifted only by different initial conditions, i.e., by different PCO2(H). In addition, the water

34 from anomalous drip seems to have firstly degassed to equilibrium with PCO2(air) and then precipitated calcite. This sequence is especially plausible considering kinetics of these processes. Therefore, the most probable cause of the anomalous properties seems to be prior calcite precipitation in some caverns above sampling site.

3.3 MODELING LIMESTONE DISSOLUTION AND MG/CA EVOLUTION IN EPIKARST The main source of minor (e.g. Mg) and trace (e.g. Sr) elements in dripwaters is limestone dissolution. The mechanism and dynamics of Mg release from carbonates is generally neglected in paleoclimatic studies despite using Mg/Ca ratio as an important proxy. In order to examine this assessment, a model of theoretical evolution of Mg/Ca ratio during limestone dissolution under conditions expected in epikarst was composed and compared with dripwater data from Punkva Caves in Moravian Karst and other caves worldwide. The study was focused especially on the effect of Mg-calcite and dolomite dissolution dynamics. Limestone dissolution was studied using a dynamic model of conjoined dissolution of two separate minerals, Mg-calcite and dolomite, while calcite and magnesite were initially not present in solution, but were allowed to precipitate when the solution became supersaturated with respect to the minerals during simulation. Series of model calculations were run for epikarstic conditions of T

= 10 °C and log PCO2 = −1.5, fixed rock/solution ({L}/V) and water/atmosphere ({S}/V) boundary area and a range of dolomite/Mg-calcite ratios (D/C). For detailed description, how the ratios were estimated, see Appendix 3. The model is an open system where CO2 from soil/epikarst air enters the solution. The

CO2 exchange is derived from the two-layer model (Liss & Slater 1974; Stumm & Morgan 1996). The detailed descriptions of Punkva Caves limestones and dripwaters compared to the model solutions are presented in Appendix 3. The model showed that the reaction path of dissolution of composite carbonates (dolomite or Mg-calcite) initially follows a straight line determined by mineral stoichiometry indicating congruent dissolution. When the solution reaches calcite saturation, the character of dissolution changes to incongruent

35 – this is illustrated in the model by the reaction path becoming non-linear and increasing the slope up to negative values. It is due to decline in release of Ca compared to Mg and subsequent decrease of Ca concentration, while Mg concentration is continuously increasing. The incongruence is caused by precipitation of calcite, to which the solution is supersaturated, while still dissolving the Mg-bearing mineral. The composition of Mg-calcite determines shape of the reaction path. Model of concurrent dissolution of Mg-calcite and dolomite mix (representing limestone) shows very similar development with reaction path dependent on percent of Mg in calcite and dolomite/calcite ratio. Nevertheless, the effect of dolomite component is almost indistinguishable from pure Mg-calcite dissolution if the D/C ratio is less than 1 as the Mg-calcite kinetics are much faster. Therefore, in most limestones is the Mg-calcite composition a key factor determining Mg/Ca evolution. Compared to real dripwater data, most of the dripwaters follow the reaction paths defined by the dissolution model. The most probable cause of the data lying elsewhere is that the site conditions are different from conditions in Moravian Karst, upon which the model is based. For example, if the water was formed under higher PCO2, the resulting solution would show enhanced Ca concentrations. Another defining parameter is temperature because of the rate constant dependence. The wide range of Ca concentrations and relatively narrow range of Mg concentrations indicates dissolution of very low-Mg calcites and only minor dolomite component. This seems to be the case not only of Punkva Caves’ dripwaters but also of other cave systems, where the Ca concentrations in dripwaters show much wider range than Mg concentrations (Immenhauser et al. 2010; Riechelmann et al. 2011). In contrast, other sites evolve along the reaction paths for dissolution of limestones with a significant dolomite component or dolostone (Fairchild et al. 2000; Wong et al. 2011). These paths are gradually straighter with increasing dolomite component and the reaction path of dolomite is a straight line. Furthermore, dynamics of the modeled reaction paths are determined by the {L}/V ratio, the limestone-solution boundary. In addition, the ratio modifies

36 the reaction path shape in transition from congruent to incongruent dissolution. Generally speaking, very high {L}/V ratio can be expected in epikarst, where the water is stored in pores and spaces between grains/clasts and fractures, whereas in fissures in deeper parts of vadose zone the {L}/V is much lower. Moreover, the perched aquifer forms an open system, where CO2 consumed by carbonates dissolution can be replenished from soil/sediments, whereas fissure flow is more of a closed system with very limited supply of free

CO2 in solution. These conditions seem to contribute to dissolution dominantly occurring in epikarst – as illustrated by the limestone corrosion diminishing downward the vadose zone. Therefore, although the value of {L}/V is generally unknown, it seems to be of high importance. Thus, the actual composition of dripwater is given not only by the reaction path shape, but also by its position along the reaction path, i.e., how far the dissolution proceeded (Appendix 3: Fig. 2 and 3). This underlines the importance of residence time in karst water formation. Under expected {L}/V ratio, the model of epikarst dissolution of Moravian Karst limestone indicates the residence times for regular dripwaters in the range from 100 to 150 days. These values are very well plausible compared to previously estimated residence times (Kamas et al. 2015; Faimon et al. 2016). Intriguingly, interpretation of the high Mg/Ca ratio in the anomalous drip based on the model would indicate extreme residence times in the range of hundreds of days. Nevertheless, the anomalous properties are most probably a result of PCP as previously explained (Appendix 1). It shows that application of the model on real dripwater data could be limited because additional processes may participate on water hydrogeochemistry evolution in addition to dissolution. The arguably most studied process possibly influencing Mg/Ca ratio is the prior calcite precipitation (PCP) (Fairchild et al. 2000). Precipitating calcite consumes dissolved Ca and therefore increases the Mg/Ca ratio. It is believed that PCP is promoted by arid and warm climate, whereas the changes in Mg/Ca might be also caused by rapid rainfall events. During such events, infiltrating water might ‘flush out’ the water stored in semi-isolated reservoirs. Such water could show enhanced Mg/Ca ratio due to prolonged residence time. In addition, PCP leads to decrease

37 in Ca concentration, whereas Mg concentration remains the same, while the dissolution dynamics affect both Ca and Mg concentrations. This might be an important distinguishing factor in dripwater, although it is indistinguishable in precipitated speleothems. Other mechanisms participating on dripwater Mg/Ca ratios might be e.g. preferential Mg leaching from fresh surfaces (McGil- len and Fairchild 2005; Morse et al. 2007; Sinclair 2011), cation capture on mineral surfaces (e.g., on clay minerals present in epikarst) or dissolution of additional Mg-rich minerals (e.g., evaporites present in limestone). In paleoenvironmental studies evaluating the Mg/Ca ratio, the effect of PCP on dripwater formation (and therefore speleothem composition) is usually favored above any dissolution dynamic influence. This study identifies further significant factors besides incongruent dissolution due to different dissolution dynamics of carbonate minerals. These are (1) the limestone composition (Mg- calcite composition and dolomite content), (2) the residence time, (3) the ratio of limestone surface area to water volume (i.e., the karst system structure) and

(4) the epikarstic PCO2. With exception of PCO2, all these factors are intrinsically water-flow-path dependent. Therefore, they may change both spatially and temporarily with the natural evolution of the karst system, independently on climate changes. Finally, it is important to note, that very similar effects can be expected from mixing of waters formed at contact with different limestones.

3.4 ANTHROPOGENIC CO2 INFLUENCE ON DRIPWATERS AND SPELEOTHEM CORROSION Cave air ventilation was monitored in the Bear Chamber of the Výpustek Cave in Moravian Karst during events with large visitor numbers. The natural cave air CO2 levels in the Výpustek Cave (0.05–0.10 vol%, i.e., PCO2 = 10−3.32 to

10−2.98) are comparable with the Punkva Caves (0.06–0.32 vol%, i.e., PCO2 = 10−3.31 to 10−2.49), albeit being lower. The difference might be caused by seasonally more representative dataset from Punkva Caves. The study in the

Výpustek Cave showed that the anthropogenic influx of CO2 could lead to increased steady state concentrations up to PCO2(anthrop) = 10−2.22 (0.61 vol%), depending on the number of visitors and their duration of stay in the cave (for modeling details see Appendix 4).

38 Additionally, dripwaters were sampled for a study of hydrogeochemical properties. Due to limited occurrence of suitable permanent drips in the cave, only two, D1 and D2, were studied in detail. Both drips showed little seasonal variations. The drip D1 had higher specific conductivity (736–775 µS cm−1) compared to drip D2 (360–395 µS cm−1). Both dripwaters showed supersaturation with respect to calcite with higher values for D1 (SI = 1.03–1.19) than for D2

(SI = 0.58–0.64). Partial pressures of CO2 corresponding to aqueous carbonate species, PCO2(W), and hypothetical PCO2 under which the water was formed in soil/epikarst, PCO2(H), were calculated (for calculation background see chapter

3.2 and Appendix 2). Whereas the values for D1 (PCO2(H) = 10−(1.56–1.85)) were very similar to properties of other dripwaters in Moravian Karst (e.g., regular dripwaters in Punkva Caves show PCO2(H) = 10−(1.49–1.77)) the values for D2 (PCO2(H) = 10−(2.27–2.51)) were somewhat lower (Pracný et al. 2016). The ventilation model was compared to hydrogeochemical model and showed that the peak PCO2(anthrop) could principally exceed dripwater PCO2(H) turning the water into solution aggressive with respect to calcite. Further modeling and analysis was focused on drip D2 because it requires much lower CO2 concentrations than dripwater D1 to be converted into aggressive solution and is therefore of greater environmental concern. Depending on the ventilation mode the duration of visitors’ stay in cave required for PCO2(anthrop) to outreach

PCO2(H) is between 4.24–7.25 h for the ordinary visitor group (50 people) and 0.42–0.73 h for the enhanced attendance (500 people). For comparison, an individual guided tour usually stays in the chamber for 0.25 h and this time would require thousands of visitors to balance the partial pressures. Such attendance is unreal given the cave capacity. Although an event with a large attendance and longer duration of stay (e.g., concert, performance or wedding) might cause conversion of some dripwater to a solution aggressive with respect to calcite, it seems that the cave dripwater cannot be converted by ordinary tours under current regime. Nevertheless, if condensation waters on cave walls are considered, the probability of corrosion increases. The condensing water dissolves gaseous CO2

39 and becomes aggressive to calcite. Therefore, any increase in cave air PCO2 results in increasing potential of speleothem corrosion, although not via dripwater conversion.

40 4 CONCLUSIONS

The aim of the presented research of hydrogeochemical properties of dripwaters in Moravian Karst was to contribute to general understanding of conditions and mechanisms of dripwater formation. Insight into these processes pro- vides (1) consolidation of basic karst genesis concepts, (2) better paleoenvironmental analyses and (3) knowledge applicable in more efficient karst/cave protection. The long-term dripwater monitoring in Punkva Caves showed no immediate relation of meteorological conditions on karst surface (temperature, precipitation) to dripwater properties. Dripwater discharge showed seasonal variation with minima in summer/fall and maxima in winter/spring. The water amount seemed to be strongly affected by evapotranspiration and significantly recharged by snowmelt. Furthermore, not all drips show the same hydrological behavior, as some dripped only seasonally while other showed only very slight discharge variations. The hydrogeochemical properties of dripwaters in Moravian Karst were a foundation for the drips’ division into two groups: (a) the regular drips (high

SIcalcite and EC, low Mg/Ca and Sr/Ca ratios) and (b) an anomalous drip (low

SIcalcite and EC, high Mg/Ca and Sr/Ca ratios). The δ2H and δ18O isotopic compositions of drips in both groups were the same, but the δ13C enrichment indicates anomalous dripwater degassing. In other properties (esp. discharge) the anomalous drip does not differ from regular drips. Prior calcite precipitation was identified as the most probable cause of the anomalous properties, followed by water mixing in the vadose zone. In addition, although the anomalous drip shows low saturation with respect to calcite, it seems incapable of becoming substantially undersaturated and corroding speleothems.

The study of CO2 in the air and water in Punkva Caves showed (I) seasonality of the cave air PCO2 with maxima in summer and minima in winter with the period slightly shorter than 1 year. The proposed cause is the cave ventilation driven by the difference of cave temperature and external temperature. Seasonality, albeit with different periods and less conclusive correlations, was

41 also showed by (II) the PCO2 corresponding to aqueous carbonates. The controlling factor seems to be water degassing dynamics determined by the CO2 concentration gradient between the water and cave air. A geochemical model was used to reconstruct (III) the hypothetical PCO2 in epikarst – the partial pressure of CO2 under which the water was formed. The results showed very stable values without significant seasonality, indicating independence on surface conditions. Thus, a CO2 source is probably situated deeper in epikarst/vadose zone rather than in karst soils that are controlled seasonally. Possibility of speleothem corrosion via dripwater conversion under increased CO2 concentration due to anthropogenic influx was studied in the Výpustek Cave. Modeling showed that although it is possible to reach concentrations causing undersaturation of dripwater with respect to calcite during events with enhanced attendance, ordinary guided tours seem to be incapable of substantially affecting dripwaters. Nevertheless, the effect on cave condensation waters might be of importance. The comparison of dripwater data with results of geochemical modeling of limestone dissolution under epikarstic conditions showed that Mg/Ca ratios are indeed controlled by incongruent dissolution of Mg-calcite, lesser of dolomite. Initially, the dissolution proceeds congruently following a straight path. When the solution reaches saturation with respect to calcite, the dissolution changes to incongruent, as part of material released from Mg-calcite or dolomite is precipitated as calcite, while Mg accumulates and the Mg/Ca ratio slowly increases. The slope of reaction path nonlinearly increases and becomes negative, resulting in bent overall reaction path. Most of compared real dripwater data follow these reaction paths with a few exceptions caused probably by distinct conditions in individual karst systems (e.g., temperature and PCO2). The shape of reaction path (and resulting Mg/Ca ratio) primarily depends on dissolved rock composition. Besides, the Mg/Ca ratio is given by the dynamics of system evolution, i.e., by the distance along the reaction path that the system covers during dissolution. This distance is controlled by the extent of limestone/water boundary and by the overall water residence time.

42 It is apparent that the Mg/Ca ratios are dependent on additional factors beside the incongruent dissolution. These are (1) the limestone composition (ratio of calcite/dolomite component and Mg-calcite composition), (2) the time of water/rock interaction (water residence time), and (3) the ratio between water volume and rock surface during dissolution (effectively the karst structure). Independently on climatic conditions, all these factors might change with the variations in water flow paths. The results lead to the conclusion that numerous processes and factors usually considered negligible in paleoclimatic reconstructions might be of substantial importance. These processes may disrupt desirable features (e.g., dissolution dynamics affecting Mg/Ca ratio in the same fashion as PCP, or crystal growth and crystal habitus) and should be at least considered in paleorecon- structions. For example, both the model of Mg/Ca evolution and the difference in hydrogeochemical properties of the anomalous drip compared to the regular drips show much higher importance of water flow path in dripwater formation than anticipated. The flow path not only determines residence times, the limestone/water boundary, possibility of prior calcite precipitation or water mixing in vadose zone. More importantly, it can change both spatially and temporally independently on climatic variables as the karst system naturally evolves. A more complex analysis of various factors impacting dripwater forming speleothem may result in more relevant interpretation of paleoenvironmental data.

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53 APPENDIX 1

This appendix presents following research paper:

Pracný, P., Faimon, J., Sracek, O., Kabelka, L., & Hebelka, J. (2016). Anomalous drip in the Punkva caves (Moravian Karst): relevant implications for paleoclimatic proxies. Hydrological Processes, 30(10), 1506–1520. http://doi.org/10.1002/hyp.10731

54 HYDROLOGICAL PROCESSES Hydrol. Process. 30, 1506–1520 (2016) Published online 30 November 2015 in Wiley Online Library (wileyonlinelibrary.com) DOI: 10.1002/hyp.10731

Anomalous drip in the Punkva caves (Moravian Karst): relevant implications for paleoclimatic proxies

Pavel Pracný,1* Jiří Faimon,1,2 Ondra Sracek,2 Ludvík Kabelka3 and Jiří Hebelka4 1 Department of Geological Sciences, Faculty of Science, Masaryk University, Kotlářská 267/2, 611 37, Brno, Czech Republic 2 Department of Geology, Faculty of Science, Palacky University, 17. listopadu 12, 771 46, Olomouc, Czech Republic 3 Hydrochemical Laboratories, GEOtest, a.s., Šmahova 1244/112, 627 00, Brno, Czech Republic 4 Cave Administration of the Czech Republic, Svitavská 11, 678 01, Blansko, Czech Republic

Abstract: The anomalous drip in the Punkva caves (Moravian Karst) shows specific hydrogeochemical properties such as low 3 1 13 SIcalcite ~ 0.14 ± 0.11 (standard deviation), low mineralization (4.53 ± 0.42) × 10 mol l , and enhanced values of δ C(7.85 to 8.35‰ VPDB), Mg/Ca × 1000 ratio (45.7 ± 3.3), and Sr/Ca × 1000 ratio (0.65 ± 0.06). By these properties, the anomalous drip significantly differs from other regular drips in the same cave and other caves in the region. The study suggests that the anomalous drip properties are a consequence of prior calcite precipitation or/and water mixing along the water flow path. As the former processes are spatially controlled, the knowledge of dripwater flow path seems to be necessary for correct paleoclimatic/paleoenvironmental reconstructions. Copyright © 2015 John Wiley & Sons, Ltd.

13 18 KEY WORDS cave dripwater; stable isotope C and O; Mg/Ca and Sr/Ca ratios; prior calcite precipitation; karst water mixing Received 24 June 2015; Accepted 21 October 2015

INTRODUCTION comprehensive characterization of several drips from the same cave/region is a critical step in defining possible Autochthonous cave deposits (speleothems) are frequent- differences in stalagmite records. ly used as terrestrial archives of paleoclimatic data (e.g. The goal of the study is to introduce an anomalous Fairchild et al., 2006 or Fairchild and Treble, 2009, for a dripwater from the Punkva caves (Moravian Karst), the review). This utilization is based on an assumption that composition of which is strongly influenced by processes external surface conditions are projected/encoded into in the epikarst/vadose zone. Its hydrology and chemistry speleothems (Verheyden et al., 2003; Li et al., 2005; Cruz is compared with the ‘regular dripwaters’ from different et al., 2007; Verheyden et al., 2000; Griffiths et al., sites in the same cave and in other caves in the same 2010). Unfortunately, the link between the surface region. The anomalous properties of water are expected to conditions and speleothems is influenced by many factors change dramatically morphological/geochemical proper- on the water reaction–transport path that may disturb the ties of a potentially formed speleothem: growth fabrics climatic signal (Baldini et al., 2006; Karmann et al., (lamina thickness), trace elements contents (Mg, Sr), and 2007; McDonald et al., 2007; Miorandi et al., 2010; stable isotopes values (δ13C). A prospective Sherwin and Baldini, 2011). Therefore, it is extremely paleoenvironmental study based on such properties might important to study recent processes in detail to see how lead to incorrect interpretations. they influence the climatic proxies. Some studies have shown that coeval stalagmites from the same cave may exhibit different trace element patterns (Roberts et al., SITE OF STUDY 1999; Finch et al., 2003). This either suggests that some stalagmite properties are not representative of climatic The Punkva caves are the best known and most popular conditions or that different drip sites preserve distinct show caves in the Czech Republic. They are situated components of the climate signal. Therefore, a more approximately 20 km northwards from Brno in the Moravian Karst, the largest karst region in the country. The caves have developed in very pure Devonian *Correspondence to: Pavel Pracný, Department of Geological Sciences, limestone of the Macocha formation, specifically in the Faculty of Science, Masaryk University, Kotlářská 267/2, 611 37 Brno, ž Czech Republic. La ánky and Vilémovice Limestones (Faimon et al., E-mail: [email protected] 2012; Blecha and Faimon, 2014; Pracný et al., 2015).

The study was conducted during the period from Meteorological data are from the meteorological station February 2012 to March 2013 in the Punkva caves situated close to the Macocha Abyss Upper Bridge and (Moravian Karst). Dripwaters were sampled in a corridor run by the Czech Meteorological Institute and the Cave behind Přední Chamber (drips CP1, CP2, and CP3), Administration of the Czech Republic (Figure 1). Tunnel Corridor (drips TC1/the anomalous drip, TC2), A set of archive data was provided for a comparison. and Zadní Chamber (drip ZD) in the so-called dry part of The data were collected from June 2003 to May 2004. the caves (Figure 1). Waters in the CP1, CP2, CP3, and The data come from a remote part of the Punkva caves TC2 sites have dripped from small straw stalactites on the and three different caves in the Moravian Karst (Figure 1). corridor ceiling. The water of TC1 has dripped from a Two drips, one from a stalactite, MD1, and one from a curtain about 30 cm wide developed on the edge of a straw stalactite, MD2, are situated in the Masaryk Dome crevasse. The water of ZD dripped on the top of Chamber, Punkva caves. Another drip comes from a stalagmite ‘Vase’ from the height of about 20 m. The straw stalactite in Big Foch’s Dome Chamber in the sampling was conducted twice per month: 126 samples in Balcarka cave (BC). Two drips are in the Amatérská cave; total were collected during 26 sampling courses. Because the ﬁrst one drips from a straw stalactite in the crossroad of their extremely low drip rate during part of the year, the of the adit and Javor Corridor (AC1), and the second one drips TC2 and ZD were not sampled in all cases. comes from a curtain in the Rozlehlá Corridor (AC2).

Figure 1. Sketch map of the Punkva caves system and its localization within the Moravian Karst and the Czech Republic. For explanation of the drip site acronyms, see the text

There were nine sampling courses with chemical analysis regimes: the drip CP2 was very stable (variation for all the archive drips and six courses with just drip coefficient 17.6%; Figure 2e), whereas drips CP3 and discharge measurement. TC1 were more variable (variation coefficients 59% and 43.2%, respectively, Figures 2f and g). More variable rates were found for the drip CP1 (from 6 to MATERIALS AND METHODS 174 drops min1, variation coefficient 87.1%; Figure 2d) 1 Drip discharge was measured by counting drips during a and ZD (rate from 3.5 to 167 drops min , variation fi given time period. The drip size was found by measuring coef cient 149.8%; Figure 2c). The drip TC2 was fi the weight of a water drop caught into a plastic container extremely variable (variation coef cient 230%). Its rate < (<0.9 g) using digital scales. An average value based on was extremely slow ( 1 drop per 5 m) or zero during a three measurements from individual drip sites was used substantial part of the monitoring period. for drip rate calculations. The archive data (not presented in Figure 2) show the – 1 Immediately in the cave, the basic hydrogeochemical most variable drip AC2 (the rate 50 400 drops min , fi parameters were determined: pH, specific electrical con- variation coef cient 77.6%). Other drips are much slower – 1 ductivity (EC), alkalinity (by acidimetric titration with (the drip BC with rate of 0.5 4 drops min , variation fi evaluation of titration curve via the Gran’s function, coef cient 60.9%; the drips MD1 and AC1 with the rates – 1 fi Stumm and Morgan, 1996), and Ca concentration of 8 18 drops min , variation coef cient below 25%; the – 1 (complexometric microtitration using calcein as indica- drip MD2 with the rate of 3 5 drops min ; variation fi tor). Other analyses (Mg, Sr) were conducted in a coef cient below 25%). laboratory using ICP (iCAP 6000 by Thermo Scientific). Saturation indices were calculated using the PHREEQC Hydrogeochemistry code (Parkhurst and Appelo, 2013). Totally, 126 new Specific ECs of the common drips TC2, CP1, CP2, CP3, water samples were analysed and compared with 45 and ZD (representing water mineralization) varied in the analyses from the archive dataset. range 604 ± 32 μScm1 (standard deviation) (Figure 3a). Dripwaters for analyses of stable isotopes were Mean values of other parameters were as follows: sampled twice, in spring (April) and fall (November) of pH ~ 8.06 ± 0.13, alkalinity ~ (5.68 ± 1.34) × 103 eq l1, 2014, and stored in airtight sealed plastic bottles. Isotope Ca concentrations ~ (3.38 ± 0.19) × 103 mol l1, Mg con- δ18 δ2 values O and H in water were determined at the centrations ~ (5.71 ± 0.31) × 105 mol l1, and Sr concen- Czech Geological Survey in Prague, Czech Republic, trations ~ (1.03 ± 0.08) × 106 mol l1. The Mg/Ca × 1000 using a Los Gatos Research laser absorption spectrom- and Sr/Ca × 1000 ratios for the drips were 17.0 ± 1.4 eter. The results were normalized to the standard Vienna (Figure 3c) and 0.31 ± 0.02 (Figure 3e), respectively. Standard Mean Ocean Water (VSMOW) and reported in In contrast, the drip TC1 showed lower EC ~ 297 δ the -notation. The reproducibility of measurements was ± 22.2 μScm1 (Figure 3a), alkalinity ~ (2.14 ‰ δ2 ‰ δ18 δ13 0.5 for H and 0.08 for O. Precipitates for C ±0.28)×103 eq l1, and Ca concentrations ~ (1.47 ± 0.13) analyses were prepared by adding NaOH and BaCl2 and mol l1. The ratios Mg/Ca × 1000 ~ 45.7 ± 3.3 and then the precipitate was filtered out. In the next step, the 13 Sr/Ca × 1000 ~ 0.62 ± 0.06 were enhanced. The mean Mg BaCO3 precipitate was dissolved by H3PO4, and δ C 5 ‰ and Sr concentrations, (6.71 ± 0.37) × 10 and (9.59 ± was measured with a precision better than 0.05 . Results 0.74) × 107 mol l1, respectively, were roughly compara- were expressed with respect to the Vienna Pee Dee ble with those in other drips. Belemnite (VPDB) standard. Archive data showed the parameters similar to the common drips. The EC values were slightly reduced: the RESULTS dripwaters of the Masaryk Dome Chamber from the Punkva Caves showed EC 437 ± 30 μScm1. The drips in Hydrology the Amatérská cave showed EC 495 ± 35 μScm1.Asan During the monitoring period, daily precipitation exception, the occasional drip in the Balcarka cave reached up to 30 mm. Overall precipitation for the entire showed very low EC 311 ± 14 μScm1, which was near monitoring period was 641.7 mm (Figure 2a). External the values of the drip TC1 (Figure 3b). The Mg/Ca × 1000 daily temperatures ranged from 9.3 °C to 25.9 °C. The ratios varied between 10 and 25 with the only exception mean temperature for the entire period was 7.89 °C for the drip AC1 showing much higher ratio (63 ± 3.9) (Figure 2b). Drip discharges in the cave reached up to (Figure 3d). In this case, however, the drip AC1 showed 180 drops min1 with the local minima in winter enhanced Mg concentrations, (2.53 ± 0.50) × 104 mol l1. (November to February) and the local maxima in spring For the archive data, the Sr/Ca × 1000 ratios were not (March to May). Studied drips showed different flow available.

Figure 2. Meteorological and hydrological data. Daily rainfall (a), external maximum and minimum temperatures (b), drip discharges of ZD (c), CP1, 2, 3 (d, e, f), anomalous TC1 (g), and TC2 (h). The orange columns and the light blue columns indicate the maximum daily temperature over the freezing point and the minimum temperature below the freezing point, respectively, during winter/spring period

Figure 3. Selected hydrogeochemical data of drip waters: electrical conductivity, main dataset (a) and archive data (b); (Mg/Ca) × 1000 ratios, main dataset (c) and archive data (d); (Sr/Ca) × 1000 ratios, main dataset (e)

The results for all dripwater samples on δ2H and δ18O from 71.16 to 73.35‰ VSMOW for δ2H and from isotopes are in Figure 4. The spring values ranged from 10.17 to 10.41‰ VSMOW for δ18O. The results for 70.70 to 74.20‰ VSMOW for δ2H and from 10.30 δ13C are in Figure 5. Majority of dripwater samples to 10.60‰ VSMOW for δ18O. The fall values ranged showed the δ13C values in the range from 10.34 to

Figure 4. Dripwater isotopes δ2D and δ18O

Figure 5. Dripwater alkalinity versus δ13C

10.94‰ VPDB for the spring season and from 10.39 to 10.67‰ VPDB for the fall period. An exception is the drip TC1 with the δ13Cvaluesof7.85‰ VPDB in the spring season and 8.35‰ VPDB in the fall period.

DATA ANALYSIS Precipitation trends Five periods with different precipitation rates (from mean 0.27 to 5.64 mm day1 with the totals from 8 mm to 296 mm) were distinguished during the monitored season Figure 6. Cumulative curves of precipitation (a), infiltration (b), and (Figure 6a and Table I). Punkva caves drips sites drip rates (c) After subtracting potential evapotranspiration (calculation based on Hargreaves and Allen, 2003), just water (below 30 drips per minute) and to 0.093 ml in case of fi in ltration and retention remain (Figure 6b). If the higher drip discharges (over 30 drips per minute), the retention is ignored, the curves in Figure 6b indicate that rates were recalculated into volume units. It was found fi ground water was recharged by in ltrating water during that the drip discharge systematically decreased with the the period from September/October 2012 to March 2013. slope from 6.5 to 23.7 ml day2 during spring– summer–fall period. The only exception was drip CP2, Drip discharge dynamics the rate of which increased with the slope of 3.3 ml day2 Based on the experimentally determined relations that (Table II and Figure 2). In contrast to the infiltrated water, one drop corresponds to 0.102 ml for low drip discharges the cave drip discharges substantially increased in the

Table I. Periods of rainfall with different precipitation

Period Date range Slope [mm day1] Totals [mm] R2 p

#1 23 February 2012 31 May 2012 0.71 65 0.91 <0.001 #2 1 June 2012 17 June 2012 5.64 88 0.97 <0.001 #3 18 June 2012 10 November 2012 2.01 296 0.99 <0.001 #4 11 November 2012 15 December 2012 0.27 8 0.84 <0.001 #5 16 December 2012 12 March 2013 2.1 185 0.95 <0.001 period between January and February 2013 (Figure 2 Trace elements – correlation analysis and 6c). The correlations of trace element ratios with precipi- From a hydrological point of view, there are no tation (cumulative amount of precipitation for 14-day fi signi cant differences between the regular drips and the period before drip sampling) and temperature (mean fi anomalous drip TC1. According to the classi cation of external temperature per 14-day period before drip Smart and Friederich (1986) and Baker et al. (1997), both sampling) are not statistically significant at α =0.05 the anomalous drip TC1 and the regular drip CP2 show (Table III). The only exception is the strong negative fl fl seepage ow regime (stable yearlong out ow with minor correlation of Mg/Ca × 1000 ratio versus temperature for variations), whereas the regular drips TC2, CP3, and ZD the anomalous drip TC1 (r = 0.78). For the same drip, are seasonal drips (they are not active during the whole the cross-correlation analysis considering a time lag fl year and show high ow variability). Even though the between variables showed a negative and statistically fl ow regime of drip CP1 belongs into seasonal drips, it significant correlation, r = 0.45, with the delay of shows yearlong discharge with a higher variability 28 days. In case of the regular drips, the cross-correlation (Figure 7, Table I and II). analysis did not bring any improvement in the presented correlations. The correlations of Mg/Ca ratios with drip discharge (Table III) are mostly negative (the only exception is the drip TC2) but statistically insignificant at α = 0.05, except for the CP3 drip. The correlations of Sr/Ca ratios with drip discharge (Table III) are both positive and negative and completely insignificant at α = 0.05. Mutual correlations of trace elements, Sr versus Mg, are given in Table IV. Just for the drip TC1, there is a positive and statistically significant correlation (α = 0.05; r = 0.40). It is without any time lag, as the cross- correlation analysis has shown. The correlations for other drips are weak, either negative or close to zero. The only statistically significant correlation (α = 0.05; r = 0.49) is for the CP2 drip.

Saturation indices

Figure 7. Hydrological classification of the dripwaters from the Punkva The saturation indices with respect to calcite (SIcalcite) caves (Moravian Karst) based on their mean discharge and variability (PS for common drips in the Punkva caves ranged from 0.8 to – percolation stream, ShF – shaft flow, VF – vadose flow, ScF – subcutaneous flow, SpF – seepage flow). Based on Smart and Friederich 1.1. In contrast, the anomalous drip TC1 varied close to (1986) and Baker et al. (1997) equilibrium (SIcalcite = 0.2 to 0.3; Figure 8a). The

Table II. Trends in drip discharge dynamics

Drip Date range Slope [ml day2] R2 p

TC1 20 June 2012 29 January 2013 16.3 0.83 <0.001 CP1 20 Jun 2012 19 December 2012 23.7 0.68 <0.001 CP2 20 June 2012 19 December 2012 3.3 0.38 0.025 CP3 10 May 2012 19 December 2012 12.1 0.85 <0.000 RD 25 April 2012 19 December 2012 6.5 0.84 <0.001

Table III. Pearson’s correlations of trace element ratios with rainfalls, external temperatures, and drip discharges

Correlation coefﬁcient TC1 n = 26 CP1 n = 26 CP2 n = 26 TC2 n = 17 CP3 n =23 ZD n =8 r (Mg/Ca × 1000 ratio vs. rainfall) 0.26 0.02 0.21 0.18 0.01 0.47 r (Sr/Ca × 1000 ratio vs. rainfall) 0.26 0.18 0.01 0.21 0.30 0.35 r (Mg/Ca × 1000 ratio vs. temperature) 0.78 0.18 0.29 0.02 0.02 0.30 r (Sr/Ca × 1000 ratio vs. temperature) 0.33 0.26 0.29 0.03 0.02 0.83 r (Mg/Ca × 1000 ratio vs. drip discharge) 0.36 0.37 0.11 0.40 0.45 0.72 r (Sr/Ca × 1000 ratio vs. drip discharge) 0.00 0.27 0.28 0.33 0.15 0.20 n – number of samples Rainfalls – cumulative rainfalls for 14-day period before drip sampling Temperature – mean external temperature per 14-day period before drip sampling Highlighted correlation is signiﬁcant at α = 0.05

Table IV. Pearson’s correlations of trace elements with each other

Correlation coefﬁcient TC1 n = 26 CP1 n = 26 CP2 n = 26 TC2 n = 17 CP3 n =23 ZD n =8

Sr versus Mg 0.40 0.21 0.49 0.06 0.13 0.17 n – number of samples Highlighted correlations are signiﬁcant at α = 0.05

Figure 8. Drip water saturation indices with respect to calcite: main data (a); archive data (Moravian Karst) (b)

18 archive data show SIcalcite in the range of 0.65–1.1 for all In the anomalous drip TC1, the values of δ O are similar the drips except the drip AC2 ranging from 0.4 to 0.8 to other drips, indicating the lack of evaporation in soil (Figure 8b). zone. In contrast, there is a signiﬁcant difference in enriched δ13C values about 3‰ in the anomalous drip Stable isotopes (Figure 5). The plot δ13C versus alkalinity (Figure 5) There is a small but noticeable difference in δ18O shows similar isotopic composition for majority of values between spring and fall samples, where the fall dripwaters. However, the drip TC1 differs from these samples are slightly enriched by about 0.2‰ (Figure 4). regular drips.

DISCUSSION Hydrogeochemistry and isotopes Hydrology Based on hydrogeochemical properties, the studied dripwaters were divided into two groups. The major Both the anomalous drip TC1 and regular drip CP2 group, including TC2, CP1, CP2, CP3, and ZD drips, (Figure 7) show a seepage flow that results from the further referred to as the regular drips, is characterized by slow movement of water through the epikarst aquifers the higher values of both the EC (Figure 3a) and SI of a small-scale primary or secondary permeability calcite fi (Figure 8a), and by the lower values of both the Mg/Ca (Baldini et al., 2006). Other drips are classi ed as (Figure 3c) and Sr/Ca ratios (Figure 3e). All the drips seasonal drips, but two drips (CP3, CP1) are quite show similar δ2H and δ18O values with no signs of fl close to the seepage ow. Therefore, as all the drips evaporation and slight isotopic enrichment in autumn show low/moderate discharges, they can be further samples (Figure 4). Values of δ13C (Figure 5) are divided into two classes based on their variability: (i) consistent with dissolution of calcite under closed system the drips of medium variability and (ii) the drips of conditions (Clark and Fritz, 1997). high variability (Figure 7). It is obvious that this study Only the drip TC1 differs: it represents a special one- fi does not con rm the conclusion of Baldini et al. member group referred to as an anomalous drip. This drip (2006) that faster drips tend to exhibit more variable shows systematically lower EC (Figure 3a), lower SIcalcite discharge. (Figure 8a), and enhanced both the Mg/Ca (Figure 3c) The decreasing discharges during the substantial part of and Sr/Ca (Figure 3e) ratios in comparison with the the season (Figure 2 and Table II) show that the relevant regular drips. The enhanced ratios are primarily given by perched aquifers in epikarst feeding the drips are low Ca concentrations. Whereas Mg and Sr concentra- insufficiently replenished during the period from May to tions in the anomalous drip TC1 are roughly consistent December 2012 and are gradually emptied. Similar with those in the regular drips, the Ca concentrations in behaviour was already described by Baker et al. (1997). the anomalous drip are markedly lower. Values of δ2H An exception is the drip CP2, the perched aquifer of and δ18O are similar to those of regular drips (Figure 4), which is replenished during the entire year (Figure 2e and but there is a strong enrichment in δ13C (Figure 5), Table II). No drips show a direct response to summer probably as result of a water degassing along the storm events, probably because of increased evapotrans- flow path. piration (Figure 6b). Even though water infiltration has Archive data show properties close to that of the increased from mid-September 2012 (see the line A in regular drips in the Punkva caves. On the contrary, they Figure 6), no significant response in drip discharges was show wider variability with particular abnormalities. All observed until mid-December 2012 (Figures 2 and 6c). the drips were supersaturated with respect to calcite, albeit From the time, the slowly increasing discharge is their values of SI varied in a wider range, SIcalcite = 0.4– observed for drips CP1 and ZD (Figure 6c). The increase 1.1. The drip AC1 shows some anomalous properties, is linked to the beginning of the period #5 (see the line B especially high Mg/Ca ratio (Figure 3d). Nevertheless, in Figure 6). A stronger response to winter precipitation this drip differs significantly from the drip TC1: its and an increase in drip discharge is observed during the electrical conductivity (Figure 3b) and SIcalcite (Figure 8b) period #5 for TC1 and TC2 drip sites (see the line C in values are consistent with the regular drips. This indicates Figure 6). Also, for the drip ZD, there is an obvious that the water of the drip AC1 is modified by dissolution enhanced discharge at the same time (Figure 6c). of minerals rich in Mg, possibly clay minerals. Another Discharges of the drips CP1, TC1, TC2, and ZD increased abnormality was observed in the drip BC: its low EC during the winter–spring period (from half December to (Figure 3b) is comparable with the EC value of the March) when maximum external temperature exceeded anomalous drip TC1 (Figure 3a). However, the drip BC is freezing point (Figure 2). It indicates that the drips are much more supersaturated with respect to calcite predominantly fed from snowmelt. The drips CP2 and (Figure 8b) and shows lower Mg/Ca ratio (Figure 3d). CP3 are an exception. Their perched aquifers might be Therefore, BC is possibly regular drip with a shorter somehow isolated and replenished only under specific residence time resulting in lower overall mineralization. conditions (Perrin et al., 2003; Tooth and Fairchild, Indeed, the BC discharge is quite variable, even though 2003). On the other hand, a time inconsistency is obvious extremely low (0.5 to 4 drops min1). Some abnormality between water infiltration (Figure 6b) and drip discharges can also be observed in the drip AC2 that shows the (Figure 6c) corresponding to the delay of 3–4 months. lowest SIcalcite from all regular drips (Figure 8b). This This indicates complex paths of water from the surface could be a result of less intense CO2 degassing. into perched aquifers, e.g. water is diverted by other The δ18O value of cave drip waters is a function of the predominant paths outside the aquifers during a substan- seasonality of recharge and modification within the soil tial part of season. and epikarst. Typically, drip water δ18O variability is

18 1.5 attenuated relative to precipitation δ O because of calcite at PCO2 ~10 . The water A may represent a mixing in the soil zone and epikarst (e.g. Mattey et al., water directly passing through limestone without a 2008 and the references therein). The δ18O value of drip contact with soils (in sites without vegetation/soil cover), waters may be increased by evaporation in caves with low whereas the water B represents an ordinary water passing relative humidity or fast air circulation (see Lachniet, soil/epikarst. Modelling of both waters mixing shows that 2009, for background). The stable isotope delta values for the resulting mix is undersaturated with respect to calcite the regular drips are roughly consistent with those in drips (SIcalcite < 0) in a wide range of mixing ratio (Figure 9). of the caves of similar latitude and altitude (e.g. Spötl As observed, the mixing ratio A/B ~ 3/7 shows similar et al., 2005; Riechelmann et al., 2011). properties as the anomalous dripwater TC1 (pH ~ 7, Both the regular and anomalous dripwaters are clustered Ca~2.45×103 mol l1 , EC ~ 307 μScm1 , along the Global Meteoric Water Line (Figure 4), SIcalcite ~ 0.17). The slightly positive SIcalcite values of indicating fast infiltration without the impact of evapora- the anomalous dripwater TC1 might be a result of tion. The difference in δ18O values between summer and subsequent water degassing in cave (compare Figure 3 fall samples (Figure 4) suggests some contribution of the and 8). Nevertheless, the WM alone is not capable to infiltration originating from summer precipitation, and, explain the enhanced Mg/Ca and Sr/Ca ratios, even thus, incomplete mixing in the epikarst zone. In contrast, though Roberts et al. (1999) suggested a model in which the δ13C value of 7.85‰ and the alkalinity value of trace element variations in speleothems are a result of 2.15 meq l1 of the anomalous drip TC1 (Figure 5) are very hydrological mixing of waters having interacted with two different from the regular drip values (10.39 to 10.94‰ geochemically distinct source rocks, calcite and dolomite. 1 and alkalinity > 5 meq l ). The reason is most likely However, the extremely pure limestones of the Macocha fi degassing of CO2 before the water enters the cave (Grif ths formation are not the case. et al., 2010; Frisia et al., 2011). Based on all the former criterions, the prior calcite precipitation (PCP) in the zone above the drip site, as Conceptual models illustrated in Figure 11, seems to be a best explanation of A plausible explanation of low water mineralisation and low calcite saturation is an enhanced water dynamics in karst profile. It is based on the idea that rapidly infiltrating water does not have enough time to attain equilibrium with calcite in the epikarst zone (Figure 11). After reaching a cave, such dripwater should show (1) lower mineralization and saturation with respect to calcite, (2) lower Mg/Ca ratio (as result of low Mg concentration due to slow dolomite dissolution dynamics and shorter time of interaction), and (3) more variable discharge (as a result of a tight relation to rainfall infiltrating quickly into epikarst). Based on all these indicators, it is obvious that the anomalous drip TC1 cannot be the result of the former mechanism. Except a lower mineralization/saturation, the drip does not fit two remaining indicators: its variability is low (Figures 2g and 7) and its Mg/Ca and Sr/Ca ratios are high (Figure 3). The low variability is contradicting a rapid rainwater influx; it rather indicates water supply from a large epikarstic aquifer with long water residence time and only minor effects of individual rainfall events. An alternative explanation of some anomalous properties is water mixing (WM). It is well known that the mixing of two waters of different chemistry leads to the decrease of saturation index (Bögli, 1971, 1980; Appelo and Postma, 2005). Such WM would be possible in the fracture system of vadose zone above the drip site Figure 9. Mixing model of two waters. The waters A and B are result of (Figure 11). As an example, let us consider two water calcite equilibrating at logPCO2 ~ 3.5 and logPCO2 ~ 1.5, respectively. SIcalcite is saturation index for calcite; P(CO2)W is partial pressure of the types: the water A is equilibrated with calcite at gaseous CO that is at equilibrium with aqueous carbonate species. Based 3.5 2 PCO2 ~10 and the water B is at equilibrium with on the PHREEQC calculations

Copyright © 2015 John Wiley & Sons, Ltd. Hydrol. Process. 30, 1506–1520 (2016) 64 1516 P. PRACNÝ ET AL. the TC1 anomalous properties. In general, PCP requires from water is temperature dependent, while the Sr well ventilated voids (upper cave floors, various air partitioning into calcite is temperature independent pockets, shafts, etc.) that allow water to degas CO2 and (Roberts et al., 1999). Recently, it is believed that reach supersaturation with respect to carbonate minerals seasonal low water saturation induces CO2 degassing and (Fairchild et al., 2000; Tooth and Fairchild, 2003; calcite precipitation in vadose zone/epikarst (a type of Fairchild and McMillan, 2007; Meyer et al., 2014). The temporal PCP) is responsible for elevated δ13C, Mg/Ca water that has underwent the PCP would show lower and Sr/Ca ratios in both the dripwaters and subsequently overall mineralization/saturation (low EC, low SIcalcite, in speleothems (Fairchild and McMillan, 2007). These lower alkalinity, and lower calcium concentrations) in parameters are therefore used rather as hydrological comparison with the regular drips. In addition, PCP indicators. In addition, some authors perceive the varying typically leads to enhanced Mg/Ca and Sr/Ca ratios trace element concentrations as indicators of changing 13 (Figure 3) and enriched δ C (DIC) values (Figure 5). elemental sources (e.g. Ayalon et al., 1999). This is the reason why we tend to explain the anomalous Despite the pure and low-Mg limestones of the properties of TC1 drip water via the PCP process, even Macocha formation (Lažánky and Vilémovice Lime- though some of its features could also result from water stones), the infiltrating water feeding the drips in the mixing. The PCP in the dripwater TC1 was also Punkva caves is enriched in trace elements. The resulting supported by speleological exploration: above the drip Mg/Ca and Sr/Ca ratios of the regular drips reflect an site, there is a crevice over 20 m high with rich impact of (i) meteoric precipitation, (ii) soil zone clays, speleothem growth (Glozar, 1984). and (iii) host rock carbonates (Figure 3c and e). In case of the anomalous drip TC1, however, the Mg/Ca and Sr/Ca ratios are more than two and half times larger (Figure 3c IMPLICATIONS FOR CLIMATIC PROXIES and e). The partitioning of trace elements between water and speleothem is quantified by a partition coefficient. In recent years, speleothems have been used for the For a species X, it is defined by the equation reconstruction of past terrestrial climate changes over a variety of time scales (see, e.g. Fairchild et al., 2006, for a = = ðÞX Ca c ¼ KxðÞX Ca w (1) review). Stable isotope ratios (δ18O and δ13C) have been widely used as proxies for palaeorainfall (e.g. Drysdale where Kx is partition coefficient, (X/Ca)w and (X/Ca)c are et al., 2005) and less frequently as proxies for the molar ratios in water and calcite, respectively. The palaeotemperature (e.g. Gascoyne, 1992). The using of value of Kx could be dependent on temperature, precipi- speleothem trace element ratios (e.g. Mg/Ca and Sr/Ca) as tation rate, lattice siting of X, water chemistry, etc. The paleoclimate indicators has been developed during the temperature-dependent KMg values are 0.019 and 0.012 for last two decades (Roberts et al., 1998, 1999; Finch et al., T = 15 °C and 6.6 °C, respectively. The KSr values range 2001). The presented study shows how the flow path between 0.057 and 0.078, depending on the calcite growth connected with PCP or/and WM can control dripwater rate (Huang and Fairchild, 2001). Equation 1 shows how parameters regardless of external climatic conditions. the anomalous Mg/Ca and Sr/Ca ratios of the dripwater The anomalous drip water properties such as TC1 would transfer into a potentially formed speleothem. All partition coefficients are lower than unity, which means mineralization/SIcalcite, enhanced Mg/Ca ratio, Sr/Ca ratio, and δ13C value would transfer into potentially that only a fraction of a given trace element enters the formed calcite speleothems. The using of such calcite during its precipitation, and the Mg/Ca and Sr/Ca speleothems for paleoclimate/paleoenvironmental recon- ratios increase in the residual solution. It is consistent with structions would lead to false conclusions. the enhanced ratios in the water subjected to the PCP. The higher partition coefficients for Sr indicate that Sr enters more readily into growing calcite than Mg. At low Trace elements speleothem growth rates, Sr may substitute for Ca in the As stated previously, the Mg/Ca and Sr/Ca ratios of calcite lattice (Boch, 2008; Fairchild et al., 2001) because speleothems are widely used for palaeoclimate recon- of the same valence and similar ionic radius of Ca2+ structions (see Fairchild and Treble, 2009 and the (1.08 Å) and Sr2+ (1.44 Å) (Gabitov and Watson, 2006; references therein). In the past, there were attempts Huang and Fairchild, 2001). At higher rates, Sr occupies to interpret Mg variations for reconstructing palaeo- interstitial and defect lattice sites, the abundance of which temperatures (e.g. Gascoyne, 1983) as the Mg partition increases with growth rate (Boch, 2008). Therefore, fast coefficient into calcite depends on temperature and not on growth rates may increase the Sr/Ca ratios in speleothems growth rate (e.g. Huang and Fairchild, 2001). The Mg/Sr (Huang and Fairchild, 2001; Treble et al., 2003). Model of ratio in speleothem calcite seems to be a better drip waters evolution during calcite precipitation is given palaeothermometer as the Mg partitioning into calcite in Figure 10. It is based on Equation 1 and partition

Figure 10. Evolution of trace element ratios during calcite precipitation. Evolution of Mg/Ca ratio (a): initial aqueous concentrations are [Ca] 3 1 5 1 =4×10 mol l and [Mg] = 6 × 10 mol l ; partition coefficient KMg = 0.014. Evolution of Sr/Ca ratio (b): initial aqueous concentrations are [Ca] 3 1 6 1 =4×10 mol l and [Sr] = 1.2 × 10 mol l ; partition coefficient KSr = 0.05 coefficients KMg = 0.014 and KSr = 0.05. As observed, the dating. These findings can be interpreted in terms of data roughly follow the theoretical curves, and the glacial or interglacial periods (Niggemann et al., 2003). anomalous drips are clearly separated from the regular Therefore, the speleothems as a product of the anomalous drips. drips of low saturation would be false proxies for climate 18 The correlations of the 1000 × (Mg/Ca)w ratio with the and environmental reconstruction. However, as δ O and drip discharges (Table III) are unconvincing and do not δ2D are similar for both the regular and anomalous drips confirm the negative correlation found in the Ernesto cave (Figure 4), they could be used as a test of the applicability by Fairchild et al. (2000). Thus, the presumed relation of overgrowths as a proxy. between 1000 × (Mg/Ca)w ratio with dry/wet changes (as it results from temporally PCP with a longer residual time Stable isotopes during dry seasons) is questioned. Similarly, expected In the study, the values of δ18O are similar for both covariance of the Mg and Sr ratios was not confirmed fi anomalous and regular drips (Figure 4), which could not (Table IV). The signi cant positive correlation for TC1 is induce any discrepancies in speleothems. However, the a consequence of the spatially controlled PCP in the enriched δ13C values in the anomalous drip would result anomalous drip. in even more enriched value in precipitated calcite. According to Dreybrodt and Scholz (2011), the δ13C Dripwater saturation value of drip water is determined mostly by the soil CO2 In principle, supersaturation with respect to carbonates composition. The exchange with the cave air has (to calcite in case of calcite speleothems) is a thermody- exchange time Tex about 3000 s, but calcite precipitation ’ namic force driving speleothems growth: the higher has time constant τp of several 100 s. When drip interval supersaturation, the higher growth rate. As supersatura- is < 0.1τp, the precipitated calcite in stalagmite reflects the tion increases, crystal habits are expected to change from isotopic composition of drip water. In such case, the prismatic to spherulitic (Zhang and Nancollas, 2008). The calcite isotopic composition may provide information calcite fabrics have been discussed in many recent works about paleoclimate conditions (Genty et al., 2003; Cruz (e.g. Frisia et al., 2000 or Riechelmann et al., 2014). et al., 2006). Changes in calcite fabrics and crystal habits have the In principal, the δ13C data in speleothems are used as potential to provide a record of supersaturation and an indicator of CO2 biological activity in soils. Under changes in water availability through time (e.g. Genty and warmer climate conditions, the fingerprint of isotopically Quinif, 1996). For example, the periods of growth activity depleted organic matter increases and dominates the and stagnation may be recognized/identified based on the isotopically enriched fingerprint of soil carbonates, with speleothem calcite individual overgrowths (calcite fab- resulting depleted δ13C values (Genty et al., 2003; rics, layers, and lamina) and compared with absolute Kaufmann and Dreybrodt, 2004). Such significant

Copyright © 2015 John Wiley & Sons, Ltd. Hydrol. Process. 30, 1506–1520 (2016) 66 1518 P. PRACNÝ ET AL. variations in the content of isotopically depleted organic CONCLUSIONS carbon in the source area of different drips in the Punkva An anomalous drip was studied in the Punkva caves and caves seem hardly possible. It is much more likely a compared with the regular drips in the same cave and consequence of degassing followed by calcite precipita- other caves in the region (Moravian Karst). The study tion (Fairchild et al., 2006). showed that properties of the anomalous drip substantially differ from the regular drips especially by lower Spatial versus temporal control on PCP and WM saturation with respect to calcite, lower mineralization, The permanent impacts of PCP and/or WM on the enhanced Mg/Ca and Sr/Ca ratios, and increased δ13C anomalous dripwater TC1 properties indicate that these values. In other properties, e.g. in the mean drip discharge phenomena are controlled spatially (Figure 11). Based on with low variability, the anomalous drip does not differ this finding, it may be deduced that spatial conditions are from the regular drips. The results of the study suggest an important factor in addition to the temporal conditions, that the anomalous properties are a result of prior calcite e.g. dry/wet seasons, as some authors have considered precipitation or/and water mixing in vadose zone above (Fairchild et al., 2000; Verheyden et al., 2003; Cruz et al., the cave. Both processes would undoubtedly affect 2007). The effect of spatial PCP and WM is hardly environmental proxies in formed speleothems. As the removable from the climatic signal. Despite a currently mentioned processes are controlled mostly spatially, the spatial character, the PCP/WM may appear and disappear study indicates that the knowledge of the reaction- in long-scale periods in dependence on the changes of transport path of water feeding the drips is necessary conditions. In case of the PCP, the changes could be before the proxies of relevant speleothems are used for linked to (1) an extinction/formation of the voids paleoclimatic reconstructions. necessary for degassing or (2) changes in the ventilation regime by e.g. closing/opening old/new air pathways. The void extinction may proceed via deposition of calcite or ACKNOWLEDGEMENTS clay sediments. In turn, new voids may be formed by fi The work was supported by a donation from GEOtest, a.s., karsti cation or tectonic processes. Similarly, the spatial š WM may evolve over a long-term period by changing Brno, Czech Republic. We thank Franti ek Buzek from the water flow path caused by tectonics, sedimentation, Czech Geological Survey for stable isotopes analyses and karstification, and other processes. two anonymous reviewers for their helpful comments.

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This appendix presents following research paper:

Pracný, P., Faimon, J., Kabelka, L., & Hebelka, J. (2016). Variations of carbon dioxide in the air and dripwaters of Punkva Caves (Moravian Karst, Czech Republic). Carbonates and Evaporites, 31(4), 375–386. http://doi.org/10.1007/s13146-015-0259-0

70 Carbonates Evaporites (2016) 31:375–386 DOI 10.1007/s13146-015-0259-0

ORIGINAL ARTICLE

Variations of carbon dioxide in the air and dripwaters of Punkva Caves (Moravian Karst, Czech Republic)

Pavel Pracny´1 • Jirˇ´ı Faimon1 • Ludvı´k Kabelka2 • Jirˇ´ı Hebelka3

Accepted: 14 July 2015 / Published online: 19 August 2015 Ó Springer-Verlag Berlin Heidelberg 2015

Abstract Carbon dioxide (CO2) was studied in Punkva CO2 degassing and calcite precipitation is demonstrated in Caves in the Moravian Karst (Czech Republic) during a detail in a geochemical model. The study presents new data one-year period from February 2012 to March 2013. Partial indicating that the CO2 source might be deployed in deeper pressures of the CO2 corresponding to aqueous carbonates, parts of karst profile (epikarst) in addition to karst soils. -2.91 -2.35 PCO2ðWÞ (10 –10 , i.e., 0.12–0.45 vol%), and those Keywords Carbon dioxide (CO ) Cave Degassing participating in the initial dripwater formation, PCO2ðHÞ 2 (10-1.77–10-1.49, i.e., 1.7–3.2 vol%), were calculated from Dripwater Model Periodicity dripwater hydrogeochemistry, and compared with the -3.31 -2.49 partial pressure in cave air, PCO2ðairÞ (10 –10 , i.e., Introduction 0.06–0.32 vol%). Both the PCO2ðairÞ and PCO2ðWÞ showed clear seasonal variations with maxima in summer and Carbonate rocks cover large parts of Earth’s surface. minima in winter. In contrast, the PCO ðHÞ was very stable 2 Besides their essential significance as a water reservoir or without any significant seasonality, which could indicate its biotope, the carbonate karst landscapes provide wide range independence on surface conditions. As an exception, one of scientific information. Karsology, speleology, geo- anomalous drip with significantly lower and varying chemistry, sedimentology, paleontology, hydrology, and P , P , and SI was recognized as a result of CO2ðWÞ CO2ðHÞ calcite various environmental sciences benefit from research in the prior calcite precipitation. Evolution of dripwater during areas of carbonate karst sediment. With growing interest in climate, change the importance of paleoclimate reconstructions based on cave speleothems has increased. Car- & Pavel Pracny´ bon dioxide (CO2) plays a crucial role in wide range of [email protected] carbonate karst processes. It is believed that most of the

Jirˇ´ı Faimon karst CO2 is produced by biodegradation of organic matter [email protected] or root respiration in soil/epikarst (Kuzyakov 2006). The

Ludvı´k Kabelka CO2 concentrations in karst soils usually reach up to [email protected] * -2 1 vol% (PCO2 10 ) with strong seasonal fluctuations Jirˇ´ı Hebelka leading to highest values in summer and lowest in winter [email protected] (Faimon and Licˇbinska´ 2010; Sanchez-Can˜ete et al. 2011; Plestenjak et al. 2012; Blecha and Faimon 2014a). Part of 1 Department of Geological Sciences, Faculty of Science, Masaryk University, Kotla´rˇska´ 2, 611 37 Brno, the soil CO2 is dissolved in percolating water and trans- Czech Republic ported downwards into lower parts of karst profile. As the 2 GEOtest, a.s., Sˇmahova 1244/112, 627 00 Brno, water flows through the carbonate rock, the dissolved CO2 Czech Republic controls carbonate dissolution (Tooth and Fairchild 2003). 3 Cave Administration of the Czech Republic, Svitavska´ 11, The water entering cave as a dripwater gradually degasses 678 01 Blansko, Czech Republic (releases CO2 into the cave atmosphere). This process is 123

71 376 Carbonates Evaporites (2016) 31:375–386 driven by the difference between the partial pressure of corridor, 0.3–0.5 m above the ground level), and (3) the gaseous CO2 corresponding to aqueous carbonates Zadnı´ Chamber with the drip site ZD (drip falling from the

(PCO2ðWÞ) and the actual partial pressure of gaseous CO2 in height of ca. 10 m on stalagmite Vase (Fig. 1). The terrain above the cave system is a part of the the cave atmosphere (PCO2ðairÞ). The water–PCO2ðairÞ system exponentially approaches the partial equilibrium given by Macocha Plateau. It is circumscribed by the Macocha % Abyss from the northeast side and by Pusty´ Valley from the the condition PCO2ðWÞ PCO2ðairÞ. This evolution leads to 2- west side. The highest point has altitude of 491 m. The pH increase and to enhanced CO3 activity. In response, the water becomes supersaturated by calcite that precipi- surface is forested with dominant beeches and minor tates in the form of calcite speleothems. Although degas- spruces. The southeastern part is covered with an asphalt sing and precipitation run together, the second process is parking lot (see Fig. 1). The soil type is mull rendzic somewhat slower than the ﬁrst one, which causes near Leptosol (under deciduous trees) and gray rendzic Leptosol constant calcium concentration during the degassing (under coniferous trees). The soil layer thickness varies (Dreybrodt 2008). Even though soils are widely accepted from 30 to 80 cm (Faimon and Licˇbinska´ 2010). The cli- as a main source of carbon dioxide in karst systems, much mate in Moravian Karst is moderate with the long-term mean annual temperature about 10 °C with maxima in higher CO2 concentrations were measured in vadose zone (2–6 %, see Benavente et al. 2010). This contradiction is July/August and minima in January/February. The mean conﬁrmed by results of inverse modeling of karst water annual precipitation is about 700 mm, almost half of which composition (Faimon et al. 2012; Peyraube et al. 2012, falls from June to September. 2013; Milanolo and Gabrovsˇek 2015) that indicates similar

CO2 concentrations (3–10 vol%). The aim of this study is Methods to show the long-term variations in (i) cave air PCO2 , and

(ii) PCO2 deduced from dripwater chemistry based on one- year monitoring campaign in the Punkva Caves (Moravian Data were collected between February 2012 and March Karst, Czech Republic). 2013. The monitoring was realized twice per month with almost regular step (about 14 days). Totally, 126 samples Site of study of dripwater from 6 cave drip sites collected during 26 sampling events were studied. The study was conducted in the Punkva Caves (Moravian Directly in cave, immediately after sampling, the drip- Karst, Czech Republic, see, e.g., Absolon 1970; Blecha and water basic hydrogeochemical parameters were determined: ± Faimon 2014a). Moravian Karst is the largest karst region pH (WTW 330i; precision 0.005 pH), specific electrical ± in the Czech Republic with surface area of 92 km2. Punkva conductivity (Cond 3110; precision 0.5 %), alkalinity Caves system is developed in very pure Devonian lime- (acidimetric titration using 0.05 M HCl with potentiometric stone of Macocha Formation—specifically in the dark-gray pH measurement; calculation based on Gran’s function, mostly heavy-bedded Lazˇańky limestone (Givetian) and the Stumm and Morgan 1996) and calcium concentration bright-gray to gray, very pure, massive to heavy-bedded (complexometric microtitration using calcein as an inner Vile´movice limestone (Frasnian). The thickness of over- indicator). In addition, the PCO2 levels in cave air were laying limestone reaches up to 120 meters. Genetically, the measured using the device Almemo 2594-4S linked with the cave system was formed by an underground flow of the sensors FYA600CO2 or FYA600–CO2H (CO2) (all devices ± river Punkva. The caves are structured into a dry upper by Ahlborn; precision 2 %). Drip rate was measured by cave part and the lower cave part through which currently counting drips during a given time period. flows the river. The river outflow (altitude of 350 m) is The hydrometeorological data come from the meteoro- situated near the entrance into the dry part (altitude of logical station Macocha situated above the caves 360 m). The Punkva Caves are one of the most popular (150–200 m eastward). The station is used by the Czech tourist destinations in the Czech Republic with as much as Hydrometeorological Institute and the Cave Administration 200 thousand visitors per year. of the Czech Republic. Details on the study cave sites are as follows: (1) the Saturation indices were calculated using PHREEQC Tunnel Corridor with the drip sites TC1 (drapery on a modeling software (Parkhurst and Appelo 1999). The same P P chimney edge at the ceiling in the central part of the corri- PC code was used to calculate CO2ðWÞ and CO2ðÞH (hy- dor) and TC2 (a small straw-stalactite on the corridor ceil- pothetical PCO2 participating in the initial water formation) ing), both dripping from the height of 4–5 m, (2) the corridor from hydrogeochemistry (Faimon et al. 2012). The Statis- behind the Prˇednı´ Chamber with the drip sites CP1, CP2, and tica software package was used for statistical data analysis CP3 (small straw-stalactites on a slanting wall of the (Statsoft 2014).

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Fig. 1 Sketch map of the Punkva Caves with the studied drip sites, meteorological station (MS) and surface features (a); a simpliﬁed cave proﬁle based on Hromas (2009)(b) and map with localization of the site in the Czech Republic (c)

Results 8.42 ± 0.16 °C (the site CP) and 9.05 ± 0.11 °C (the site ZD). Climate and cave microclimate The values of partial pressure of carbon dioxide in cave

atmosphere, PCO2ðairÞ, varied with temperature: The lower - \ Mean daily external temperatures ranged between 9.3 PCO2 values correspond to the conditions at T(max) MAT, ° and 25.9 C (Fig. 2a). Mean annual temperature (MAT) whereas the enhanced PCO2 are linked to the conditions at ° [ was 7.89 C during the entire monitored period. There T(min) MAT (Fig. 2b). PCO2ðairÞ in Prˇednı´ Corridor (all were 141 days when external temperature maxima were the CP sites) ranged from 10-3.31 to 10-2.63 (0.23– below MAT (the blue areas in Fig. 2) and 105 days when 0.05 vol%; 10-3.05 ± 0.08 conﬁdence interval at a = 0.05). the external temperature minima were above MAT (the The CO2 partial pressure in Tunnel Corridor (both the TC yellow areas in Fig. 2). The temperature in cave ranged * -3.31 sites) was virtually similar, PCO2ðairÞ from 10 to from 7 °C (winter) to 10 °C (summer). The mean tem- 10-2.63 (0.23–0.05 vol%; 10-3.05 ± 0.07 conﬁdence inter- peratures in cave were 7.64 ± 0.15 °C (the site TC), val). The CO2 partial pressure in Zadnı´ Chamber (ZD site)

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than 1 to 174 drops min-1 (*8.5–1450 mL h-1) (see Table 1). The low rates occurred from November to February, whereas enhanced rates were observed from February to May. The variation coefﬁcients were from 43.2 % for TC1 and 59 % for CP3 to 87.1 % for CP1. The drips TC2 and ZD were very unstable: the periods of standard activity were alternated by the periods, during which the dripping was either extremely slow (\1 drop per 5 min) or zero. On the contrary, the drip CP2 shows very stable drip rate with variation coefﬁcient 17.6 %. The lowest total discharge calculated for the monitored period was 403.6 L for drip TC2, while the highest was 2372.1 L for the drip CP1.

Hydrogeochemistry

Hydrogeochemical properties of dripwaters show only minor variations over time. The highest average concentration of calcium and alkalinity, [Ca] *3.5 mmol L-1 and [Alk] *6.23 meq L-1, respectively, was found in the drip CP1 and the drips CP2 and CP3 show very close values (Table 1). The drip TC1 shows chemical composition with Fig. 2 External and cave air temperatures (a) and cave air CO2 in the lowest concentrations of calcium, [Ca] = 1.47 ± 0.05 TC, CP, and ZD cave sites (b). MAT represents external mean annual -1 ± -1 temperature. The blue and yellow areas mark the conditions, under mmol L , and alkalinity, [Alk] = 2.14 0.1 meq L a = which T(max) \ MAT and T(min) [ MAT, respectively (conﬁdence interval at 0.05). Besides the carbonates and calcium, the electrical neu- trality of the solution is balanced by other anions (esp. - - ? ? ? - - 2 2 ranged from 10 3.22 to 10 2.49 (0.32–0.06 vol%; SO4 and Cl ) and cations (esp. Mg ,K and Na )as 10-2.86±0.09 conﬁdence interval). illustrated in Table 1. Complete chemical composition of

the solution was used to calculate both SIcalcite and PCO2ðWÞ

Hydrology and PCO2ðHÞ, respectively. The drips monitored in the Prˇednı´ Corridor (CP1, 2, 3) Overall precipitation during the monitored period was were supersaturated with respect to calcite, as indicated by 641.7 mm. Daily precipitations reached the maximum at saturation index values in the range of SIcalcite = 0.76–1.28 -1 35.6 mm day . The drip rates of TC1, CP1 and CP3 (mean value SIcalcite = 1.04 ± 0.03, conﬁdence interval signiﬁcantly change during the year; they ranged from less at a = 0.05). Less supersaturated were the drips TC2

Table 1 Basic hydrogeochemical properties of monitored drips Site TC1 TC2 CP1 CP2 CP4 ZD n/nQ 26/26 17/26 26/26 26/26 23/26 8/23 Drip rate (min-1) 30.8 ± 5.12 10.2 ± 9.03 46.3 ± 15.5 17.96 ± 1.2 14.95 ± 3.4 25.6 ± 14.7 Annual volume (L) 1604.6 403.6 2372.1 1001.1 838.9 1133.5 EC (lScm-1) 297 ± 8.5 551.2 ± 2.9 628 ± 2.3 621.7 ± 2 616.4 ± 5.5 551.5 ± 1.9 pH 7.95 ± 0.06 7.99 ± 0.05 8.03 ± 0.05 8.10 ± 0.04 8.11 ± 0.05 8.09 ± 0.06 Ca (mmol L-1) 1.47 ± 0.05 3.07 ± 0.04 3.50 ± 0.02 3.48 ± 0.02 3.49 ± 0.03 3.05 ± 0.02 Sum of cations (meq L-1) 3.2 ± 0.27 6.35 ± 0.15 7.22 ± 0.08 7.18 ± 0.1 7.2 ± 0.14 0.226 ± 0.01 Alkalinity (meq L-1) 2.14 ± 0.10 5.33 ± 0.04 6.23 ± 0.07 6.12 ± 0.05 6.14 ± 0.06 5.67 ± 0.04 Sum of anions (meq L-1) 3.29 ± 0.3 6.44 ± 0.12 7.4 ± 0.2 7.29 ± 0.16 7.32 ± 0.16 6.47 ± 0.1

SI(calcite) 0.14 ± 0.04 0.83 ± 0.09 1.01 ± 0.05 1.06 ± 0.04 1.07 ± 0.05 0.92 ± 0.04 n number of analyzed samples, nQ number of drip rate measurements All the conﬁdence intervals calculated for a = 0.05

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= ± = ± = -1.57 ± 0.01 a = (SIcalcite 0.83 0.09) and ZD (SIcalcite 0.92 0.04). PCO2ðHÞ 10 (confidence interval at 0.05; In contrast, the drip TC1 in Tunnel corridor varied close to 2.7 vol%), were found for all the drips CP. Somewhat equilibrium with calcite, in the range of SI from = -1.66 ± 0.01 calcite lower but similarly stable values, PCO2ðHÞ 10 -1.75 ± 0.01 -0.20 to 0.32 (mean value SIcalcite = 0.14 ± 0.04). = (2.2 vol%) and PCO2ðHÞ 10 (1.8 vol%), were found for the drips ZD and TC2, respectively. In contrast, Carbon dioxide significantly lower values with somewhat higher variability = -2.95 ± 0.07 were found for the drip TC1, PCO2ðHÞ 10 Partial pressure of the gaseous CO2 corresponding to aque- (0.11 vol%). The calculated PCO ðWÞ and PCO ðHÞ values for ous CO (P ) was calculated from water hydrogeo- 2 2 2 CO2ðWÞ the drips CP1, 2, 3 and TC1 are presented in Fig. 3 and chemistry. It ranged from 10-2.91 to 10-2.35 (0.12– compared with the PCO ðairÞ values measured in cave 0.45 vol%) for the drips CP1, CP2, and CP3 2 atmosphere. The distribution of P values for the drips * -2.64±0.03 a = CO2 (PCO2ðWÞ 10 , confidence interval at 0.05), -2.8 -2.48 TC2 and ZD is not presented, as their time series are from 10 to 10 (0.16–0.33 vol%) for the drip TC2 incomplete due to the drip small discharges. * -2.75 ± 0.01 -2.87 -2.46 (PCO2ðWÞ 10 ), and from 10 to 10 P * -2.61 ± 0.04 (0.13–0.35 vol%) for the drip ZD ( CO2ðWÞ 10 ). Data analysis

The PCO2ðWÞ values for the drip TC1 were lowest: they - - ranged from 10 3.25 to 10 2.60 (0.06–0.25 vol%; Correlations * -2.95 ± 0.07 PCO2ðWÞ 10 ). The partial pressures potentially participating at given Correlations between the CO2 partial pressures in cave air, water formation, hypothetical partial pressures PCO2ðHÞ, external temperature and cave temperature are presented in were calculated from water chemistry as PCO2 , at which Table 2. In the upper part of the table, there are correla- degassed water would return back to equilibrium with tions of the variables in raw data set. As can be seen, all the calcite (Faimon et al. 2012). Very stable values, variables are positively correlated (r * 0.42–1.00,

Fig. 3 Actual partial pressure of CO2 in cave air, PCO2ðairÞ, potential partial pressure of gaseous CO2 that would be at equilibrium with aqueous CO2,

PCO2ðWÞ, and hypothetical partial pressure of CO2 at water formation, PCO2ðHÞ. The partial pressures are given for the sites TC1 (a), CP1 (b), CP2 (c), and CP3 (d)

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= Table 2 Correlations of cave air PCO2 and external daily temperatures (n 26) ° ° ° D ° ° Variable logPCO2ðairÞ (TC) logPCO2ðairÞ (CP) logPCO2ðairÞ (ZD) T(max) ( C) T(min) ( C) T(mean) ( C) T ( C) T(cave) ( C]

Raw data

logPCO2ðairÞ (TC) 1 0.97 0.89 0.58 0.63 0.60 0.60 0.46

logPCO2ðairÞ (CP) 0.97 1 0.91 0.59 0.60 0.60 0.60 0.48

logPCO2ðairÞ (ZD) 0.89 0.91 1 0.80 0.79 0.80 0.80 0.42

T(max) (°C) 0.58 0.59 0.80 1 0.92 0.96 0.96 0.46

T(min) (°C) 0.63 0.60 0.79 0.92 1 0.91 0.91 0.5

T(mean) (°C) 0.60 0.60 0.80 0.96 0.91 1 1.00 0.44 DT (°C) 0.60 0.60 0.80 0.96 0.91 1.00 1 0.44

T(cave) (°C) 0.46 0.48 0.42 0.46 0.5 0.44 0.44 1 Stationary data* 2 - - - logPCO2ðairÞ (TC) 1 0.84 0.27 0.46 0.34 0.40 0.40 0.23 - - - - logPCO2ðairÞ (CP) 0.84 1 0.48 0.25 0.40 0.32 0.32 0.19 - - - - - logPCO2ðairÞ (ZD) 0.27 0.48 1 0.17 0.23 0.24 0.24 0.3

T(max) (°C) -0.46 -0.25 -0.17 1 0.56 0.78 0.78 -0.18

T(min) (°C) -0.34 -0.40 -0.23 0.56 1 0.49 0.49 0.02

T(mean) (°C) -0.40 -0.32 -0.24 0.78 0.49 1 1.00 -0.17 DT (°C) -0.40 -0.32 -0.24 0.78 0.49 1.00 1 -0.17

T(cave) (°C) 0.23 0.19 -0.3 -0.18 0.02 -0.17 -0.17 1

DT = T(mean) - TMAT * Data detrended by polynomial of 6th order The highlight correlations are signiﬁcant at a = 0.05

\ p 0.05). To eliminate potential effect of trends, the raw between PCO2ðairÞ and PCO2ðHÞ, and between PCO2ðWÞ and data were transformed into stationary data set: Nonlinear PCO2ðHÞ. In case of stationary data (detrended by the trends were determined by regression with a polynomial polynomial of 6th order, see the right side of Table 3), function of sixth order (P6). The trend was subtracted from most of the correlations are insignificant at a = 0.05. An the raw data set. The obtained residues were used as a exception is the positive correlation between PCO2ðHÞ and stationary data set. = PCO2ðairÞ in TC1 site (r 0.43). Correlations of the stationary data are in the lower part of Table 2. It is obvious that these correlations are less CO2 periodicity strong in comparison to the raw data. The correlations of CO2 with external temperature are negative, however, the Autocorrelation can help verify the presence of cycles and a = D statistically significant at 0.05 are only T(mean), T, determine their duration. The autocorrelation function and T(max) versus CO2 on the TC site and T(min) versus CO2 measures the correlation of a signal x(t) with itself shifted in the CP sites. Correlations between CO2 on the ZD site by a time delay called a lag time s. The maximum value of and external T are statistically insignificant. There is strong correlation coefficient will always be at a zero lag, since a correlation between the P on the CP site and P on CO2 CO2 signal is always perfectly correlated with itself. Other the TC site. The P on the ZD site correlates with P CO2 CO2 peaks in the autocorrelation indicate the periods, at which on the CP site. Cave temperature does not correlate with the signal quasi-repeats. In other words, autocorrelation is any other variables. based on the idea that a quasi-periodic signal will resemble Correlations between the partial pressures of all types of itself in the time domain when time shifted by duration CO (CO in cave air, P ,CO corresponding to 2 2 CO2ðairÞ 2 equals to the period. Before the analyses, the raw data (time P aqueous carbonates, CO2ðWÞ, and hypothetical CO2, series) were transformed to equidistant data with a step of

PCO2ðHÞ) are shown in Table 3. In case of raw data (see the 15.2 days by linear extrapolation. left side of Table 3), there were found positive and statis- Autocorrelation analysis of cave air partial pressures tically signiﬁcant correlations just between P and CO2ðairÞ shows a periodicity in PCO2ðairÞ with the period about = \ PCO2ðWÞ (r 0.48–0.79, p 0.05) for all the drip sites. For 304 days, as indicated by the half of period (HP) of the drip TC1, there were found also positive correlations 152 days between the lag 5 and 15 (lag *1 corresponds to

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Table 3 Correlations of all logPCO2ðairÞ logPCO2ðWÞ logPCO2ðHÞ logPCO2ðairÞ logPCO2ðWÞ logPCO2ðHÞ studied types of PCO2 TC1 (raw data) TC1 (stationary data, P6) - logPCO2ðairÞ 1 0.79 0.92 1 0.11 0.43 - - logPCO2ðWÞ 0.79 1 0.79 0.11 1 0.21 - logPCO2ðHÞ 0.92 0.79 1 0.43 0.21 1 CP1 (raw data) CP1 (stationary data, P6) - logPCO2ðairÞ 1 0.59 0.17 1 0.28 0.02

logPCO2ðWÞ 0.59 1 0.13 0.28 1 0.13 - logPCO2ðHÞ 0.17 0.13 1 0.02 0.13 1 CP2 (raw data) CP2 (stationary data, P6) - logPCO2ðairÞ 1 0.48 0.31 1 0.30 0.31 - logPCO2ðWÞ 0.48 1 0.29 0.30 1 0.13 - - logPCO2ðHÞ 0.31 0.29 1 0.31 0.13 1 CP3 (raw data) CP3 (stationary data, P6) - logPCO2ðairÞ 1 0.54 0.11 1 0.11 0.18

logPCO2ðWÞ 0.54 1 0.36 0.11 1 0.11 - logPCO2ðHÞ 0.11 0.36 1 0.18 0.11 1 The statistically signiﬁcant correlation (a = 0.05) is highlighted P6—data detrended by polynomial of 6th order n = 26 for TC1, CP1, CP2 and n = 23 for CP3

15.2 days) with negative correlations (Fig. 4a, d, g, j). The The term C in Eq. (1) represents relationship same period of 304 days is observed at P (Fig. 4b) 2 3 CO2ðWÞ 2 4 K c and PCO2ðHÞ (Fig. 4c) for dripwater TC1. In contrast, the 6 2 HCO 7 4 3 5; C ¼ 2 ð2Þ PCO2ðWÞ of the drip CP1 (Fig. 4e) shows the HP of c Kc KH K1 ðÞCa2þ 122 days. Autocorrelation for drips CP2 (Fig. 4h) and CP3 (Fig. 4k) is more inconclusive with HP of 136 days and where Kc is the solubility product of calcite, K1 and K2 are 182.4 days, respectively. Besides, no PCO ðHÞ periodicity is 2 the carbonate dissociation equilibrium constants, KH is the visible for the drips CP1, CP2 and CP3 (Fig. 4f, i, l). equilibrium constant between CO2(g) and CO2(aq), and the Results of the autocorrelation analysis are summarized in c symbols are the activity coefﬁcients for respective ions. Table 4. An useful visualization of Eq. (1) is a plot in which the Autocorrelation of mean daily temperatures in studied SIcalcite is the y-axis coordinate and the log PCO2ðWÞ is the x- period (Feb 2012–Mar 2013) shows HP of 180 days. axis coordinate (Fig. 5). Let us assume that the point A in this model represents an initial solution formed by calcite Hydrogeochemical model of CO2 degassing and calcite =- dissolution under log PCO2 1.5 (3.2 vol%) and tem- precipitation perature T = 10 °C into equilibrium in soil/epikarst.

In a cave, this water degasses CO2 and evolves along the Based on the chemical composition of dripwater in cave, it degassing line with the slope of * -1, based on Eq. (1). is possible to estimate the CO2 concentration participating The linearity requires pure degassing, i.e., the invariant in water formation during limestone dissolution in term 3log(Ca2?) ? logC. The more realistic degassing line epikarst/vadose zone above cave systems. It is represented is slightly curved as illustrated by a line calculated using by the so-called hypothetical CO partial pressure, P . 2 CO2ðHÞ the PHREEQC code (Fig. 5). This nonlinearity is caused Using an inverse model, the P may be reconstructed CO2ðHÞ by evolution of the term logC due to a change in activity as the PCO2 , at which given degassed dripwater returns coefﬁcients. Let us assume that this water is fully degassed back to equilibrium with calcite. The simpliﬁed relation = at the point B, when PCO2ðWÞ PCO2ðairÞ. This water is between SI , P , and Ca2? activity (Peyraube calcite CO2ðWÞ strongly supersaturated with respect to calcite (SI *1.5). et al. 2012; Milanolo and Gabrovsˇek 2015) is: During the calcite precipitation, it would evolve along the 2þ SIcalcite ¼logðÞþPCO2 3 log Ca þ log C ð1Þ line perpendicularly toward the point C in an open system

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Fig. 4 Autocorrelation of PCO2 in cave air (a, d, g, j), PCO2 corresponding to aqueous carbonates (b, e, h, k), and PCO2 hypothetical (c, f, i, l) 123

78 Carbonates Evaporites (2016) 31:375–386 383

Table 4 Overview of the P P P autocorrelation analysis results CO2ðairÞ CO2ðWÞ CO2ðHÞ Site HP (days) Autocorrel. HP (days) Autocorrel. HP (days) Autocorrel.

TC1 152 S 152 S 152 S CP1 152 S 122 S NP IS CP2 152 S 136 IC NP IS CP3 152 S 182 IC NP IS HP half of period, NP no period S signiﬁcant at a = 0.05 IC inconclusive IS insigniﬁcant at a = 0.05

Fig. 5 Saturation indices of dripwaters in Punkva Caves as function of log PCO2ðWÞ. At the intersection of degassing line with x-axis, the values of

PCO2ðHÞ can be read. The points A, B, C, D show important evolution state of system; see the text for more detailed information. The measured range of PCO2 in cave air in winter/summer season is highlighted

= (under conditions that PCO2ðairÞ PCO2ðWÞ) or toward the Discussion point D in a closed system (when PCO2ðWÞ is not invariant), as showed by a simulation calculated using the PHREEQC Cave carbon dioxide seasonality code (Fig. 5). This model can be used to reconstruct the conditions (partial pressures), under which was karst water The partial pressure of CO2 in cave air, PCO2ðairÞ, shows formed. A degassing line with slope -1 is drawn over the clear seasonality with maximum in summer and minimum real data projected into the plot. The intersection of the line in winter (Fig. 3). The period calculated from the half of with the x-axis corresponds to PCO2ðHÞ. All data excluding period (Fig. 4a, d, g, j; Table 4) is about 304 days—which is less than 1 year. It is probable that the period is a result drip site TC1 head to the logPCO2ðHÞ values in the range from -1.8 to -1.5 (1.6–3.2 vol%). However, this estimate of cave ventilation driven by difference between exterior is valid only under the condition that no calcite was pre- and interior temperature. Because the mean annual tem- cipitated, i.e., that degassing was faster than calcite pre- perature in exterior (MAT) roughly represents the mean cipitation. If the calcite is precipitated during the annual cave temperature (see, e.g., Domıńguez-Villar et al. degassing, the evolution path is shifted below the degassing 2013), the daily temperature maxima below MAT indicate line. As a result, the line projected over the data does not daily upward airflow ventilation mode (UAF mode), whereas the daily temperature minima above MAT indicate intersect with x-axis in the original PCO2ðHÞ value but at a lower value. daily downward airflow ventilation mode (DAF mode) in a

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cave (Fig. 2). Under assumption that the air residence time Table 3). This indicates that PCO2ðWÞ is connected with in cave is shorter than 24 h, those days represent the PCO2ðairÞ via degassing. periods of cave active ventilation (Faimon et al. 2012). However, a tight relationship of PCO2ðairÞ and PCO2ðWÞ Totally, there were 141 days when the cave persisted in was not confirmed by the stationary data analysis: no UAF ventilation mode and, probably, also in the winter correlations statistically significant at a = 0.05 were found active ventilation period, and 105 days when the cave (see the right side of the Table 3). The reason could be that persisted in DAF mode and in the summer active ventila- PCO2ðWÞ is additionally affected by the drip rate variations. tion period. The sum of the days is lesser than the period of If the drip rate is low, the drop hanging on the tip of spe- 304 days mentioned formerly. leothem can degas more extensively in comparison with As shown in Fig. 2, the ventilation modes very likely the drop that hanged on the tip for a shorter time. As seen control the cave air CO2. The systematically lower PCO2ðairÞ in case of the most stable drip CP2, the PCO2ðWÞ periodicity in UAF mode is a result of a transport of air from exterior is on the edge of statistical significance, but the period is through the lower cave entrances, whereas the higher close to PCO ðairÞ (Table 4). HP for drip CP1, which is a P in DAF mode is a result of a transport of CO 2 CO2ðairÞ 2 very variable drip with high discharge, is shorter than HP enhanced air from some upper openings, i.e., from vents for PCO ðairÞ, whereas HP for drip CP3, a variable drip with leading through epikarst and soils. Similarity in the trends/ 2 low discharge, is longer than HP for P (Table 4). In seasonality of cave CO and external temperatures is CO2ðairÞ 2 principle, dripwater degassing is also influenced by documented by the correlation of raw data (upper part of hydrochemistry, i.e., by the initial P at reaching the Table 2). Nevertheless, the stationary data do not confirm CO2ðWÞ cave. Another reason could be a short time shift between this relationship (see the lower part of Table 2): the cor- the variables. Unfortunately, this shift cannot be proved, relations between PCO and external temperature are neg- 2 e.g., by cross-correlation method, because of the long ative and just few of them (e.g., T vs. log P in (mean) CO2ðairÞ monitoring step. the TC site or T(min) vs. logPCO2ðairÞ in the CP site) are statistically significant. This contradiction could be result Hypothetical P in epikarst of the time shifts between the variables being shorter than CO2 the relatively long monitoring steps (about 15.2 day). In In contrast to P , the hypothetical partial pressure, addition, the internal temperatures in cave do not correlate CO2ðWÞ P , shows very stable values for all regular drips with P CO2ðHÞ CO2ðairÞ. during the whole monitored period (Fig. 3b, c, d). Besides, no seasonality is indicated by the autocorrelation analysis Dripwater hydrogeochemistry and P CO2ðWÞ (Table 4; Fig. 4f, i, l). In the caves of Moravian Karst, the hypothetical partial pressure of CO varies in the range of Based on hydrogeochemical properties, the studied drip- 2 1–11 vol% (Faimon et al. 2012). Other studies indicate waters were divided into two groups, which were discussed similarly high CO concentrations (3–10 vol%) (Peyraube separately. The bigger group comprises the drips CP1, 2 et al. 2012, 2013; Milanolo and Gabrovsˇek 2015). This is CP2, CP3, TC2 and ZD that are further referred to as the contradictory to low concentrations of soil CO with sea- regular drips. These drips show high Ca2? concentrations 2 sonal character (Faimon and Licˇbinska´ 2010; Sanchez- and alkalinity, high EC as well as strong supersaturation Can˜ete et al. 2011; Plestenjak et al. 2012; Blecha and with respect to calcite (Table 1). The second group con- Faimon 2014a), although supported by some measurements tains only the drip TC1 and it is called the anomalous drip. in vadose zone and caves (Otava 1995; Benavente et al. It shows lower conductivity/mineralization and a nearly 2010). Independence of P on surface conditions full saturation by calcite (Table 1). CO2ðHÞ (temperature variation, rainfall, amount of water in perched Evolution of the P is similar to some extent to the CO2ðWÞ collectors etc.) could indicate the source of CO was situ- cave air P evolution, with maxima in summer and 2 CO2ðairÞ ated in epikarst or vadose zone sufficiently deep below minima in winter. This is clearly visible at sites TC1 and surface thus it was not subjected to seasonal temperature CP1, but is less recognizable at sites CP2 and CP3 (Fig. 3). variations. Major role of epikarstic reservoir in karst water P shows periodicity, even though less pronounced CO2ðWÞ composition was speculated (Spo¨tl et al. 2005; Aquilina P than CO2ðairÞ (compare Fig. 4a, d, g, j with Fig. 4b, e, h, k) et al. 2006; Fairchild et al. 2006). Questionable is the origin and with somewhat distinct periods (Table 4). The simi- of epikarstic CO2 as (I) the availability and sources of larity in the PCO2ðWÞ and PCO2ðairÞ seasonality/trends is organic carbon are unknown and (II) the suitability of documented by the positive and statistically significant conditions for biodegradation processes are unknown as correlations of the raw data (see the left side of the well.

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80 Carbonates Evaporites (2016) 31:375–386 385

In principle, an alternative explanation of the invariant precipitation. That is in contrast with drips on straw spe-

PCO2ðHÞ is a water mixing. Waters of different age and leothems. Such drip has a quite extensive surface for composition are mixed in the perched collector or vadose degassing compared to the interaction surface with calcite. zone above the cave and their seasonality is ‘‘wiped off’’. An alternative explanation in a mix of the waters of dif- However, such mechanism is hardly conceivable, as it would ferent ages and compositions is improbable as was require very stable mixing ratio, constant transport paths, and explained before. probably rainfall without any variations. However, this is contradicted by drip hydrology and rainfall distribution in the area. Finally, one should take into consideration that the Conclusions contradiction between the soil PCO2 and PCO2ðHÞ could be caused by a methodological problem (systematic error) of Carbon dioxide and its variations in air and water were measurement in soil (Blecha and Faimon 2014b). studied in Punkva Caves in the Moravian Karst (Czech Republic) during the period between February 2012 and In contrast, the PCO2ðHÞ reconstructed from the anomalous drip TC1 shows very different behavior: there is a March 2013 with a monitoring frequency twice per month. Additional data from the hydrometeorological station on clear seasonality similar to PCO2ðairÞ and PCO2ðWÞ (Fig. 3a). Similarly, also correlations indicate a mutual relation (see surface above the cave complemented the cave data. Table 3). The periodicity resulting from the autocorrelation Cave air CO2 showed usual seasonality with maximum analysis (Fig. 4c; Table 4) suggests the same. This indi- in summer and minimum in winter with the period shorter than one year. This behavior was to some extent given by cates a distortion of original PCO ðHÞ values during the 2 the cave ventilation controlled by the difference between water evolution (probably by degassing together with cal- cave and external temperature. cite precipitation). This is demonstrated in the models of CO corresponding to equilibrium with aqueous car- geochemical evolution of dripwater in Fig. 5 (Peyraube 2 bonates showed similar seasonality, even though, less et al. 2012; Milanolo and Gabrovsˇek 2015). Projected data conclusive and with different periods. It seems plausible of the regular drip are aligned along the degassing line that it is controlled by CO concentration in the cave pointing to unique value of P regardless of the season 2 CO2ðHÞ atmosphere and dripwater degassing dynamics, partially when the data point was measured, although the winter influenced by dripwater hydrochemistry and drip rate. samples are usually more degassed. That might be caused CO concentration hypothetically participating in the by increased ventilation causing lower values of winter 2 initial formation of regular dripwaters showed stable values P . If the P is stable throughout the year (see CO2ðairÞ CO2ðHÞ without any obvious seasonality, which indicates its inde- the variations in Fig. 4f, i, l), the enhanced winter pendency on the surface conditions. Therefore, it is pos- D = PCO2 PCO2ðWÞ–PCO2ðairÞ would lead to faster and more sible that the source of CO2 is not in karst soils, as it is extensive degassing. However, regardless of the season, the generally believed. Instead, it could be situated deeper in - slope of degassing line is roughly 1 for all regular drips. epikarst or vadose zone. The intersections of the lines with x-axis (logP )at CO2ðWÞ The problem of calcite prior precipitation was demon- = SIcalcite 0 correspond to hypothetical partial CO2 pres- strated as an example of anomalous drip. This paper should P sure, CO2ðHÞ. As can be seen, the lines for drips TC2 and contribute to a better understanding of the limestone/cal- ZD show the same slope of * -1, but the lines are shifted, cite–water interaction in the carbonate karst. indicating different PCO2ðHÞ and, thus, different conditions at water formation in comparison with the drips CP. In Acknowledgments The research was supported by funding from addition, the data for the anomalous drip TC1 are much GEOtest, a.s. and Masaryk University. Laboratory analyses were kindly provided by GEOtest, a.s. The authors would like to thank the P more scattered and the calculated values of CO2ðHÞ on the anonymous reviewer for helpful comments. line with slope of -1 would probably not correspond to the original partial pressures. The model indicates that the dripwater TC1 firstly degassed to equilibrium with the cave References PCO2ðairÞ and then precipitated calcite. It is possible that the anomalous dripwater TC1 flows in the chimney above Absolon K (1970) Moravsky´ kras (in Czech). Academia, Praha, p 415 sampling site as a thin film on the wall, which allows it to Aquilina L, Ladouche B, Do¨rflinger N (2006) Water storage and degas much faster than the regular dripwater degassing in transfer in the epikarst of karstic systems during high flow form of a drop hanging on a straw speleothem. Further- periods. J Hydrol 327:472–485 Benavente J, Vadillo I, Carrasco F, Soler A, Liań C, Moral F (2010) more, the total area of interaction surface between water Air carbon dioxide contents in the vadose zone of a mediter- film and limestone (calcite) enables substantial calcite ranean karst. Vad Zone J 9:126–136

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81 386 Carbonates Evaporites (2016) 31:375–386

Blecha M, Faimon J (2014a) Spatial and temporal variations in carbon Otava JR (1995) Isotopic analyses and origin of CO2 in some dioxide (CO2) concentrations in selected soils of the Moravian Moravian caves. Acta Carsol 24:439–446 Karst (Czech Republic). Carbonates Evaporites 29:395–408 Parkhurst D, Appelo CAJ (1999) User’s guide to PHREEQC (Version Blecha M, Faimon J (2014b) Karst soils: dependence of CO2 2)—a computer program for speciation, batch-reaction, one- concentrations on pore dimensions. Acta Carsol 43:55–64 dimensional transport, and inverse geochemical calculations. US Domıńguez-Villar D, Fairchild IJ, Baker A, Carrasco RM, Pedraza J Geol Surv Water Resour Inv Rep 99–4259:310 (2013) Reconstruction of cave air temperature based on surface Peyraube N, Lastennet R, Denis A (2012) Geochemical evolution of atmosphere temperature and vegetation changes: Implications groundwater in the unsaturated zone of a kostic massif, using the for speleothem palaeoclimate records. Earth Planet Sci Lett PCO2—SIc relationship. J Hydrol 430–431:13–24 369–370:158–168 Peyraube N, Lastennet R, Denis A, Malaurent P (2013) Estimation of 13 Dreybrodt W (2008) Evolution of the isotopic composition of carbon epikarst air PCO2 using measurements of water d CTDIC, cave 13 and oxygen in a calcite precipitating H2O–CO2–CaCO3 solution air PCO2 and d CCO2. Geochim Cosmochim Acta 118:1–17 and the related isotopic composition of calcite in stalagmites. Plestenjak G, Eler K, Vodnik D, Ferlan M, Cˇ ater M, Kanducˇ T, Geochim Cosmochim Acta 72:4712–4724 Simoncˇicˇ P (2012) Ogrinc N (2012) Sources of soil CO2 in Faimon J, Licˇbinska´ M (2010) Carbon dioxide in the soils and calcareous grassland with woody plant encroachment. J Soils adjacent caves of the Moravian Karst. Acta Carsol 39:463–475 Sediment 12:1327–1338. doi:10.1007/s11368-012-0564-3 Faimon J, Licˇbinska´ M, Zajıćˇek P, Sracek O (2012) Partial pressures Sanchez-Can˜ete EP, Serrano-Ortiz P, Kowalski AS, Oyonarte C, of CO2 in epikarstic zone deduced from hydrogeochemistry of Domingo F (2011) Subterranean CO2 ventilation and its role in permanent drips, the Moravian Karst, Czech Republic. Acta the net ecosystem carbon balance of a karstic shrubland. Carsol 42:47–57 Geophys Res Lett 38:L09802 Fairchild IJ, Smith CL, Baker A, Fuller L, Spo¨tl C, Mattey D, Spo¨tl C, Fairchild IJ, Tooth AF (2005) Cave air control on dripwater McDermott F (2006) Modification and preservation of environ- geochemistry, Obir Caves (Austria): implications for speleothem mental signals in speleothems. Earth-Sci Rev 75:105–153 deposition in dynamically ventilated caves. Geochim Cos- Hromas J (2009) Jeskyneˇ - Chrańeˇna´ u´zemı´ Cˇ R, svazek XIV. AOPK mochim Acta 69:2451–2468 & ECB, Praha StatSoft, Inc. (2013) Electronic statistics textbook. Tulsa, http://www. Kuzyakov Y (2006) Sources of CO2 efflux from soil and review of statsoft.com/textbook/ partitioning methods. Soil Biol Biochem 38:425–448 Stumm W, Morgan JJ (1996) Aquatic chemistry: chemical equilibria Milanolo S, Gabrovsˇek F (2015) Estimation of the average carbon and rates in natural waters, 3rd edn. Wiley, New York, p 1040 dioxide flux degassing in a cave by the geochemistry of Tooth AF, Fairchild IJ (2003) Soil and karst aquifer hydrological percolating and pool water. Chemie der Erde—Geochemisty— controls on the geochemical evolution of speleothem-forming (submitted) drip waters, Crag Cave, southwest Ireland. J Hydrol 273:51–68

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82 APPENDIX 3

This appendix presents following research paper:

Pracný, P., Faimon, J., Všianský, D., & Kabelka, L. (2017). Evolution of Mg/Ca Ratios During Limestone Dissolution Under Epikarstic Conditions. Aquatic Geochemistry, 23(2), 119–139. http://doi.org/10.1007/s10498-017-9313- y

83 Aquat Geochem (2017) 23:119–139 DOI 10.1007/s10498-017-9313-y

ORIGINAL PAPER

Evolution of Mg/Ca Ratios During Limestone Dissolution Under Epikarstic Conditions

Pavel Pracny´1 • Jirˇ´ı Faimon1,2 • Dalibor Vsˇiansky´1 • Ludvı´k Kabelka3

Received: 1 June 2016 / Accepted: 21 February 2017 / Published online: 4 March 2017 Ó Springer Science+Business Media Dordrecht 2017

Abstract The Mg/Ca ratios in karst water are generally believed to comprise information on climate, and, being encoded in speleothems, they are utilized as paleoenvironmental proxy. However, the mechanism and dynamic of Mg release from limestone during dissolution is not well understood. A theoretical evolution of the Mg/Ca ratios during lime- = ° =- stone dissolution under epikarstic conditions (T 10 C, log PCO2 1.5) was studied via a dynamic model. The results were compared with (1) the dripwater data set collected in Punkva Caves (Moravian Karst, Czech Republic) during one-year period and (2) the published data from various locations worldwide. The modeling showed that the Mg/Ca ratios are governed by composition of Mg-calcite present in limestone. Two distinct stages in the dissolution dynamics were recognized: (1) an initial congruent dissolution with stoichiometric release of Ca and Mg and, subsequently, (2) an incongruent dissolution demonstrated by the gradual release of Mg with simultaneous Ca decrease via calcite precipitation. Additional identiﬁed factors inﬂuencing the reaction path and Mg/Ca ratio evolution were the dolomitic component of limestone and the ratio of limestone/solution boundary area to water volume. Finally, the water–rock interaction time controls the resulting Mg/Ca ratio in dripwater determining how far the dissolution proceeds along the reaction path. Thus, the study results indicate that Mg/Ca ratio depends on many factors in addition to climatic variables.

& Pavel Pracny´ [email protected] Ludvı´k Kabelka [email protected]

1 Department of Geological Sciences, Faculty of Science, Masaryk University, Kotla´rˇska´ 267/2, 611 37 Brno, Czech Republic 2 Department of Geological Sciences, Faculty of Science, Palacky´ University, 17. listopadu 1192/12, 779 00 Olomouc, Czech Republic 3 GEOtest, a.s., Sˇmahova 1244/112, 627 00 Brno, Czech Republic 84 123 120 Aquat Geochem (2017) 23:119–139

Keywords Moravian Karst Limestone dissolution Kinetic model Cave dripwater Mg-calcite

1 Introduction

Autochthonous cave deposits precipitated from karst waters (speleothems) are extensively utilized as terrestrial archives of paleoclimatic data (Fairchild et al. 2000). Beside stable isotopes (McDermott 2004; Verheyden et al. 2008; Lachniet 2009), trace elements are studied as a paleoclimatic proxy (e.g., Verheyden et al. 2000; Huang and Fairchild 2001; Cruz et al. 2007; Wong et al. 2011; Jochum et al. 2012; Sinclair et al. 2012). For example, Mg and Sr are proposed as rainfall indicators (e.g., McMillan et al. 2005; Tre- maine and Froelich 2013). It is believed that the content of trace elements in a speleothem and their ratios, as e.g., Mg/Ca or Sr/Ca, is controlled by climatic conditions projected into (a) dripwater hydrogeochemistry and then also into (b) speleothems. However, a direct link between surface conditions and speleothem composition is inﬂuenced by many factors disturbing the paleoclimatic interpretations (see, Fairchild et al. 2006a for a review). One approach to understand the mechanism of proxy transfer better is to compare selected climatic proxies with recent dripwater hydrogeochemistry (Riechelmann et al. 2014). Another approach is experimental or theoretical study of the mechanisms of recent karst processes with respect to selected proxies. Whereas trace elements in speleothems are widely documented (e.g., Frisia et al. 2012; Tan et al. 2014; Orland et al. 2014; Casteel and Banner 2015 and references therein), the release of trace and minor elements during limestone dissolution is little understood (Fairchild and Treble 2009). Thus, the goal of this work was (1) simulating limestone dissolution under epikarstic conditions based on a dynamic model with emphasis on Mg release and (2) a comparison of the model results with dripwater data from Moravian Karst caves and caves worldwide. In addition, special emphasis was placed on the incongruent dissolution of Mg-calcites.

2 Methods

2.1 Site of Study

Dripwater hydrogeochemistry was studied in the Punkva Caves in Moravian Karst (eastern part of the Czech Republic, see, e.g., Blecha and Faimon 2014) (Fig. 1). The cave system was formed by the underground river Punkva, which currently flows through a system of water passages from the Macocha Abyss to its outflow near the Punkva Cave visitor center in the Pusty´ Valley. The old cave level is currently dry (the dry part of the caves) and connects the entrance with the Macocha Abyss (Fig. 1a). The Moravian Karst is composed of folded and faulted Devonian limestones. Substantial part of the Punkva Caves develops in Lazˇańky Limestone, whereas upper parts of chambers and cave ceilings are developed in Vile´movice Limestone and the uppermost part of the overburden is composed of Lazˇańky Limestone (Fig. 1b). The limestone thickness above the cave ceiling is approximately 100 m in total (Hromas et al. 2009). Dripwaters were sampled twice a month from February 2012 to March 2013 in the dry part of the cave system (Fig. 1a). In total, 26 campaigns were conducted resulting in 127

85 123 Aquat Geochem (2017) 23:119–139 121

Zadní Chamber

30 m N Přední CP1 CP3 Chamber TC1 ZD Entrance CP2 Tunnel Corridor TC2 Macocha Abyss b altitude 500 m

Reichenbach Chamber 400 m Přední Zadní Chamber Chamber Macocha Abyss

Tunnel Corridor 350 m SW Vilémovice Lmst. Lažánky Lmst. NE c Sloup

Šošůvka

karst boundary Holštejn cave system Cave System build-up area Amaterská

Czech Republic Ostrov u Macochy 0 100 km Prague

Punkva Caves MK Vilémovice N N North Part 1 km of the Moravian Karst

86 123 122 Aquat Geochem (2017) 23:119–139 b Fig. 1 Sketch map of the dry part of the Punkva Caves with sampling sites (a); cave cross section (b); Czech Republic with the northern part of Moravian Karst (c). (based on Hromas et al 2009)

samples from 6 sites presented in this study. The dripwaters are issued from small straw stalactites in a corridor behind the Prˇednı´ Chamber (drips CP1–CP3), from a drapery (TC1), and straw stalactite (TC2) in the Tunnel Corridor and from a stalactite in the Zadnı´ Chamber (drip ZD) (Fig. 1a) (see, Pracny´ and Faimon 2013; Pracny´ et al. 2016a).

2.2 Dripwaters

Selected parameters were determined directly in the cave: pH (WTW 330i; precision ±0.005 pH), speciﬁc electrical conductivity EC (WTW Cond 3110; precision ±0.5%), alkalinity (acidimetric titration using 0.05 M HCl with potentiometric pH indication) and calcium concentration (complexometric microtitration with calcein as inner indicator). Further chemical analyses were conducted consecutively in the laboratory by ICP-OES (iCAP 6000 by Thermo Scientiﬁc).

2.3 Limestone Samples

Chemical composition (Ca, Mg, Fe, Mn, and Sr contents) of two samples representing the two limestone types was determined using AAS (TJA Solutions) after acid dissolution (HCl). Carbonate contents were estimated from the heat loss. Additional chemical components were determined using electron microprobe (WDX-analysis, Cameca SX100). Mineral composition of the samples was determined using powder X-ray diffraction (PXRD Bruker D8 Advance diffractometer). Minor minerals were identiﬁed via their habitus and stoichiometry using the electron microscope and microprobe equipped with EDXA (Cameca SX100).

2.4 Modeling and Calculations

The saturation index (SI) was calculated as a ratio of the actual ion activity product of aqueous components (Q) and the equilibrium constant (K), SI = log(Q/K), using PHREEQC code (Parkhurst and Appelo 2013) with the PHREEQC thermodynamic database. The kinetic module of this software was used to simulate the evolution of aqueous solution during limestone dissolution. In the simulation, the partial pressure of gaseous CO2 and temperature were adjusted to the average values prevailing in epikarst (Faimon et al. 2012; Peyraube et al. 2013; Pracny´ et al. 2016a). The Mg/Ca ratios used in data analysis were calculated from molar concentrations of the respective elements.

3 Results

Basic characteristics of the limestone samples as well as the mean chemical composition are presented in Table 1. Results of the XRD analysis showed that the samples are made of ca. 99% of calcite. Detailed composition of the bulk calcite determined by electron microprobe point analyses showed Mg (0.39 wt% in Lazˇa´nky Limestone, based on 6 measurements, and 0.14 wt% in Vile´movice Limestone, based on 3 measurements), Fe, and Mn in lower concentrations than their respective total concentration in the samples.

87 123 Aquat Geochem (2017) 23:119–139 123

Table 1 Characteristics, results of chemical analysis and XRD analysis of the limestones that constitute the cave overburden Lmst. Abrev. Appearance FeO MgO MnO SrO (Mg/Ca) 9 Minerals (wt%) (wt%) (ppm) (ppm) 1000 (XRD)

Lazˇa´nky Laz Dark-gray mostly 0.04 0.58 84 190 14.66 Calcite heavy-bedded (98.8), dolomite (0.9) Vile´movice Vil Bright-gray to gray, 0.01 0.32 42 103 7.99 Calcite very pure, (99.7) massive to heavy- bedded

The dolomite component was proved by XRD analysis only in Lazˇańky Lmst. (0.9%). Its identification in microprobe was difficult because of its similarity to calcite: BSE images often show dolomite only as a slight shadow in the calcite bulk. Additional minor minerals identified by using electron microprobe were quartz, clay minerals, mica, K-feldspar, and pyrite. The molar Mg/Ca 9 1000 ratios were 7.99 in the Vile´movice Lmst. and 14.66 in the Lazˇańky Lmst. (Table 1). Based on calculation, an ideal pure calcite limestone with 1% of dolomite should have the Mg/Ca 9 1000 ratio of 10. The hydrochemical parameters of dripwater samples are presented in Table 2. Most of the drips show minor variations through the year, whereas some drips (e.g., TC1) show more variable properties (for details see Pracny´ and Faimon 2013; Pracny´ et al. 2016b). The EC values varied in range 270–638 lScm-1. Calcium concentrations and alkalinity were in the range (1.31–3.79) 9 10-3 mol L-1 and (1.82–6.73) 9 10-3 eq L-1, respectively. The Mg concentrations were (5.3–7.4) 9 10-5 mol L-1. Majority of the drips monitored in Punkva Caves (regular drips) was supersaturated with respect to calcite (the mean saturation index ranged from 0.83 to 1.07) (Table 2). The only exception was the anomalous drip TC1 ranging from 0.21 to 0.32 with the mean value of SIcalcite = 0.14 ± 0.04 (Pracny´ et al. 2016b). The values of the saturation index with respect to dolomite (Table 2) were very low for drip TC1 (SIdolomite =-1.20 ± 0.09) in contrast to other drips where they were closer to zero (indicating equilibrium). The drip TC2 was slightly unsaturated (SIdolomite =-0.20 ± 0.10), and the drips CP1, CP2, and CP3 were slightly supersaturated (SIdolomite = 0.17 ± 0.08). Saturation index for the drip ZD was SIdolomite =-0.02 ± 0.07. The molar Mg/Ca 9 1000 ratios in the dripwater ranged from 14 to 60. The highest mean ratio, 46.2 ± 1.4, was found for the TC1 drip. The ratios for other dripwaters were substantially lower from 16.1 to 19.2 (see Table 2).

4 Modeling

The theoretical Mg/Ca evolution during limestone dissolution (reaction path) was simulated via a dynamical model. In the model, the limestone was regarded as a mix of two independent minerals: Mg-calcite and dolomite. Further carbonate minerals with zero initial mass (calcite and magnesite) were included in the model to be able to precipitate 88 123 124 123

Table 2 Chemical properties of the dripwaters from Punkva Caves (Pracny´ et al. 2016b) Sample CP1 CP2 CP3 ZD TC1 TC2

n 26 26 24 8 26 17 - EC (lScm 1) 628 ± 2 622 ± 2 616 ± 6 552 ± 2 297 ± 8 551 ± 3 pH 8.03 ± 0.05 8.1 ± 0.04 8.11 ± 0.05 8.09 ± 0.06 7.95 ± 0.06 7.99 ± 0.05 - Ca (mmol L 1) 3.50 ± 0.02 3.48 ± 0.02 3.49 ± 0.03 3.05 ± 0.02 1.47 ± 0.05 3.07 ± 0.04 - Alkalinity (mmol L 1) 6.23 ± 0.07 6.12 ± 0.06 6.14 ± 0.06 5.67 ± 0.04 2.14 ± 0.11 5.34 ± 0.04 - Mg (lmol L 1)58± 1.3 56 ± 1.0 57 ± 0.9 54 ± 1.7 67 ± 1.4 59 ± 0.9 (Mg/Ca) 9 1000 16.7 ± 0.4 16.1 ± 0.3 16.2 ± 0.3 17.7 ± 0.5 46.2 ± 1.4 19.2 ± 0.4

SIcalcite 1.01 ± 0.05 1.06 ± 0.04 1.07 ± 0.05 0.92 ± 0.04 0.14 ± 0.04 0.83 ± 0.05

SIdolomite 0.12 ± 0.10 0.19 ± 0.09 0.23 ± 0.10 -0.02 ± 0.07 -1.20 ± 0.09 -0.20 ± 0.10 qa ece 21)23:119–139 (2017) Geochem Aquat The confidence intervals are calculated for the level of significance of a = 0.05 n number of analyzed samples 89 Aquat Geochem (2017) 23:119–139 125 during the simulation. Simulation was conducted under epikarstic conditions (T = 10 °C, =- log PCO2 1.5; see Pracny´ et al. 2016a)—because epikarst is presumed to be the site of principal limestone dissolution (Williams 2008)—and for various dolomite/Mg-calcite ratios and fixed ratios of both the limestone/solution and the water/atmosphere boundary areas to aqueous volume. The processes considered in the model were based on the modified equations of Plummer et al. (1978), Chou et al. (1989), and Busenberg and Plummer (1982):

k !c1 CaCO þ Hþ Ca2þ þ HCO ð1Þ 3 3 kc1

k !c2 CaCO þ H CO Ca2þ þ 2HCO ð2Þ 3 2 3 3 kc2

k !c3 CaCO Ca2þ þ CO2 ð3Þ 3 3 kc3

kMgc1 ! Ca Mg CO þ Hþ xCa2þ þ yMg2þ þ HCO ð4Þ x y 3 3 kMgc1

kMgc2 ! Ca Mg CO þ H CO xCa2þ þ yMg2þ þ 2HCO ð5Þ x y 3 2 3 3 kMgc2

kMgc3 ! Ca Mg CO xCa2þ þ yMg2þ þ CO2 ð6Þ x y 3 3 kMgc3

k !d1 CaMg(CO Þ þ 2Hþ Ca2þ þ Mg2þ þ 2HCO ð7Þ 3 2 3 kd1

k !d2 CaMg(CO Þ þ 2H CO Ca2þ þ Mg2þ þ 4HCO ð8Þ 3 2 2 3 3 kd2

k !d3 CaMg(CO Þ Ca2þ þ Mg2þ þ 2CO2 ð9Þ 3 2 3 kd3

k !m1 MgCO þ Hþ Mg2þ þ HCO ð10Þ 3 3 km1

k !m2 MgCO þ H CO Mg2þ þ 2HCO ð11Þ 3 2 3 3 km2

90 123 126 Aquat Geochem (2017) 23:119–139

k !m3 MgCO Mg2þ þ CO2 ð12Þ 3 3 km3 where ki and k-i are the forward and backward rate constants for dissolution of i-carbonate mineral, respectively. Aqueous complexes were assumed to be at equilibrium with main * aqueous species. The specie H2CO3 represents the sum of carbonic acid and aqueous carbon dioxide. The system of equations was supplemented by the equation describing the CO2 exchange between water and atmosphere under open system conditions:

k !CO2 H O þ CO ðgÞ H CO; ð13Þ 2 2 2 3 kCO2

where kCO2 and kCO2 are the forward and backward rate constants. Under conditions of pH [ 5.5, the carbonate dissolution is pH independent and governed by the surface processes (for review see e.g., Morse and Arvidson 2002). Because of the very rapid increase of initial pH value, the overall reaction was simplified to Eqs. 3, 6, 9, and 12. Equilibrium constants corresponding to the equations were recalculated for 10 °C. Kcalc and Kdolo were calculated using PHREEQC and its default database while Kmagn was calculated from thermodynamic data (Robie and Hemingway 1995). Dissolution equilibrium constants of various Mg-calcites were recalculated for 10 °C from published data (Plummer and Mackenzie 1974) using dissolution enthalpies (Bischoff 1998). All used Keq are given in Table 3. The rate equations of the processes were derived in conventional manner. Back rates were implemented by relevant affinity terms (1 - Qi/Ki). The final rate equations are fLg Qcalc Rcalcite ¼ ackc3 1 ð14Þ V Kcalc

Table 3 The equilibrium constants for simulated processes Equation # Equilibrium constants calculated for T = 10 °C

a -9 3 Kcalc 3.89 9 10 b -9 6 K0.3%Mg-calc 3.98 9 10 b -9 6 K3%Mg-calc 5.26 9 10 b -9 6 K5%Mg-calc 6.43 9 10 b -9 6 K7%Mg-calc 8.45 9 10 b -8 6 K10%Mg-calc 1.38 9 10 a -17 9 Kdolo 1.91 9 10 c -9 12 Kmagn 2.05 9 10 -3 -1 9 -4 13 KCO2 (mol m Pa ) 5.21 10 a Calculated using PHREEQC database derived from works by Harned and Scholes (1941), Harned and Davis (1943), Larson et al. (1973), Reddy et al. (1981) b Constants for Mg-calcite interpolated from Plummer and Mackenzie (1974) and recalculated using Van’t Hoff equation with enthalpies from Bischoff (1998) c Calculated from thermodynamic data (Robie and Hemingway 1995)

91 123 Aquat Geochem (2017) 23:119–139 127 fLg QMgcalc RMgcalcite ¼ aMgckMgc3 1 ð15Þ V KMgcalc fLg Qdolo Rdolomite ¼ adkd3 1 ð16Þ V Kdolo fLg Qmagn Rmagnesite ¼ amkm3 1 ð17Þ V Kmagn where Ri represents the change of Ca and Mg concentrations in solution over time, respectively. {L}/V is the ratio of limestone/solution boundary area to water volume (m2 -1 L ). The mineral activities, ai, represent the molar proportion of calcite, Mg-calcite, dolomite, and magnesite in limestone. Rate constants kMgc3 for 10 °C were estimated using corresponding KMgcalc (Table 3) and backward dissolution rate kc3 for pure calcite estimated from Kcalc and kc3 (see Tables 3 and 4) based on relationship from Plummer et al. (1978). Values of used rate constants recalculated for 10 °C are presented in Table 4. The CO2 exchange between water and air is described by an equation derived from two- layer model (Liss and Slater 1974; Stumm and Morgan 1996)as fSg ÀÁ R ¼ k Hccc c ; ð18Þ CO2 V CO2 CO2ðgÞ CO2ðaqÞ where {S}/V is the ratio of water/atmosphere boundary area to water volume (m2 L-1), -1 cc kCO2 is the transfer coefﬁcient (m s ), and H is dimensionless Henry’s constant. The -3 terms cCO2ðgÞ and cCO2ðaqÞ represent CO2 concentrations (mol m ) in the air and water, respectively. To show principle reaction paths, simulations with the ratios {S}/V = 2.5 9 10-3 m2 L-1 and {L}/V = 0.01 m2 L-1 were conducted. For dissolution of composite carbonates (dolomite or Mg-calcite), the path is initially a straight-line reﬂecting mineral stoichiometry and indicating congruent dissolution. After the solution reaches saturation with respect to calcite, the process continues as an incongruent dissolution: The

Table 4 The rate constants used in simulation Rate constants recalculated for T = 10 °C Based on

-1 9 -6 kCO2 (m s ) 9.82 10 Stumm and Morgan (1996) Hcc 1.23 Sander (2015) -1 -2 -9 kd3 (mol s m ) 9.87 9 10 Busenberg and Plummer (1982) -1 -2 -10 a km3 (mol s m ) 1.68 9 10 Chou et al. (1989) -1 -2 -6 kc3 (mol s m ) 1.05 9 10 Plummer et al. (1978) -1 -2 -6 b k0.3%Mgc3 (mol s m ) 1.29 9 10 Plummer et al. (1978) -1 -2 -6 b k3%Mgc3 (mol s m ) 1.41 9 10 Plummer et al. (1978) -1 -2 -6 b k5%Mgc3 (mol s m ) 1.73 9 10 Plummer et al. (1978) -1 -2 -6 b k7%Mgc3 (mol s m ) 2.27 9 10 Plummer et al. (1978) -1 -2 -6 b k10%Mgc3 (mol s m ) 3.72 9 10 Plummer et al. (1978) a Recalculated for 10 °C utilizing Arrhenius equation with activation energy from Saldi et al. (2010) b Calculated using corresponding dissolution constant KMgcalc and k-c3 based on Plummer et al. (1978) and Kcalc (Table 3)

92 123 128 Aquat Geochem (2017) 23:119–139 increase of Ca concentration starts to slow down relatively to Mg and then, after reaching a sufﬁcient supersaturation, the Ca concentration decreases due to secondary calcite precipitation. During the process, Mg concentration steadily increases because of further primary mineral dissolution. At this stage, the reaction path is strongly nonlinear and its slope quickly increases and gets into negative values (Fig. 2). The paths corresponding to the dissolution of pure calcite and magnesite respectively are straight lines leading along the Ca or Mg axis to equilibrium of given mineral with solution. For limestone (modeled as a Mg-calcite/dolomite mix), the reaction path evolves in similar manner as in the case of pure minerals (Fig. 2). After reaching the partial equilibrium with calcite, the congruent dissolution changes into incongruent. The detailed path depends on Mg-calcite composition and the dolomite/calcite ratio (D/C ratio) (Fig. 3). In the case of limestones with less than 50% of dolomite (D/C \ 1), the effect of dolomitic component is almost indistinguishable from pure Mg-calcite dissolution. This is result of much faster dissolution kinetics of Mg-calcite in comparison with dolomite. Therefore, the composition of Mg-calcite is a key factor determining reaction path and Mg/Ca ratio evolution. The evolution of Mg/Ca ratio during the simulated Moravian Karst limestone dissolution (mix of 0.3% Mg-calcite and dolomite with D/C ratio of 1/99) was calculated with ﬁxed parameters of {L}/V = 0.01 m2 L-1 and {S}/V = 2.5 9 10-3 m2 L-1 (Fig. 4). These parameters were chosen by trial and error in order to emulate site conditions. In general, the actual value of the {L}/V ratio participating in karst processes is unknown. However, it may be estimated independently for (1) epikarstic aquifer and (2)

Fig. 2 Theoretical reaction 6.5E-03 paths for dissolution of pure 3% Mg-calcite carbonate minerals under 6.0E-03 5% Mg-calcite epikarstic conditions 7% Mg-calcite =- = ° (log PCO2 1.5, T 10 C) 5.5E-03 with highlighted partial 10% Mg-calcite equilibria. Simulated at ﬁxed 5.0E-03 dolomite ratios of mineral surface area to eq. calcite water volume, {L}/V 4.5E-03 eq. dolomite 2 -1 * 0.01 m L , and water eq. magnesite

table area to water volume, ] 4.0E-03 - {S}/V * 0.0025 m2 L 1 -1 3.5E-03 [mol L 3.0E-03 Tot

Mg 2.5E-03

2.0E-03

1.5E-03

1.0E-03

5.0E-04

0.0E+00 5.0E-04 1.0E-03 1.5E-03 2.0E-03 2.5E-03 3.0E-03 3.5E-03 4.0E-03 0.0E+00 -1 CaTot [mol L ] 93 123 Aquat Geochem (2017) 23:119–139 129

1.8E-03 dolomite Limestone 99/1 calcite 90/10 1.6E-03 5% Mg- + dolomite 50/50 Mg-calcite A 1.4E-03 B C D 1.2E-03 E ]

-1 F G 1.0E-03 H

[mol L I J Tot 8.0E-04 ite K m D/C = 99/1

Mg L dolo M 6.0E-04

4.0E-04 D/C = 50/50

2.0E-04 Mg-calcite D/C = 90/10 0.0E+00 5.0E-04 1.0E-03 1.5E-03 2.0E-03 2.5E-03 3.0E-03 3.5E-03 4.0E-03 0.0E+00 -1 CaTot [mol L ]

Fig. 3 Theoretical reaction path for limestones of different dolomite vs. 5% Mg-calcite ratios (D/C). Fixed ratios of limestone surface area to water volume, {L}/V * 0.01 m2 L-1, and water table area to water * 2 -1 =- = ° volume, {S}/V 0.0025 m L , under epikarstic conditions (log PCO2 1.5, T 10 C). The data points represent mean values measured in dripwater from various caves worldwide. The bars denote the data set span between minimum and maximum values: A Ballynamintra Cave (Baldini et al. 2012); B Bunker Cave (Immenhauser et al. 2010; Riechelmann et al. 2011); C Brown’s Folly Mine, F5 (Fairchild et al. 2006b); D Brown’s Folly Mine, B (Fairchild et al. 2006b); E Clamouse Cave (Fairchild et al. 2000); F Ernesto Cave (Fairchild et al. 2000); G Hollow Ridge Cave, Ballroom (Tremaine and Froelich 2013); H Hollow Ridge Cave, Duece (Tremaine and Froelich 2013); I Inner Space Cavern (Musgrove and Banner 2004); J Natural Bridge Caverns, group 1 (Wong et al. 2011); K Natural Bridge Caverns, group 2 (Wong et al. 2011); L Natural Bridge Caverns, group 3 (Wong et al. 2011); M Natural Bridge Caverns (Musgrove and Banner 2004) fissures in vadose zone. In epikarst, the {L}/V ratio was estimated according to properties of the sediment forming the perched epikarstic aquifer. It is evident that porosity and rock clast geometry are principal. The porosity defines water volume V in the spaces between rock clasts (Fig. 5), whereas the clast dimension/shape defines rock/limestone surface area {L}. Results of such calculations for feasible porosities and spherical shape of clasts range from 0.0014 to 0.1 m2 L-1 (Table 5). For modeling, the value of 0.01 m2 L-1 was chosen. In fissures, the {L}/V ratio would be given by the fissure wall area and the fissure diameter. It could be estimated from an etalon with the area of 1 m2 and diameter d (Fig. 5). As the fissure diameters are generally estimated as a fraction of a millimeter (Gabrovsˇek et al. 2004; Kaufmann et al. 2010), the value of {L}/V ratios could principally range from 1000 to 1 m2 L-1 (see the examples in Table 6). Nevertheless, such {L}/V ratios were ignored (see the discussion below). The value of the {S}/V ratio was deduced from the possible water body shapes and porosities (Table 7). The calculations showed the {S}/V ratio in the range of 0.05–0.0025 m2 L-1. As the ratio has negligible effect on reaction paths (see the discussion below), the value 2.5 9 10-3 m2 L-1 was set for all calculations. 94 123 130 Aquat Geochem (2017) 23:119–139

Fig. 4 Mg and Ca evolution 8 E-05 (reaction path) simulated for Moravian Karst limestones (0.3% Anomalous 7 E-05 drip 150th day Mg-calcite and dolomite with Regular D/C ratio of 1/99) with drips highlighted composition of the 6 E-05 anomalous drip and regular dripwaters from Punkva Caves. 5 E-05 ]

Fixed ratios of limestone surface -1 93th day area to water volume, 2 -1 {L}/V * 0.01 m2 L-1, and 4 E-05 {L}/V = 0.01 m L water table area to water volume, [mol L 2 -1 {S}/V * 0.0025 m L , under Tot 3 E-05

epikarstic conditions Mg (log P =-1.5, T = 10 °C) calcite eq. CO2 2 E-05

1 E-05 MgTot =F(CaTot ) 4.4th day 0 E+00 5.0E-04 1.0E-03 1.5E-03 2.0E-03 2.5E-03 3.0E-03 3.5E-03 4.0E-03 0.0E+00 -1 CaTot [mol L ]

Fig. 5 Schematic illustration of a perched epikarst aquifer and the

air relation between {L} and V in pore different positions of the proﬁle

1 m

1 m n ~ 10-30% {L}/Vป 0.01

primary reservoir 0.1 mm

1 m

fissure flow

1 m {L}/Vป 10

5 Discussion

5.1 Modeling Approach

Only the stoichiometric dissolution of individual carbonates (the ratio of released elements into solution is consistent with mineral stoichiometry) was assumed at modeling. This approach was chosen although some works point to a preferential release of trace elements,

95 123 Aquat Geochem (2017) 23:119–139 131

Table 5 Estimates of {L}/V ratio (m2 L-1) in epikarstic aquifer based on the properties of its ﬁlling (porosity and spherical clast dimensions) c.d.a Porosity (%)b

(mm) 10 15 20 25 30

5 0.108 0.068 0.048 0.036 0.028 10 0.054 0.034 0.024 0.018 0.014 20 0.027 0.017 0.012 0.009 0.007 50 0.011 0.0068 0.0048 0.0036 0.0028 100 0.005 0.0034 0.0024 0.0018 0.0014 a Clast diameter b Range of values from Williams (2008)

Table 6 Estimates of {L}/V ra- - Aperture (mm) {L}/V (m2 L 1) tio in karst fissure based on fissure diameter 0.001 1000 0.01 100 0.05 20 0.1 10 0.2 5 0.25 4 0.5 2 11 e.g., Sr or Mg (e.g., Busenberg and Plummer 1982; Fairchild et al. 2000; McGillen and Fairchild 2005; Morse et al. 2007; Sinclair 2011). We believe that this effect is merely an artifact of dissolution of fresh surfaces, which could be created by physical weathering processes. Nevertheless, chemical weathering is generally believed to play dominant role in limestone weathering (Fookes and Hawkins 1988). In principle, dissolution of fresh surfaces must become stoichiometric during the advanced dissolution stages due to the decrease of surface activity of the elements and bulk mineral dissolution. The dissolution in epikarst may be understood as an analogy of repeated dissolution in batch reactor: The old solution may periodically be replaced by a fresh one, but the activities of elements on mineral surface remain the same (in a steady state). Albeit dissolution of composite mineral is stoichiometric, it becomes incongruent during advanced stages (Fig. 2) due to concurrent precipitation of secondary calcite as discussed below. It is important to note that some simplifications were necessary during modeling: (1) a possible slight incorporation of Mg into precipitated calcite was ignored, (2) the mineral activities were set as constant parameters, (3) the activities of virtual minerals (calcite and magnesite with zero amounts at the beginning of dissolution) were set to be equal to Mg- calcite activity, and (4) the ratios {L}/V and {S}/V were estimated. Another uncertainty 0 arises from a weaker knowledge of various constants (Keq, k(dissol), Hr ), especially for Mg- calcites. Despite extensive studies with a wide variety of relevant parameters (for sum- mary, see Plummer and Mackenzie 1974; Mackenzie et al. 1983; Morse and Mackenzie 1990), these findings may not be easily applicable in karst conditions. In addition, a 96 123 132 Aquat Geochem (2017) 23:119–139

Table 7 Estimates of {S}/V ra- Porosity (%) r (m) h (m) tio of epikarstic aquifer (m2 L-1) based on its filling porosity and 0.5 1 2 aquifer shape: paraboloid, cone, and spherical cap Paraboloid 30 2.5 0.0133 0.0067 0.0033 30 5 0.0133 0.0067 0.0033 30 10 0.0133 0.0067 0.0033 20 2.5 0.020 0.010 0.005 20 5 0.02 0.01 0.005 20 10 0.02 0.01 0.005 10 2.5 0.04 0.02 0.01 10 5 0.04 0.02 0.01 10 10 0.04 0.02 0.01 Cone 30 2.5 0.02 0.01 0.005 30 5 0.02 0.01 0.005 30 10 0.02 0.01 0.005 20 2.5 0.03 0.015 0.0075 20 5 0.03 0.015 0.0075 20 10 0.03 0.015 0.0075 10 2.5 0.06 0.03 0.015 10 5 0.06 0.03 0.015 10 10 0.06 0.03 0.015 Spherical cap 30 2.5 0.013 0.006 0.003 30 5 0.013 0.007 0.003 30 10 0.013 0.007 0.003 20 2.5 0.02 0.01 0.005 20 5 0.02 0.01 0.005 20 10 0.02 0.01 0.005 10 2.5 0.04 0.02 0.01 r radius of the aquifer/reservoir 10 5 0.04 0.02 0.01 water table 10 10 0.04 0.02 0.01 h depth of the aquifer/reservoir nucleation of the virtual minerals is somewhat problematic. They were assumed to be able to precipitate on primary Mg-calcite, which allowed definition of their fixed activity in agreement with Mg-calcite. Notwithstanding all the uncertainties and simplifications, we believe that our approach is adequate for demonstration of temporal and spatial evolution of Mg/Ca ratios during karst rock dissolution.

5.2 Reaction Paths

There is a universal understanding that evolution of Mg in solution should reflect its content in the mineral structure, i.e., that trace elements follow the dissolution reaction path of the dominant cation, unless some processes inhibit the release (see, e.g., Dreybrodt and Eisenlohr 2000). However, the dissolution of Mg-calcite and dolomite is congruent 97 123 Aquat Geochem (2017) 23:119–139 133 only during the initial stage of dissolution—until calcite saturation is reached (Fig. 2). Then, the dissolution becomes effectively incongruent—the resulting solution is gradually enriched in Mg in comparison with mineral stoichiometry (Plummer and Mackenzie 1974). In principle, the reaction paths lead to a stoichiometric saturation of solution given by Mg- calcites (see, e.g., Morse and Mackenzie 1990). However, the paths will probably be terminated by precipitation of dolomite or magnesite (Fig. 2). Therefore, it seems that Mg- calcite undergoes an irreversible dissolution and re-precipitates into calcite and dolomite (or even magnesite). Nevertheless, specifics of this mechanism are beyond the scope of this study. In the case of limestones as a mixture of Mg-calcite and dolomite, the reaction paths are similar to the path of Mg-calcite. The dissolution dynamic of Mg-calcite exceeds and overlaps the dissolution of dolomite. Only if dolomite significantly prevails over Mg- calcite in limestone (D/C [[1), the resulting paths are substantially influenced by dolomite. This is consistent with the general view that the effect of dolomite dissolution is very low and requires very long residence times to enhance the Mg/Ca ratio significantly (Fairchild et al. 2006a).

5.3 Comparison of Dripwater Hydrogeochemistry with the Model

Most of the dripwaters in Punkva Caves show uniform hydrogeochemical properties (Table 2). As an exception, there is the anomalous drip TC1, which shows considerably distinct hydrochemical properties (Pracny´ et al. 2016b). It is close to equilibrium with calcite and shows high Mg/Ca ratio (Table 2) with a strong reduction of calcium content. The interpretation of anomalous drip’s properties is ambiguous in terms of our dissolution models as the water was clearly affected by additional processes beside dissolution and is rather a result of a prior calcite precipitation. The Mg/Ca ratios in regular dripwaters in Punkva Caves are low; however, they show a slight enrichment in Mg compared to stoichiometry of the Vile´movice and Lazˇańky Lmsts., the parent rocks at given site (see Table 1 and Table 2). This enrichment is consistent with our model showing the increase of Mg concentrations after reaching calcite saturation (Fig. 4). In addition to data from Punkva Caves, dripwater compositions from various caves worldwide were compared with the model. The plot in Fig. 3 shows that most of the dripwaters lie in the area defined by the reaction paths for the mix of 5% Mg-calcite and dolomite. The data maxima/minima lying outside the region delimited by reaction paths (Figs. 3, 4) are probably result of conditions different from those applied in the model. If the - limestone dissolved under the log PCO2 higher than 1.5, the reaction paths would extend along the x-axis toward higher Ca concentrations [e.g., Group 2 in Wong et al. (2011)as well as some Punkva Cave dripwaters, Fig. 3]. In addition, the Ca concentrations exceeding the predicted values might be a consequence of secondary mineral dissolution (e.g., evaporites as gypsum). Besides, the distinct temperatures may cause data inconsistency with the model due to rate constants’ temperature dependence. At higher temperatures, the area defined by reaction paths might broaden and the lines describing highly dolomitic limestones move closer to y-axis. Although most of the publications do not present detailed description of limestone composition (Mg content in limestone is usually represented by Mg/Ca ratio, which is inapplicable in context of this study), some relevant observations can be made. The wide range of Ca concentrations connected with relatively narrow range of Mg concentrations indicates dissolution of limestones with minor dolomitic or Mg-calcite components similar to Punkva Caves, e.g., the data from Bunker Cave

98 123 134 Aquat Geochem (2017) 23:119–139

(Immenhauser et al. 2010; Riechelmann et al. 2011). On the contrary, the wide range of Mg concentrations suggests an evolution along the reaction paths corresponding to higher dolomitic or Mg-rich calcite content in limestone, e.g., the data from Natural Bridge Cavern (Wong et al. 2011, Group 1), or even evolution of dripwaters from caves developed in dolostone, e.g., Clamouse Cave (Fairchild et al. 2000).

5.4 Dynamics of Mg/Ca Evolution

Based on our model, the Mg/Ca ratio in the regular dripwaters corresponds to the residence time between 100 and 150 days, which is the more plausible value compared to usual residence times (Kamas et al. 2015; Faimon et al. 2016). The residence time of water in karst profile plays important role. As water composition evolves along the relevant reaction path, the time of interaction determines the resulting dripwater composition. In addition, {L}/V ratio has very similar effect as the water residence time: It changes dissolution dynamics substantially and determines the instantaneous position on reaction path. What is more, the ratio modifies the reaction path shape determining the transition from congruent to incongruent dissolution (Fig. 6). The dependence of reaction path on the {L}/V ratio might be due to dissimilar variations in dissolution/precipitation dynamics. This demonstrates that the ratio is of much higher importance than is generally expected. We believe that the {L}/V ratios in range 0.01–0.005 m2 L-1 in epikarstic perched aquifer (Fig. 5; Table 5) are representative for entire karst profile. The enhanced {L}/V ratio in the fissures of vadose zone does not seem very important: The water entering the fissure is probably already close to equilibrium with bedrock limestone and epikarstic gaseous CO2. The water flowing through the fissure represents a closed system with respect to gaseous CO2. The CO2 consumed at limestone dissolution cannot be replenished from the pore air overhead as the water flow velocity in fissure exceeds the rate of aqueous CO2 diffusion. That is consistent with the idea of prevailing corrosion in epikarst diminishing downward the vadose zone.

Fig. 6 Variations of reaction 3.0E-03 paths for 5% Mg-calcite with changing {L}/V ratio: (a) 0.2 m2 -1 2 -1 2 L ;(b) 0.1 m L ;(c) 0.01 m 2.5E-03 5% Mg-cacite L-1;(d) 0.001 m2 L-1. Fixed water table area to water volume, 2 -1 {S}/V * 0.0025 m L , under 2.0E-03 epikarstic conditions =- = ° -1 (log PCO2 1.5, T 10 C)

[mol L ] 1.5E-03

Tot

Mg 1.0E-03

(a) 5.0E-04 (b) (c) (d) 0.0E+00 5.0E-04 1.0E-03 1.5E-03 2.0E-03 2.5E-03 3.0E-03 3.5E-03 4.0E-03 0.0E+00 CaTot [mol L-1 ] 99 123 Aquat Geochem (2017) 23:119–139 135

Modeling showed that the {S}/V ratios had negligible effect on the reaction path. The most significant influence is a brief decrease in PCO2ðaqÞ (virtual PCO2ðgÞ corresponding to concentration of aqueous CO2) during initial stages of dissolution (Fig. 7). It should be noted that the {S}/V ratio might additionally be influenced by properties of the capillary fringe that might vary seasonally.

5.5 Comparison with Other Processes Affecting Trace Element Ratios

The presented model describes the Mg/Ca ratio evolution in dripwaters via effects of Mg- calcite/dolomite incongruent dissolution and its dynamics. However, the ratio may be to some extent influenced by additional processes. One such phenomenon might be a non- stoichiometric dissolution of calcite/limestone causing difference in the element ratios between host rock and solution as discussed above. Additional phenomenon changing the Mg/Ca ratio might be (1) a capture of specific cations on rock/mineral surfaces along water flow path (e.g., on clay minerals) or (2) a dissolution of non-carbonate minerals rich in Mg (e.g., evaporites as gypsum). Another important factor changing the Mg/Ca ratios in dripwater is the process of prior calcite precipitation (PCP, Fairchild et al. 2000). It is based on the idea that aqueous Ca concentration decreases during calcite precipitation and, thus, water Mg/Ca ratio increases. Resulting Mg/Ca ratio is then stored in growing speleothems. PCP occurs whenever aqueous CO2 is allowed to degas into vented spaces within the vadose zone between epikarst aquifer and dripsite (e.g., Wong et al. 2011). As it is believed that PCP is promoted by drier climatic conditions, the Mg/Ca ratio is frequently tested as a paleoenvironmental proxy (Fairchild et al. 2000, 2006a; McMillan et al. 2005; Tremaine and Froelich 2013). The presented study shows that incongruent dissolution is similarly productive of the increase in the Mg/Ca ratios as PCP. Although these effects are possibly distinguishable in

Fig. 7 Variations of PCO2ðaqÞ -1.3 during dissolution of 5% Mg- calcite with changing {S}/V ratio: (a) 0.1 m2 L-1;(b) 0.01 m2 L-1; -1.4 5% Mg-calcite (c) 0.001 m2 L-1. Fixed ratio of limestone surface area to water (a) - -1.5 volume, {L}/V * 0.01 m2 L 1, under epikarstic conditions (log P =-1.5, T = 10 °C) (b) CO2 -1.6

CO2(aq) -1.7

logP

-1.8

-1.9 (c)

-2 1E-01 1E+00 1E+01 1E+02 1E+03 1E+04 elapsed time [day] 100 123 136 Aquat Geochem (2017) 23:119–139 dripwater (e.g., based on anomalous Mg concentration in comparison with its standard values on site), they are indistinguishable in speleothems.

5.6 Implications for Paleoenvironmental Studies

Presented model shows that changes in Mg/Ca ratio caused by precipitation of secondary calcite may occur not only by variations in climatic conditions as PCP but also during incongruent dissolution of limestone under constant conditions in epikarst. In contrast to PCP showing negative correlation to rainfall (Fairchild et al. 2000; Wong et al. 2011), the increase caused by dissolution dynamics may be initiated by intensive rainfalls and may show positive correlation (see, e.g., Baldini et al. 2012). The rainwater inflow may cause overflow of semi-isolated epikarstic perched reservoirs, in which the water is rich in Mg due to extended residence time leading to very advanced position along the reaction path. In addition to the incongruent dissolution, this study identifies further factors influencing the reaction path and, consecutively, Mg/Ca ratios in water: (1) the composition of limestone (Mg-calcite composition, dolomite content), (2) the time of water-mineral interaction, (3) the ratio of rock surface area to water volume (given by karst structure), and

(4) the epikarstic PCO2 . With exception of the PCO2 , all the factors are intrinsically dependent on water flow paths that may naturally change during karst evolution independently on climate changes. The individual reaction paths mirror the composition of Mg-calcites and dolomite content in bedrock limestones (Fig. 3). Therefore, if different limestone types were present in the karst profile above a cave (the case of the Punkva Caves in Moravian Karst), the resulting reaction path would be dependent on the limestone type in contact with the percolating water. A similar effect would have mixing of waters formed at contact with different limestones (e.g., waters from different epikarst aquifers). It should be emphasized that water flow paths and mixing in vadose zone may change both spatially and temporary. All presented effects should be considered in prospective attempts to interpret the paleoclimate signal in speleothems.

6 Conclusion

The evolution of Mg/Ca ratios during limestone dissolution under epikarstic conditions was studied theoretically using dynamic models. Results of the modeling were compared with actual dripwater data from Punkva Caves (Moravian Karst, Czech Republic) and published dripwater data from various caves worldwide. The presented work indicates that a key factor determining Mg/Ca ratios in dripwaters is the Mg-calcite composition. Besides, the ratio may vary in dependence on additional parameters as follows: (1) the ratio of Mg-calcite and dolomite in parent rocks, (2) the time of the interactions (water residence time in epikarst), and (3) the ratio between limestone surface area and water volume. With changing water ﬂow paths, all the parameters may change spatially and temporally, independently on climatic conditions. All the former facts may complicate proper interpretations of the climate signal preserved in speleothems.

Acknowledgements We thank GEOtest, a.s., for ﬁnancial and material support of ﬁeld work and analyses. Many thanks also belong to Masaryk University (Brno) and Palacky´ University (Olomouc) for the additional support. We would like to thank the anonymous reviewer for valuable comments that helped us to improve the manuscript.

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104 123 APPENDIX 4

This appendix presents following research paper:

Lang, M., Faimon, J., Pracný, P., & Kejíková, S. (2017). A show cave management: Anthropogenic CO2 in atmosphere of Výpustek Cave (Moravian Karst, Czech Republic). Journal for Nature Conservation, 35, 40–52. http://doi.org/10.1016/j.jnc.2016.11.007

105 Journal for Nature Conservation 35 (2017) 40–52

Contents lists available at ScienceDirect Journal for Nature Conservation

journal homepage: www.elsevier.de/jnc

A show cave management: Anthropogenic CO2 in atmosphere of Vypustek´ Cave (Moravian Karst, Czech Republic)

Marek Lang a,b,∗, Jiríˇ Faimon a,c, Pavel Pracny´ a, Sandra Kejíková a a Department of Geological Sciences, Faculty of Science, Masaryk University, Kotlárskáˇ 2, 611 37 Brno, Czechia b Department of Geology and Pedology, Faculty of Forestry and Wood Technology, Mendel University, Zemedˇ elskáˇ 3, 613 00 Brno, Czechia c Department of Geology, Faculty of Science, Palacky´ University Olomouc, 17. listopadu 1192/12, 771 46 Olomouc, Czechia article info a b s t r a c t

Article history: Anthropogenic impact on CO2 levels was studied in the Bear Chamber of the Vypustek´ Cave, a show Received 20 February 2016 cave in the Moravian Karst (Czech Republic), during a period of active ventilation and enhanced atten- Received in revised form dance. The study showed that the natural CO2 levels were controlled by (i) the natural CO2 influxes from 28 November 2016 −2 −1 soils/epikarst (up to ∼5.64 × 10 mol s ); and, (ii) the advective CO2 fluxes out of cave atmosphere (up Accepted 28 November 2016 −2 −1 to 4.66 × 10 mol s ). During visitor presence, the anthropogenic CO2 flux into the chamber reached up to −1 ∼0.13 mol s and exceeded all other CO2 fluxes. The reachable anthropogenic steady states at sufficient Keywords: duration of stay (up to 2.65 × 10−1 mol m−3) could exceed the natural CO levels by factor of more than Anthropogenic impact 2 Carbon dioxide nine based on the number of visitors. Recession analysis of anthropogenic pulses showed that intervals ∼ Dynamic model between individual visitor groups would have to be up to 6 h long if the cave environment has to return Recession analysis to natural conditions. As such pauses between individual tours are hardly realizable, a risk analysis was Response time conducted to find the consequences of breaking natural conditions. It showed that the condition under

Show cave which dripwater becomes aggressive to calcite (i.e., the point when PCO2 in cave atmosphere exceeds the −1.56 hypothetical CO2 concentrations in epikarst that has participated on the water formation, PCO2(H) = 10 ) is potentially reachable under extreme conditions only (enormous visitor stay period and visitor number).

In case of condensed water, however, any increase in CO2 concentration will cause an increase of water aggressiveness to calcite. Therefore, in the periods and sites of enhanced condensation, it is important to strive for preservation of natural conditions. © 2016 Elsevier GmbH. All rights reserved.

1. Introduction Pracny,´ Faimon, Sracek, Kabelka, & Hebelka, 2015; Milanolo and Gabrovsek,ˇ 2015). Whereas a high PCO2(H) controls saturation of Carbon dioxide (CO2) plays a key role in carbonate karst sys- percolating water with respect to calcite, the lower PCO2(C) is tem by controlling the karst processes as limestone dissolution responsible for dripwater degassing (release of the excessive CO2) (e.g., Stumm and Morgan, 1996) or speleothem growth via cal- (Holland, Kirsipu, Huebner, & Oxburgh, 1964). In principal, the cite/aragonite precipitation (e.g., Dreybrodt 1988; Frisia et al., instantaneous CO2 concentration in a cave is given by balance 2011). The driving force for the latter processes is the differ- of input and output CO2 fluxes. Whereas the input fluxes are ence in the CO2 partial pressure between (1) the soil/upper associated with soil/epikarstic sources, the output fluxes are con- epikarst, PCO2(H), and (2) the cave atmosphere PCO2(C) (White, trolled by cave airflow (Spötl, Fairchild, & Tooth, 2005; Banner, 1988; Ford and Williams, 2007). Some studies showed that Guilfoyle, James, Stern, & Musgrove, 2007; Baldini, McDermot, hypothetical PCO2(H) values can be reconstructed from the hydro- Hoffmann, Richards, & Clipson, 2008; Fernández-Cortes, Sanchez- geochemistry of cave dripwaters (Faimon, Licbinská,ˇ Zajícek,ˇ & Moral, Cuezva, Canaveras,˜ & Abella, 2009). Recently, Lang et al. Sracek, 2012; Peyraube, Lastennet, Denis, & Malaurent, 2013; (2016) showed that also a part of the input CO2 fluxes could be associated with cave airflow. This clearly illustrates the importance of cave ventilation. Theoretical foundations of cave ventilation were elaborated by Cigna (1968). His ideas were further devel- ∗ Corresponding author at: Department of Geological Sciences, Faculty of Science, oped by de Freitas, Littlejohn, Clarkson, & Kristament 1982; Pflitsch Masaryk University, Kotlárskáˇ 2, 611 37 Brno, Czechia. and Piasecki (2003), Kowalczk and Froelich (2010), Benavente, E-mail address: [email protected] (M. Lang). http://dx.doi.org/10.1016/j.jnc.2016.11.007 1617-1381/© 2016 Elsevier GmbH. All rights reserved.

106 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 41

Vadillo, Linan,˜ Carrasco, & Soler, 2011; Faimon, Troppová, Baldík, 3. Monitoring & Novotny,´ 2012; Sánchez-Canete,˜ Serrano-Ortiz, Domingo, & Kowalski, 2013; Faimon and Lang (2013), James, Banner, & Hardt, Data on the cave atmosphere variables were collected during 2015 and others. three individual monitoring campaigns between November 2013 The environment of show caves is significantly influenced by and June 2015. The campaigns were conducted in the Vypustek´ (1) human activity connected with works required for ensuring Cave during the occasional cultural events connected with a signif- the access for visitors into the cave (e.g., de Freitas, 2010), or icantly enhanced attendance. The individual events were chosen (2) by visitor presence inside the cave. The latter effect may be in order to cover different meteorological conditions. In addition divided into three main categories: (1) impact of anthropogenic to the number of visitors, the cave CO2 concentrations and the CO2 on cave environment (e.g., Pulido-Bosch, Martín-Rosales, temperatures of cave air and external air were monitored with a López-Chicano, Rodríguez-Navarro, & Vallejos, 1997; Hoyos, Soler, minute time steps. The CO2 concentration was measured in the Canaveras,˜ Sánchez-Moral, & Sanz-Rubio, 1998; Carrasco, Vadillo, Bear Chamber at 2 m above the cave floor. It was detected by a hand- Linán,˜ Andreo, & Durán, 2002 ; Lang, Faimon, & Ek, 2015b); (2) held device (FYAD00CO2B10 digital sensor linked to the ALMEMO direct monitoring of anthropogenic influence from cave CO2 con- 2290-4 V5 Ahlborn data logger) with the measuring range from 0 to centrations (e.g., Dragovich and Grose, 1990; Song, Wei, & Liang, 10,000 ppmv and accuracy ±100 ppmv + 5% of measured value. For 2000; Faimon, Stelcl,ˇ & Sas, 2006; Linán,˜ Vadillo, & Carrasco, 2008; modeling, the volume concentration (in ppmv unit) was converted Fernández-Cortes et al., 2009; Milanolo and Gabrovsek,ˇ 2009; to molar concentration (mol m−3), based on the Ideal Gas Law and Sebelaˇ et al., 2013; Lang, Faimon, & Ek, 2015a); and, (3) cave given temperature/pressure, management and cave environment conservation (e.g., Fernández, − P c [molm 3] = c [ppmv], (1) Gutierrez, Quindós, Soto, & Villar, 1986; Calaforra, Fernández- co2 6 co2 Cortés, Sánchez-Martos, Gisbert, & Pulido-Bosch, 2003; Lario and 10 RT Soler, 2010). In addition, there appeared attempts to describe where P is a barometric pressure [Pa], R is the universal gas constant −1 −1 the anthropogenic CO2 behavior by a complete dynamic model [R = 8.3145 J kg K ] and T is a temperature [K]. The temperatures (Faimon, Stelcl,ˇ Sas et al., 2006; Milanolo and Gabrovsek,ˇ 2009; Lang for T calculations were logged (i) in the exterior, approximately et al., 2015a). Despite the effort, some aspects of anthropogenic 10 m from the cave entrance, and (ii) in the Bear Chamber. Tem- influence on cave environment remain little understood. There are perature was measured by COMET S3120 data loggers (measuring ◦ ◦ questions of (1) the persistence of anthropogenic influence after range: −30 to +70 C; accuracy: ±0.4 C) (TR Instruments Inc.). The cave exposure, (2) the extent of affecting natural karst processes, visitor numbers and entering time were logged in front of the cave. and (3) the real threat to cave environment (e.g., speleothems). Dripwater samples from two drips were collected during two In this study, we discuss the anthropogenic impact on cave CO2 individual monitoring campaigns in July and December 2015. Drip levels based on (1) new data sets from the Vypustek´ Cave (Mora- D1 is situated in Císarskáˇ Chamber and the drip D2 is situated in vian Karst) and (2) a simplified dynamic model following Lang Skrapovˇ y´ Chamber. Both drips fall from small soda straw stalac- et al. (2015a). Data come from the periods of active ventilation tites ca. 4–5 m above the cave floor. Immediately in the cave the and enhanced attendance. The goals of this study were (1) to ana- specific electrical conductivity EC (Greisinger GMH 3431; preci- ± ± lyze both the natural and anthropogenic CO2 levels, (2) to describe sion 0.5%) and pH (WTW 330i; precision 0.005 pH) of dripwater dynamics of the adaptation of cave environment after the depar- were measured as well as the immediate atmospheric pressure of ture of all visitors, (3) to analyze the effect of anthropogenic CO2 CO2 in the cave, PCO2(C) (Ahlborn ALMEMO 2594-4S with the sensor levels on dripwater hydrogeochemistry, and (4) to summarize the FYAD00CO2B10; accuracy ±(100 ppmv +5% of meas. value)). Alka- principles of a better cave management. linity (acidimetric titration with potentiometric pH indication), calcium concentration (complexometric microtitration, calcein as inner indicator) and magnesium concentration (AAS; Solaar M5, TJA Solutions) were analyzed in a lab. Statistical analysis was conducted using the Statistica 12, StatSoft Inc. program. (StatSoft Inc., 2. Site of study 2015). The data on dripwater geochemistry were processed using the program PHREEQC (Parkhurst and Appelo, 2013). The site of study, the Vypustek´ Cave, is situated in the central part of the Moravian Karst in Krtinyˇ valley about 2 km south- 4. Results westward from the village Krtinyˇ (Fig. 1). The position and sketch map of the cave are illustrated in Fig. 1. The cave was formed by 4.1. Monitoring campaign I the Krtinyˇ stream in the Middle/Upper Devonian limestone of (i) the Macocha Formation (Lazánkyˇ and Vilémovice limestone) and During the Campaign I (running in the Bear Chamber from (ii) the Líseˇ nˇ Formation (Krtinyˇ Limestone). Presently, the Krtinyˇ 30 November to 2 December 2013), two unique tourist events Stream flows through the lower cave passages. The cave consists occurred: the first event covered 9 tours including 856 persons, of a complex of relatively narrow corridors and large chambers whereas the second event covered 8 tours including 776 per- comprising a total length of about 2 km. Due to a complex mor- sons (Fig. 2a). The CO2 concentrations varied in a wide range phology (two levels, four known entrances, and some presumed (2.69–5.32) × 10−2 mol m−3, i.e., 619–1225 ppmv (first event) and hidden openings), the cave shows typical dynamic air circulation. (2.82–5.48) × 10−2 mol m−3, i.e., 649–1263 ppmv (second event) In the 1960s, an underground fallout shelter and a secret command depending on the number of visitors (Fig. 2b). There are con- post of the Czechoslovak army were built in the corridor between spicuous CO2 peaks representing the anthropogenic impact of the the entrance and the Lion Chamber. Since 2008, the cave is open individual tours. As further tours followed shortly after, the cave to tourists with a visitor rate of 15,000–20,000 people per year. was not able to return to the initial values, CO2 concentrations 3 The Bear Chamber with about 7800 m of total volume situated cumulated and reached maxima by the time of the last tour in given approximately 100 m from the cave entrance was chosen as the event. In turn, the 12-h period without visitors between both events monitoring site (Fig. 1). The mean annual precipitation in the area seems to be sufficient under given conditions for return to the initial is about 700 mm; the mean annual temperature of the external values (about 600 ppmv). A temperature difference T = T ◦ exterior atmosphere is about 8 C. − Tcave during the whole monitoring campaign ranged between

107 42 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52

Fig. 1. The cave position and sketch map of the monitoring site.

−8.1 and −2.9 ◦C(Fig. 2c). The negative values of T indicate an hydrogeochemical properties were slightly different (see upward airflow (UAF) ventilation mode (Faimon, Troppová et al., Table 1). The D1 had higher specific electrical conductivity 2012; Faimon and Lang, 2013). EC = 736–775 ␮S cm−1 compared to D2 with EC = 360–395 ␮S cm−1. Only slight variations in dripwater hydrogeochemistry were 4.2. Monitoring campaign II observed in comparison of summer and winter measurements and were in consent with other dripwater properties in Moravian This 48-h campaign was run in the Bear Chamber in the Karst (Pracny,´ Faimon, Sracek et al., 2015). The calcium con- period from 29 November to 1 December 2014. Two tourist centrations were slightly higher in summer, D1 = 4.16 mmol L−1 events happened: At the first event, 828 visitors in 9 tours and D2 = 2.06 mmol L−1, than in winter, D1 = 3.87 mmol L−1 passed the chamber. At the second event, 775 visitors in 8 tours and D2 = 1.88 mmol L−1. The alkalinity was lower in summer −1 −1 passed the chamber (Fig. 3a). The CO2 concentrations varied in (5.56 mmol L ) than in winter (5.88 mmol L ) in the case of D1. the wide range from ∼ 2.60 × 10−2 mol m−3 (600 ppmv) to the In the case of D2, the alkalinity was lower in winter (2.80 mmol L−1) maximum of 5.32 × 10−2 mol m−3, i.e., 1227 ppmv (first event) or than in summer (3.18 mmol L−1). The pH values were very similar to 4.91 × 10−2 mol m−3, i.e., 1133 ppmv (second event) (Fig. 3b). with higher values in winter. The magnesium concentrations were −2 −1 Between both events, the initial CO2 values about 600 ppmv were measured only in December 2015 (4.56 × 10 mmol L for D1 reached. Based on a temperature difference T varying from −10.5 and 2.81 × 10−2 mmol L−1 for D2). The resulting Mg/Ca × 1000 to −6.9 ◦C, the cave persisted in UAF ventilation mode during the ratios were 11.18 and 17.03 for D1 and D2, respectively. The entire campaign (Fig. 3c). immediate partial pressure of the CO2 at the sampling site PCO2(C) was lower in winter than in summer (Table 1). 4.3. Monitoring campaign III

5. Cave CO2 level modeling This 48-h campaign was run in the Bear Chamber in the period from 6 to 7 June 2015. During the two tour events, 534 visitors For modeling of CO2 evolution, the Bear Chamber was con- divided into 8 tours (first event) and 94 visitors divided into 10 sidered a perfectly mixed reactor represented by a homogeneous tours (second event) passed the chamber (Fig. 4a). The CO2 concen- reservoir with both the input and output CO2 fluxes (Fig. 5). The trations in the chamber varied in relatively narrow ranges (Fig. 4b). fluxes in mol s−1 units include (1) the direct net natural flux into Whereas the first event with higher number of visitors induced rec- the chamber from epikarst, jN (covering all the diffusive/advective ognizable CO2 peaks, any impact of the second event (connected fluxes, and the flux stemming from dripwater degassing), (2) the with substantially lower visitor number) was barely recognizable. anthropogenic flux jA (resulting from a human respiration), (3) the The maximum range between natural and anthropogenic levels advective input flux from exterior or/and from an adjacent cave was about 8.67 × 10−3 mol m−3 (200 ppmv). Based on the temper- ◦ space, jin, and (4) the advective output flux out of the chamber, jout. ature difference T ranging from 2.3 to 21.9 C, the cave persisted 3 −1 The fluxes jin and jout are driven by (i) airflow v [m s ], and (ii) throughout the campaign in DAF ventilation mode (Fig. 4c). CO2 concentration in adjacent spaces (cin) or in the cave chamber (c), all in [mol m−3]. 4.4. Dripwater hydrogeochemistry The instantaneous CO2 concentration in cave atmosphere is given by the sum of all individual fluxes Both drips were hydrologically active during the sampling − campaigns with similar discharge of 8–17 drips min 1 (the sum- dnCO Vdc 2 = = j + j + v c − vc, (2) mer discharge exceeded the winter discharge). In contrast, dt dt N A in

108 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 43

Fig. 2. Monitoring Campaign I (30 November–2 December 2013; Bear Chamber, Vypustek´ Cave): the visitor number per individual tours (the column widths represent the periods of visitors’ stay in the chamber) (a), CO2 concentrations (the red dashed lines represent hypothetical recession curves in case of no further visitors) (b), and temperature difference T = Texterior − Tcave (c). See text for details. (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)

where nCO2 is the total content of CO2 in the chamber atmosphere is an initial CO2 concentration in cave atmosphere), the integration 3 [mol], t is time [s], V is the chamber volume [m ], jN is natural CO2 yields −1 −1 flux [mol s ], jA is anthropogenic CO2 flux [mol s ], v is volumet- − ric velocity of the airflow through the cave chamber [m3 s 1], c − v t in ln(j + j + v c − vc) = ln e V + ln(j + j + v c − v c0), (3) −3 N A in N A in is CO2 concentration in adjacent cave spaces [mol m ], and c is an instantaneous CO concentration in the chamber atmosphere 2 A rearrangement gives [mol m−3]. Eq. (2) was integrated on the assumption that, V, jN, jA, cin, and v jN jA jN jA − v t are constant. Under the initial conditions that c = c at t = 0 (where c c = + + c − ( + + c − c ) e V . (4) 0 0 v v in v v in 0

Table 1 Hydrogeochemical properties of sampled dripwaters.

−1 a b c d e Drip EC [␮S cm ] pH SI(calcite) Mg/Ca × 1000 log PCO2(W) log PCO2(H) log PCO2(C) July 2015 D1 736 8.02 1.03 NA −2.61 −1.56 −2.98 D2 395 8.05 0.58 NA −2.87 −2.27 −2.99

December 2015 D1 775 8.15 1.19 11.18 −2.75 −1.85 −3.30 D2 360 8.18 0.64 17.03 −3.05 −2.51 −3.32

NA − not analyzed. a Calcite saturation index. b Mg/Ca ratio in dripwater. c Logarithm of the CO2 partial pressure corresponding to aqueous carbonate species in dripwater. d Logarithm of hypothetical partial pressure of CO2 participating on water chemistry formation in soil/epikarst. e Logarithm of partial pressure of CO2 in cave atmosphere.

109 44 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52

Fig. 3. Monitoring Campaign II (29 November–1 December 2014; Bear Chamber, Vypustek´ Cave). Visitor number per individual tours (the column widths represent the staying periods of visitors in the chamber) (a), CO2 concentrations (the red lines represent hypothetical recession curves in case of no further visitors) (b), and temperature difference T = Texterior − Tcave (c). See text for details. (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)

At steady state, all the CO2 ﬂuxes into/out of the chamber are 6. Data analysis balanced and the CO2 concentration in the chamber is invariant (dc/dt = 0). From Eq. (2), the natural steady state CO2 concentration, 6.1. Regression analysis ss c(N), results in jN + jA + v For regression analysis, the terms v v cin and V were replaced by the parameters K1 and K2, respectively. Then, for the CO2 concentrations in the cave chamber, Eq. (4) yields jN jA ss = + + , − c(N) cin (5) = − − K2t, v v c K1 (K1 c0) e (7)

ss where K1 represents c and K2 represents 1/␶. Based on the visitor presence in the chamber, two main parts where all the symbols have their standard meaning. Eq. (4) indi- were distinguished on the anthropogenic CO2 peaks in the time ss cates that reaching the c(N) value is associated with decaying of data sets: (i) pulse leading edge (PLE); and, (ii) recession curve − v t the exponential term, e V at the time → ∞. Due to this uncer- (REC) (Fig. 6). PLE corresponds to the linear part of the CO2 con- tainty, a response time is defined as the time needed for substantial centrations increase (from c0(1) to the inflection) and represents v CO2 evolution during the visitor presence in the chamber. The sec- approach to the steady state. When t = V/v, the exponent V t is unity and e−1 corresponds to the value of ∼0.37. It means that the con- tions before PLE and between PLE and REC marked as transition centration reached 63% of the steady state value. Therefore, the sections (TS) represent the visitors entering the chamber and vis- response time, ␶ [s], corresponds to itors gradually leaving the chamber, respectively. Whereas at the TS1, the slope of CO2 concentrations gradually increased depending on the instantaneous visitor number in the chamber, the TS2 leads to the slope decrease and up to the complete slope inversion into ␶ = V . negative values. REC represents the period with no visitors: it leads (6) ss v from the inflection (c0(2)) to the natural CO2 steady state level, c(N).

110 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 45

Fig. 4. Monitoring Campaign III (6–7 June 2015; Bear Chamber, Vypustek´ Cave): visitor number per individual tour (the column widths represent the periods of visitor staying in the chamber) (a), CO2 concentrations (the red dashed lines represent hypothetical recession curves in the case of no further visitors) (b), and temperature difference T = Texterior − Tcave (c). See text for details. (For interpretation of the references to colour in this ﬁgure legend, the reader is referred to the web version of this article.)

During the periods between individual visitor tours, the cave CO2 concentrations are controlled by natural CO2 ﬂuxes and approach the natural steady state concentrations. Based on Equation (4), both the RECs and PLEs of each anthropogenic peak were analyzed separately by Recession analysis and Analysis of the pulse leading edges, respectively.

Fig. 6. Evolution of the cave CO2 concentrations during visitor presence (c0(1) and c are initial concentrations for the pulse modeling; css is natural steady state 0(2) (N) concentration; TS1 and TS2 are the transition sections). See text for details.

6.2. Analysis of the recession part of anthropogenic pulse

The hypothetical periods needed for cave relaxation, i.e., to Fig. 5. Conceptual model of CO2 dynamics in the cave chamber. reach the “natural” steady state level, were estimated based on the 111 46 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52

Table 2 Regression parameters of the recession curves (Bear Chamber, Vypustek´ Cave).

REC × −4 ␶ × −2 × −2 Tour number Visitor number K2 10 c0(2) 10 jN 10 v [s−1] [hours] [mol m−3] [mol s−1] [m3 s−1]

Campaign I #1 88 1.61 1.72 3.37 1.23 1.26 #2 58 1.14 2.43 3.57 0.87 0.89 #3 113 1.21 2.30 4.22 0.92 0.94 #4 79 1.29 2.16 4.48 0.98 1.00 #5 100 0.92 3.02 4.93 0.70 0.72 #6 111 0.75 3.73 5.32 0.57 0.58 #7 99 0.65 4.30 4.70 0.49 0.50 #8 93 0.82 3.40 5.06 0.63 0.64 #9 97 0.97 2.88 5.29 0.74 0.75 #10 100 1.51 1.84 3.76 1.16 1.18 #11 88 1.14 2.44 4.09 0.87 0.89 #12 99 1.01 2.75 4.51 0.77 0.79 #13 105 1.07 2.61 5.00 0.82 0.83 #14 95 0.87 3.19 5.26 0.67 0.68 #15 91 0.93 2.99 5.35 0.71 0.73 #16 87 0.76 3.64 5.42 0.58 0.60 #17 97 0.64 4.32 5.48 0.49 0.50

Campaign II #1 100 1.17 2.37 3.37 0.62 0.92 #2 82 0.93 2.99 3.64 0.49 0.73 #3 91 0.77 3.59 3.95 0.41 0.60 #4 75 0.97 2.87 4.11 0.51 0.75 #5 93 0.57 4.89 4.25 0.30 0.44 #6 89 0.69 4.05 4.41 0.36 0.54 #7 93 0.45 6.23 4.55 0.24 0.35 #8 111 0.50 5.54 5.11 0.27 0.39 #9 94 0.48 5.76 4.84 0.26 0.38 #10 92 0.97 2.86 3.40 0.52 0.76 #11 98 0.51 5.45 3.63 0.27 0.40 #12 104 0.64 4.34 3.99 0.34 0.50 #13 93 0.63 4.39 4.21 0.34 0.49 #14 90 0.90 3.10 4.54 0.48 0.70 #15 97 0.61 4.53 4.57 0.33 0.48 #16 101 0.70 3.98 4.84 0.37 0.54 #17 100 0.88 3.17 4.91 0.47 0.68

Campaign III #1 10 1.39 1.99 4.25 2.26 1.09 #2 87 3.48 0.80 4.62 5.64 2.71 #3 99 2.91 0.96 4.78 4.72 2.27 #4 74 1.89 1.47 4.78 3.07 1.48 #5 63 2.29 1.21 4.75 3.72 1.79 #6 71 1.48 1.87 4.83 2.41 1.16 #7 78 1.87 1.49 4.88 3.03 1.46 #8 52 1.63 1.71 4.91 2.64 1.27

KRECis css . 2 (N) ␶ is response time; c0(2) is initial concentration at regression; jN is natural CO2 flux; v is volumetric airflow. recession analysis. That is a well known method in hydrology airflow rates, v, varied in the ranges: 0.50–1.26 m3 s−1 (cam- (Furey and Gupta, 2000; Chapman, 2003; Chen and Krajewski, paign I), 0.35–0.92 m3 s−1 (campaign II), and 1.09–2.71 m3 s−1 2015) which was adopted for CO2 concentration evolution in this (campaign III). Then, the resulting ␶ values ranged from work. Note that the “natural” levels represent the CO2 concentra- 1.72 to 4.32 h (campaign I), from 2.37 to 6.23 h (cam- tions in the cave atmosphere in the periods without anthropogenic paign II), and from 0.80 to 1.99 h (campaign III) (Table 2). REC ± × −3 −1 influx. To simplify the regression analysis, the values of K1 (K1 Direct natural fluxes, jN, were (7.77 0.98) 10 mol s value derived from the recession part of the pulse) were esti- (campaign I), (3.86 ± 0.65) × 10−3 mol s−1 (campaign II), and mated directly from the data sets in the periods between the (3.44 ± 0.77) × 10−2 mol s−1 (campaign III). REC individual monitoring campaigns. Under given conditions, the K1 values were 2.70 × 10−2 mol m−3 (622 ppmv), 2.40 × 10−2 mol m−3 (553 ppmv), and 3.80 × 10−2 mol m−3 (876 ppmv) for the moni- REC toring campaigns I, II, and III, respectively. The K2 values (K2 6.3. Analysis of the leading edges of anthropogenic pulse value derived from the recession part of the pulse) were acquired by regression analysis; the regression parameters are given in The maximum attainable CO2 concentrations (anthropogenic Table 2. The modeled RECs based on the parameters are presented ss steady states, c(A)) when the visitors stayed for sufficiently long in Figs. 2–4, as the dashed red lines. period in the chamber were found by the analysis of PLEs. There REC × −5 The K2 values varied in the range from 6.43 10 to were expected invariant conditions in the chamber at both the PLE × −4 −1 × −5 × −4 −1 1.61 10 s (campaign I), 4.46 10 to 1.17 10 s and REC for the same pulse. Therefore, the values of KRECand v from × −4 −1 2 (campaign II), and (1.39–3.48) 10 s (campaign III). Based the REC analysis were used for the PLEs modeling. Then, the data on the fixed chamber volume, V ∼ 7800 m3, the volumetric PLE regression yielded the K1 values. The anthropogenic CO2 flux,

112 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 47

Table 3 Regression parameters from PLE analysis (Bear Chamber, Vypustek´ Cave).

PLE × −1 × −2 × −1 × −3 Tour number Visitor number K1 10 c0(1) 10 jA 10 jAP 10 [mol m−3] [mol m−3] [mol s−1] [mol s−1]

Campaign I #1 88 0.80 2.75 0.66 0.75 #2 58 0.80 3.21 0.47 0.81 #3 113 1.24 3.37 0.92 0.81 #4 79 1.11 3.83 0.84 1.06 #5 100 1.83 4.12 1.12 1.12 #6 111 1.93 4.67 0.96 0.87 #7 99 0.51 4.67 0.12 0.12 #8 93 1.56 4.44 0.82 0.88 #9 97 1.80 4.67 1.15 1.19 #10 100 1.05 2.98 0.93 0.93 #11 88 0.80 3.47 0.47 0.53 #12 99 1.05 3.86 0.62 0.62 #13 105 1.58 4.22 1.09 1.04 #14 95 1.35 4.61 0.74 0.78 #15 91 1.47 4.87 0.87 0.95 #16 87 1.33 5.00 0.63 0.73 #17 97 1.46 5.03 0.60 0.62

Campaign II #1 100 1.58 2.66 1.23 1.23 #2 82 1.71 3.20 1.07 1.30 #3 91 2.13 3.41 1.14 1.26 #4 75 1.27 3.73 0.78 1.04 #5 93 1.49 4.89 0.55 0.59 #6 89 2.62 4.03 1.27 1.43 #7 93 1.50 4.14 0.44 0.47 #8 111 1.42 4.22 0.46 0.42 #9 94 2.16 4.46 0.72 0.77 #10 92 1.68 2.83 1.09 1.18 #11 98 2.00 3.12 0.70 0.71 #12 104 2.65 3.45 1.20 1.16 #13 93 1.51 3.70 0.63 0.67 #14 90 1.13 3.82 0.62 0.69 #15 97 1.12 4.05 0.42 0.43 #16 101 2.01 4.37 0.96 0.95 #17 100 1.22 4.44 0.67 0.67

Campaign III #1 10 0.77 4.10 0.43 4.29 #2 87 0.57 4.08 0.53 0.60 #3 99 0.64 4.35 0.59 0.59 #4 74 0.81 4.37 0.63 0.86 #5 63 0.57 4.44 0.35 0.55 #6 71 0.63 4.48 0.29 0.41 #7 78 0.77 4.49 0.57 0.73 #8 52 0.64 4.51 0.33 0.63

KPLE is css . 1 (A) c0(1) is initial CO2 concentration; jA is total anthropogenic ﬂux; jAP is anthropogenic ﬂux per person.

−1 jA [mol s ], was calculated from differences between the upper 6.4. Theoretical anthropogenic steady states ss ss anthropogenic steady states,c(A), and natural steady states, c(N) as Based on Eqs. (5) and (6), the theoretical anthropogenic steady ss ␶ state CO2 concentration, c(A), and values were calculated for stan- = ss ss = PLE REC jA (c(A)-c(N))v (K1 -K1 )v. (8) dard cave tours covering 20 and 50 visitors under assumption that length of stay would not be limited. These visitor numbers correspond to standard visitor capacity of the individual tour in the The found regression parameters are presented in Table 3. ss Vypustek´ Cave. The c(A) values were calculated separately for UAF The KPLE values (representing css ) were found in a wide 1 (A) mode and DAF mode. The calculation was based on the parame- − − − range from 5.07 × 10 2 to 1.93 × 10 1 mol m 3 (1166 to ter values from the regression analysis (Tables 2 and 3). For the − − 4436 ppmv), (1.12–2.65) × 10 1 mol m 3 (2585 to 6106 ppmv), and calculation, the maxima and minima of cave airflows were used − − (5.74–8.10) × 10 2 mol m 3 (1323 to 1868 ppmv), for campaigns (0.35 and 1.26 m3 s−1 for UAF mode and 1.09 and 2.71 m3 s−1 for −3 −1 I, II, and III, respectively. Anthropogenic CO2 flux, jA, varied in DAF mode). The jN values of 5.81 × 10 mol s (UAF mode) and − − − the ranges: from 1.19 × 10 2 to 1.15 × 10 1 mol s 1 (campaign 3.44 × 10−2 mol s−1 (DAF mode) correspond to the mean values × −2 × −1 −1 I); from 4.21 10 to 1.27 10 mol s (campaign II); and, for individual ventilation modes. Anthropogenic CO2 fluxes, jA, − − (2.90–6.34) × 10 2 mol s 1 (campaign III). Based on the number of were calculated as function of visitor number and mean value of × −4 × −4 −1 × −3 −1 visitors, the individual personal flux, jAP, ranged from 1.20 10 jAP (8.47 10 mol s for UAF mode, 1.08 10 mol s for DAF − − − − to 1.19 × 10 3 mol s 1 person 1 (campaign I), from 4.17 × 10 4 mode). The resulting values are presented in Table 4. to 1.43 × 10−3 mol s−1 (campaign II), and from 4.09 × 10−4 to 4.29 × 10−3 mol s−1 person−1 (campaign III) (Table 3).

113 48 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52

Table 4 Table 6

Theoretical values of the steady state CO2 concentrations covering anthropogenic Theoretical staying periods for different visitor numbers that would be needed for ﬂuxes depending on the visitor number. balancing of the PCO2(H) in dripwater D2 with the peak anthropogenic PCO2(C) in cave atmosphere. Ventilation mode theoretical reachable css (A) Visitor number Staying period [hours] −3 [mol m ] [ppmv] log PCO2 UAF mode DAF mode 20 visitors UAF mode (3.53–8.26) × 10−2 813–1904 −3.09 to −2.72 20 18.12 10.60 DAF mode (3.78–6.87) × 10−2 873–1585 −3.06 to −2.80 50 7.25 4.24 100 3.63 2.12 50 visitors 500 0.73 0.42 × −1 − − UAF mode (0.56–1.56) 10 1278–3588 2.89 to 2.45 1000 0.36 0.21 × −2 − − DAF mode (4.98–9.86) 10 1149–2274 2.94 to 2.64 10,000 0.04 0.02

During the presence of tour with 20 visitors in the cave chamber, − − − − ss (∼7.89 × 10 4 mol s 1) and DAF mode (∼2.35 × 10 3 mol s 1) were the potential anthropogenic CO2 steady state c could reach up to (A) used (see Table 3). The resulting values of tour staying periods are 8.26 × 10−2 mol m−3 (i.e., up to 1904 ppmv or log P ∼ −2.72) at CO2 presented in Table 6. UAF mode and up to 6.87 × 10−2 mol m−3 (i.e., up to 1585 ppmv Considering different j values for individual modes, the val- or log P ∼ −2.8) at DAF mode. The presence of 50 visitors could AP CO2 ues of the periods for DAF mode exceed the values for UAF mode increase the theoretical CO steady state up to 1.56 × 10−1 mol m−3 2 almost by factor of 2. In the case of an ordinary cave tour in the (i.e., up to 3588 ppmv or log P ∼ −2.45) at UAF mode and CO2 Vypustek´ Cave (50 visitors), the required staying period was 7.25 h 9.86 × 10−2 mol m−3 (i.e., up to 2274 ppmv or log P ∼ −2.64) at CO2 (UAF mode) and 4.24 h (DAF mode). For the tours with enhanced DAF mode (Table 4). The ␶ values varied from 1.72 to 6.23 h during attendance (up to 500 visitors), the required staying period was UAF mode and from 0.80 to 1.99 h during DAF mode in dependence estimated up to 0.73 h (UAF mode) and 0.42 h (DAF mode). On the on actual airﬂows and independently on visitor number. other hand, the usual time of individual tour of 0.04 or 0.02 h would require the group of 10,000 visitors during the UAF and DAF mode, 6.5. Correlation analysis respectively.

A relation between individual parameters/variables was tested 7. Discussion by correlation analysis. The correlations between the initial CO2 ␶ concentrations used in REC, c0(2), and response time, , were strong (r = 0.81) for campaign I, weaker (r = 0.53) for campaign II, and The natural CO2 concentrations in the range of − − insignificant (r = 0.24) for campaign III. The correlations between (2.40–3.80) × 10 2 mol m 3 (553–876 ppmv) usual in the Bear the temperature difference, T, and the cave airflows, v, were Chamber indicate that the Vypustek´ Cave belongs to the caves insignificant for campaign I (r = 0.20) and for campaign II (r = 0.19), with rather lower CO2 levels. These values are comparable with and moderately strong for campaign III (r = 0.69). The correlations the values in some other Moravian Karst caves (Faimon, Stelcl,ˇ between attendance and anthropogenic flux, jA, were weak for Sas et al., 2006; Faimon, Troppová et al., 2012; Lang et al., 2015a, campaign I (r = 0.52), insignificant for campaign II (r = −0.16), and 2015b, 2016) or with some caves worldwide, see e.g., Austrian moderately strong for campaign III (r = 0.62). All the correlations Obir Cave (Spötl et al., 2005), Irish Ballynamintra Cave (Baldini, are summarized in Table 5. Baldini, McDermott, & Clipson, 2006), Spanish Nerja Cave (Linán˜ et al., 2008), or Pozalagua Cave (Lario and Soler, 2010). On the 6.6. Analysis of dripwater hydrogeochemistry other hand, the Vypustek´ Cave values are substantially lower in comparison with the values of 0.26 mol m−3 (6000 ppmv) Both dripwaters show permanent supersaturation with respect reported by Sánchez-Moral et al. (1999) for the Altamira Cave (Spain), 0.36 mol m−3 (8300 ppmv) referred by Ek and Gewelt to calcite with higher values for D1 (see the SI(calcite) values in (1985) for the Ste-Anne Cave (Belgium), extreme 1.80 mol m−3 Table 1). Two CO2 partial pressures were calculated for each water: (41,500 ppmv) presented by Bourges et al. (2001) for the Aven one corresponding to aqueous carbonate species, PCO2(W), and −3 hypothetical one influencing the water formation in soils/epikarst, d’Orgnac Cave (France), or even 2.68 mol m (62,000 ppmv) given by Batiot-Guilhe et al. (2007) for the Causse dıAumelas´ (France). PCO2(H) (the PCO2 in equilibrium with both aqueous carbonate species and calcite, see for details Faimon, Licbinskáˇ et al., 2012 and The proposed model presumes that the natural CO2 levels in the cave (the levels without any anthropogenic flux) are con- Pracny,´ Faimon, Kabelka, & Hebelka, 2015). Both calculated PCO2 are lower for dripwater D2 and for winter period (Table 1). Therefore, trolled by (i) the natural CO2 fluxes from soils/epikarst (diffusional, additional considerations are mostly related to D2. advective, and water degassing fluxes) and (ii) the advective CO2 fluxes out of cave atmosphere driven by cave airflows. The natu- Based on the values of individual personal anthropogenic CO2 ral CO flux into the Bear Chamber obtained from data modeling, flux, jAP, the staying periods of visitors needed for balancing 2 j , showed slight seasonality: whereas j values varied in the the PCO2(H) values of dripwater D2 with the PCO2(C) values in N N × −2 −1 cave atmosphere were calculated for different number of visi- range of (0.24–1.23) 10 mol s during fall/winter at UAF mode (campaigns I and II), they ranged in (2.26–5.64) × 10−2 mol s−1 tors. For the calculation, the mean jAP values during UAF mode interval during summer at DAF mode (campaign III). Based on the orthogonal projection plane of about 1100 m2 of the Bear Table 5 Chamber, the normalized mean specific natural flux per 1 m2 Correlations between selected variables. was 6.64 × 10−6 and 3.58 × 10−5 mol m−2 s−1 for UAF and DAF Variables Campaign I Campaign II Campaign III mode, respectively. These values are roughly consistent with −7 −2 −1 −5 −2 −1 c0 vs. ␶ 0.81 0.53 0.24 6.20 × 10 mol m s (UAF mode) and 1.24 × 10 mol m s T vs. v 0.20 0.19 0.69 (DAF mode) presented by Milanolo and Gabrovsekˇ (2009). In con- − A vs. jA 0.52 0.16 0.62 trast, the values are much higher in comparison with the mean The highlighted correlations are significant at ␣ = 0.05. values of 7.59 × 10−8 mol m−2 s−1 estimated for the Císarskáˇ Cave

114 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 49

−9 −2 −1 (Faimon, Stelcl,ˇ Sas et al., 2006) or 5.36 × 10 mol m s for Bal- sufficiently long period, the theoretical CO2 steady state val- ss × −1 −3 carka Cave (Lang et al., 2015b). It is important to note that such ues, c(A), could reach 2.65 10 mol m (6106 ppmv) during normalization is very simplified. It is based on an assumption that UAF mode and 8.10 × 10−2 mol m−3 (1868 ppmv) during DAF the input fluxes are distributed homogeneously across the cham- mode (Table 3). Note that these values correspond to the tours ber. In addition, it ignores the input advective fluxes from adjacent with extreme visitor numbers. If an “ordinary” cave tour is cave spaces, the projection plane area of which is outside the Bear ∼ ss assumed ( 50 people), the resulting c(A) values could reach Chamber (see Lang et al., 2015a). − − up to 1.56 × 10 1 mol m 3 (3588 ppmv) during UAF mode and The output advective CO fluxes out of the chamber atmo- − − 2 0.99 × 10 1 mol m 3 (2274 ppmv) during DAF mode (see Table 4). sphere, 6.02 × 10−3–4.66 × 10−2 mol s−1, are roughly comparable For comparison, the theoretical anthropogenic concentra- with other CO fluxes in the modeled system (Tables 2 and 3). These 2 tions css were estimated also for groups of 20 visitors values are slightly higher in comparison with peak values of advec- (A) representing the lower limit of the individual tour in the tive fluxes estimated by Lang et al. (2015a), and Lang et al. (2015b) Vypustek´ Cave. In such cases, the resulting css values would for the Balcarka Cave. The advective fluxes are a function of (1) the (A) reach 0.83 × 10−1 mol m−3 (1904 ppmv) during UAF mode and CO2 concentrations in the external/cave atmosphere, cin, and (2) 0.69 × 10−1 mol m−3 (1585 ppmv) during DAF mode (see Table 4). It the cave airflows controlled by temperature difference T = Text − × −2 −3 means that the sufficiently long presence of visitors could increase Tcave. Based on the invariant cin value (about 1.72 10 mol m , i.e., ∼400 ppmv), the advective fluxes are controlled by volumet- the natural CO2 levels in the chamber by factor of two during DAF ric velocity of the airflow through the cave chamber. Similarly for mode and by factor of ten during UAF mode. However, the presence of 20 visitors could exceed the natural CO levels in the chamber jN values, the airflow values estimated by modeling showed slight 2 seasonality. Whereas during campaigns I and II (corresponding by factor of 1.8 (DAF mode) or 3.2 (UAF mode). to UAF mode) the values varied from 0.35 to 1.25 m3 s−1, during To quantify the period necessary for cave relaxation, the cave ␶ campaign III (corresponding to DAF mode) the values varied from response time, , was proposed. It represents the time that the cave 1.09 to 2.71 m3 s−1 (Table 2). The higher airflows during DAF mode needs to reinstate natural conditions after the last visitors have left. appear contradictory to the paradigm that winter airflows exceed If the cave tours with enhanced visitor numbers are assumed, the ␶ summer airflows (Bourges, Mangin, & d’Hulst, 2001; Kowalczk resulting values varied between 1.72 and 6.23 h during UAF mode and Froelich, 2010; Duenas,˜ Fernández, Canete,˜ Pérez, & Gordo, (campaigns I and II) and between 0.80 and 1.99 h during DAF mode 2011; Faimon, Troppová et al., 2012; Faimon and Lang, 2013; Lang, (campaign III) (Table 2). These values are roughly consistent with Faimon, Godissart, & Ek, 2016). In this case, however, the higher the mean value of 4.4 h presented by Lobo (2015) for the Santana ␶ airflows during DAF mode might be a consequence of the higher Cave (Brazil). Note that represents time after which only 63% of − ss |DT| values representing driving force of cave airflows. Whereas the the excessive CO2 (c0(2) c(N)) is vented out (Fig. 7). Based on Eq. (6), maximum of |T| during DAF mode reached up to 21 ◦C(Fig. 4c), the ␶ values can vary depending on cave airflow v under assump- the |T| maxima during UAF mode were much lower: 6.7 and 8.1 ◦C tion that the chamber volume V is invariant. It means that the ␶ for campaign I (Fig. 2c), 9.0 and 10.5 ◦C for campaign II (Fig. 3c). values are independent on the number of visitors. However, the Nevertheless, the correlation analysis showed just weak depen- visitor number influences the c0(2) values on recession curve and, dence between T and cave airflows for both UAF and DAF modes thus, the response time ␶ represents different final CO2 concentra- (Table 5). The insignificance of the higher correlation coefficient tions. The relation between c0(2) values and periods required for for campaign III is a consequence of the smaller data population cave relaxation is presented in Fig. 7. Based on Eq. (4), two reces- (the CO2 peaks were identified only during the first visitor event of sion curves with the same parameters were simulated for different A B A campaign III). c0(2) values (c0 and c0 ). Whereas c0 corresponds to the cave tour B An anthropogenic impact on the natural CO2 concentrations in with enhanced visitor number, c0 represents an “ordinary” cave the Bear Chamber is evident. It is represented by peaks super- tour. As can be seen, the shapes of both curves are the same. How- A B imposed on the roughly smooth curve of “natural” CO2 levels ever, c0 and c0 are in different positions on the curve. For the same (Figs. 2b, 3b and 4b). In general, the height of a peak should time (␶), 63% of the excessive CO2 was vented out, but the final CO2 A B correspond to (i) the number of visitors in individual tour and concentrations c0 and c0 were different (Fig. 7). To reach the same (ii) staying period in the chamber. Nevertheless, the correlation B final CO2 concentrations as in case of c0 , much longer time period analysis showed insignificant or moderate correlations (r ∼ −0.16 A is required for c0 . to 0.62) (Table 5). This indicates more complicated relation An important factor influencing the response time of CO2 in cave between anthropogenic fluxes, jA, and visitor number. According atmosphere is the ventilation period (Faimon, Troppová et al., 2012; to attendance, the anthropogenic CO2 fluxes into the chamber, Lang et al., 2015a, 2015b). The ␶ values were determined for the jA, correspond to the mean CO2 flux related to one person of period of active ventilation (see the temperature differences in Figs. ∼ × −3 −1 −1 jAP 2.21 10 mol s person (Table 3). It is consistent with 2c, 3c, 4c). In the period of limited ventilation (typically in the spring × −3 −1 −1 the value of 1.49 10 mol s person reported by Dragovich and fall when external temperature approaches the temperature in and Grose (1990) for the Jenolan Caves (Australia). However, cave), much longer ␶ would be necessary for returning CO2 levels this flux belongs to higher values in comparison with the fluxes to the natural values due to lower and more varying airflow v. × −5 × −4 −1 −1 between 5.35 10 and 2.90 10 mol s person found in Principally, two possibilities exist for the cave management to other Moravian Karst Caves (Faimon, Stelcl,ˇ Sas et al., 2006; Lang maintain near-natural conditions: (1) to reduce number of visitors et al., 2015a, 2015b) or Srednja Bijambarska Cave in Bosnia and in individual tours; and, (2) to introduce sufficiently long pauses Herzegovina (Milanolo and Gabrovsek,ˇ 2009). Differences in the between individual groups. The intervals between individual tours jAP values in caves may be result of various factors such as visitors’ needed for large groups (up to 500 visitors) in the Vypustek´ Cave activity (Iwamoto, Pendergast, Suzuki, & Krasney, 1994), gender reach up to 6 h. In cases of ordinary groups (∼50 visitors), the (Sciacca et al., 2002), and age (Tormo, Bertaccini, Conde, Infante, & required intervals would correspond to significantly lower values. Cura, 2001). These values are in a good agreement with the intervals suggested The modeling allowed estimating the theoretically achiev- by Calaforra et al. (2003) for the Cueva del Agua Cave (4–5 h) or ss able anthropogenic concentrations c(A) in dependence on the Hoyos et al. (1998) for the Candamo Cave (7 h). Such a regime visitor number. If the visitors stay in the chamber for a would be insufficient for the period of limited ventilation or for

115 50 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52

−3 −1 Fig. 7. Inﬂuence of the CO2 relaxation period after visitors leave the chamber on the values of the theoretical CO2 levels. The parameters used: jN ∼ 5.81 × 10 mol s , ∼ 3 −1 ∼ × −2 −3 ∼ 3 A ∼ × −2 −3 B∼ × −2 −3 v 1.20 m s , cin 1.72 10 mol m (400 ppmv), V 7800 m , c0 4.50 10 mol m (1040 ppmv), c0 3.50 10 mol m (810 ppmv).

the extreme visitor numbers frequent in some caves, e.g., in the The modeled anthropogenic CO2 steady state maxima, ss ∼ × −1 −3 ∼ −2.45 Punkva Caves in Moravian Karst, where the annual attendance ordi- c(A) 1.56 10 mol m (i.e., 3588 ppmv or PCO2 10 ) for narily exceeds 200,000 people. In the Vypustek´ Cave, the intervals ordinary visitor group (50 visitors) are clearly lower than the PCO2(H) between individual tours are presently set to ∼1 h (Figs. 2a, 3a and (Table 1). Under such conditions, the dripwater cannot become 4a), which is a lower value in comparison with the values given by unsaturated (cannot become aggressive to calcite). However, if a recession analysis (Table 2). larger group of visitors stayed in the cave for a long time, the ss − − The impact of visitor numbers during tours on the cave CO2 con- ∼ × 1 3 anthropogenic steady state value of c(A) 2.65 10 mol m (i.e., centrations is clearly visible in campaign III (Fig. 4a,b). Whereas ∼ −2.22 6106 ppmv or PCO2 10 ) could approach the dripwater PCO2(H). cave tours of more than 50 people during ﬁrst tour event induced Therefore, under extreme conditions, the anthropogenic concen- recognizable CO2 peaks, an impact of cave tours with a maximum trations could potentially reach the values, at which water will of 15 people during second tour event is almost unrecognizable. become unsaturated and aggressive to calcite. Especially, dripwa- Therefore, the optimal visitor number per one tour in the Vypustek´ −2.27 −2.51 ter D2 showing the lower PCO2(H) values (10 and 10 , see Cave is up to 20 people in order to preserve the natural cave con- Table 1) is potentially sensitive to the increase in aggressiveness to ditions. Similar limit was suggested by Calaforra et al. (2003), who calcite by anthropogenic CO2 inﬂuence. In fact, D2 is rather atypical recommended the visitor regime in the Cueva del Agua Cave not to water as the PCO2(H) does not reach values of D1 nor the PCO2(H) val- exceed 53 people. However, these recommendations are incompa- ∼ −1.50 ues common in the Moravian Karst (PCO2(H) 10 ; see Faimon, rable with the visitor capacity of 29 visitors per day proposed by Licbinskáˇ et al., 2012 or Pracny,´ Faimon, Kabelka et al., 2015). Hoyos et al. (1998) for Candamo Cave. To quantify the risk for different visitor numbers, the theoretical The problem is that any limitations of the visitor number and staying periods were calculated, at which PCO2(H) would be bal- the relaxation periods between individual tours represent a com- anced by PCO2(C). The results showed periods of more than 7 h (UAF plication for the administration of tourist activities. Therefore, a mode) and 4 h (DAF mode) for a standard group size and 0.73 h risk analysis was implemented for the event when the natural con- (UAF mode) and 0.42 h (DAF mode) for larger groups (up to 500 vis- ditions would not be met. As the highest risk of anthropogenic itors) (Table 6). As can be seen, the values for an ordinary group in CO 2 is associated with a potential increase of dripwater aggres- both modes clearly exceed the duration of standard tour in the Bear siveness to calcite speleothem, the analysis is focused on comparing Chamber (∼0.5 h). If a tour with very high visitor numbers (up to the anthropogenic PCO2 with dripwater hydrogeochemistry. For the 10,000 visitors) would be present in the cave, the period needed for risk analysis, the dripwater entering the cave was assumed to be balancing PCO2(H) and PCO2(C) is only 0.04 h (UAF mode) and 0.02 h at equilibrium with both epikarstic CO2 (represented by PCO2(H)) (DAF mode). However, such scenario is unrealizable as such visitor and calcite. This equilibrium is broken when the water enters number exceeds the cave capacity. The former implications indicate the cave and interacts with lower PCO2(C) in cave atmosphere. A that the presence of standard tour group cannot lead to the conver- new equilibrium should be achieved by (i) degassing excessive sion of dripwater into solution aggressive to calcite. The situation CO2 and (ii) precipitating calcite. The dynamics of degassing is will drastically change in case of the water condensed from cave given by the difference between PCO2(W) and PCO2(C). The initial atmosphere inducing so-called condensation corrosion (Dreybrodt, value of PCO2(W) equals to PCO2(H) but it quickly evolves towards Gabrovsek,ˇ & Perne, 2005; Gabrovsek,ˇ Dreybrodt, & Perne, 2010). the PCO2(C) value due to a partial degassing. Both dripwaters stud- This phenomenon was well documented by many studies (e.g., ied showed partial degassing as the PCO2(W) is lower than PCO2(H) Sarbu and Lascu, 1997; Dublyansky and Dublyansky, 1998; Tarhule- (Table 1). Concurrently with CO2 degassing, the water supersat- Lips and Ford, 1998; de Freitas and Schmekal, 2005, 2006). The uration with respect to calcite increases. If PCO2(C) abruptly rises principle of the phenomenon is that the water condensed from due to the anthropogenic impact and exceeds PCO2(W), atmospheric atmosphere shows low alkalinity and becomes increasingly acidic CO2 would dissolve in water and the water supersaturation would during gaseous CO2 dissolution. In this case, any increase of CO2 decrease. Only if PCO2(C) exceeded PCO2(H), the atmospheric CO2 concentration in cave atmosphere causes increase in aggressive- would dissolve up to the point where water becomes unsaturated ness of the water to calcite, which results in speleothem corrosion. with respect to calcite, i.e., becomes aggressive to calcite (see, e.g., Therefore, despite the resistance of dripwaters against a transfer Palmer and Palmer, 2013). into undersaturation with respect to calcite, the anthropogenic

116 M. Lang et al. / Journal for Nature Conservation 35 (2017) 40–52 51

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