Acta carsologica

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Acta carsologica

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Acta carsologica
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Acta Carsologica
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Krasoslovni zbornik
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Inštitut za raziskovanje krasa (Slovenska akademija znanosti in umetnosti)
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Vol. 42, no. 2-3 (2013)
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Introduction to the symposium ( .pdf )

Carbon fluxes in Karst aquifers: Sources, sinks, and the effect of storm flow / William B. White ( .pdf )

Do carbonate karst terrains affect the global carbon cycle? / Jonathan B. Martin, Amy Brown, John Ezell ( .pdf )

Quaternary glacial cycles: Karst processes and the global CO2 budget / Erik B. Larson, John E. Mylroie ( .pdf )

A framework for assessing the role of karst conduit morphology, hydrology, and evolution in the transport and storage of carbon and associated sediments / George Veni ( .pdf )

Biological Control on Acid Generation at the Conduit-Bedrock Boundary in Submerged Caves: Quantification through Geochemical Modeling / Janet S. Herman, Alexandria G. Hounshell, Rima B. Franklin, Aaron L. Mills ( .pdf )

An approach for collection of nearfield groundwater samples in submerged limestone caverns / Aaron L. Mills, Terrence N. Tysall, Janet S. Herman ( .pdf )

Organic matter flux in the epikarst of the Dorvan karst, France / Kevin S. Simon ( .pdf )

Environmental controls on organic matter production and transport across surface-subsurface and geochemical boundaries in the Edwards aquifer, Texas, USA / Benjamin T. Hutchins, Benjamin F. Schwartz, Annette S. Engel ( .pdf )

Using isotopes of dissolved inorganic carbon species and water to separate sources of recharge in a cave spring, northwestern Arkansas, USA Blowing Spring Cave / Katherine J. Knierim, Erik Pollock, Phillip D. Hays ( .pdf )

Isotopes of Carbon in a Karst Aquifer of the Cumberland Plateau of Kentucky, USA / Lee J. Florea ( .pdf )

Organic Carbon in Shallow Subterranean Habitats / Tanja Pipan, David C. Culver ( .pdf )

Contribution of non-troglobiotic terrestrial invertebrates to carbon input in hypogean habitats / Tone Novak, Franc Janžekovič, Saška Lipovšček ( .pdf )

Physical Structure of the Epikarst / William K. Jones ( .pdf )

Using hydrogeochemical and ecohydrologic responses to understand epikarst process in semi-arid systems, Edwards plateau, Texas, USA / Benjamin F. Schwartz, Susanne Schwinning, Brett Gerrard, Kelly R. Kukowski, Chasity L. Stinson, Heather C. Dammeyer ( .pdf )

Variability of groundwater flow and transport processes in karst under different hydrologic conditions / Nataša Ravbar ( .pdf )

Spring discharge records - a case study / Carol M. Wick ( .pdf )

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O RGANIC CARBON IN SHALLO W SUBTERRANEAN HABITATS O RGANSKI OGLJIK V PLITVIH PODZEMELJSKIH HABITATIH Tanja P IPAN 1 & David C. CULVER 2 Izvleek UDK 551.44:546.26 Tanja Pipan & David C. Culver: Organski ogljik v plitvih podzemeljskih habitatih Organski ogljik je najverjetneje omejujoi dejavnik v plitvih podzemeljskih habitatih (SSH). Analizirali smo podatke o raztopljenem organskem ogljiku (DOC) v treh SSH: (1) hipo telminorejik z mezii (Nanos, Slovenija), (2) hiporejik (reka Rona, Francija in manji potoki na vznoju Nanosa), ter (3) epikras (Kitajska, Slovenija in ZDA). Hipotelminorejini habi tati so mesta s podzemeljsko vodo v globini do enega metra, ki primezi na povrje iz drobnih izvirov mezi. Hiporejini habitati so podzemeljski tokovi potokov in rek. Epikras je pre del tik pod povrjem krasa, prepreden s tevilnimi majhnimi votli nami in kanali. Povprene vrednosti organskega ogljika v hipotelminorejinih habitatih s stigobiontskimi vrstami na Nanosu so znaale 3,4 mg C L in so v razlinih obdo bjih mono nihale. V hiporejiku Rone in potokov na Nanosu so srednje vrednosti znaale med 1,4 in 3,5 mg C L V dobro proue nem poreju Rone je bila asovna variabilnost majhna. Vrednosti DOC v epikrasu v treh jamah se gibljejo med 0,70 do 1,10 mg C L (jama Shihua na Kitajskem, PostojnskoPlaninski jamski sistem in jama Organ v Zahodni Virginiji, ZDA). Iz re zultatov sklepamo, da so vrednosti organskega ogljika v vodnih SSH najnije v epikrasu. Kljune besede : raztopljeni organski ogljik, epikras, hiporejik, hipotelminorejik, mezie. 1 Karst Research Institute ZRC SAZU, Titov trg 2, SI-6230 Postojna, Slovenia, e-mail: pipan@zrc-sazu.si 2 Department of Environmental Science, American University, 4400 Massachusetts Ave. NW W ashington, DC 20016, USA, e-mail: dculver@american.edu Received/Prejeto: 15.1.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 291, POSTOJNA 2013 Abstract UDC 551.44:546.26 Tanja Pipan & David C. Culver: Organic carbon in shallow subterranean habitats Organic carbon is likely to be a limiting factor in shallow sub terranean habitats (SSHs). Data on dissolved organic carbon (DOC) in three SSHs are reviewed: (1) hypotelminorheic and associated seepage springs (Nanos Mountain, Slovenia), (2) hyporheic zones (Rhne River, France and seepage streams on Nanos Mountain, Slovenia), and (3) epikarst (China, Slovenia, and USA). Hypotelminorheic habitats are supercial ground water sites less than 1 m below the surface that exit from seep age springs. Hyporheic habitats are the underow of streams and rivers. Epikarst is the uppermost zone of karst with exten sive small cavities and channels. Nanos hypotelminorheic sites that harbored stygobiotic species had organic carbon values averaging 3.4 mg C L and temporal variability was high. For hypoheic sites in the Rhne River basin and on Nanos Moun tain, mean values ranged from 1.4 to 3.5 mg C L In the more extensively studied Rhne River basin sites, temporal variabil ity was low. Epikarst DOC ranged from 0.70 to 1.10 mg C L in three caves in China (Shihua Cave), Slovenia (Postojna Planina Cave System) and United States (Organ Cave, W est Virginia). ese results suggest that organic carbon in aquatic SSHs is lowest in epikarst. Keywords : dissolved organic carbon, epikarst, hyporheic, hy potelminorheic, seepage spring. I NTRODUCTION Shallow Subterranean Habitats (SSHs) are aphotic envi ronments that are less than 10 m below the surface (Cul ver & Pipan 2008, 2011). ey include hypotelminorheic and associated seepage springs, hyporheic, and epikarst

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ACTA CARSOLOGICA 42/2-3 2013 292 (dened below). SSHs, because of their close connec tions with the surface, are typically intermediate in en vironmental variability and in average parameter values between surface and deep subterranean habitats (Culver & Pipan 2008, 2011). ese patterns are largely based on extensive temperature records for both terrestrial (Pipan et al. 2011) and aquatic (Culver & Pipan 2011) SSHs. As a very general rule, productivity in surface ter restrial habitats is limited by availability of usable nitro gen (xed nitrogen) and productivity in surface aquatic habitats is limited by availability of phosphorus. By con trast, productivity in deep subterranean habitats (>10 m deep), especially cave streams (Simon & Beneld 2002; Huntsman, Venarsky & Bernstead 2011; Venarsky, Bern stead & Huryn 2012), seems to be controlled by the amount of organic carbon, with bottom-up control of the detrititus-based community. Shallow subterra nean habitats, with the exception of the hyporheic (see e.g. Danielopol et al. 2000; Datry et al. 2005; Marmonier et al. 2000; Vervier et al. 1992), have been little studied in this regard, and little is known about the amount of organic carbon available in SSHs. In this study, we report on dissolved organic carbon values in a variety of SSHs present on Nanos Mountain and above nearby Postojna Planina Cave System in Slo venia. W e compare these results with other studies of or ganic carbon in SSHs, and propose that organic carbon, especially dissolved organic carbon, is likely a limiting factor in many aquatic SSHs. A R EVIE W OF AQUATIC SSH s Culver and Pipan (2011) propose that aquatic SSHs share the following features: surface; subterranean habitats; Portions of most caves are in the zone of less than 10 m from the surface, but the habitats typically extend deeper and so are not included. Exceptions are lava tubes which are always quite shallow (Palmer 2007), but lava tubes rarely have any water because of the high porosity of the lava in which they occur. e three major aquatic SSHs are (1) the hypotelminorheic and associated seep age springs, (2) epikarst, and (3) hyporheic ows associ ated with rivers and streams. HYPOTELMINORHEIC e Croatian speleobiologist, Milan Metrov (19292010) applied the term hypotelminorheic to shallow groundwater habitats that are vertically isolated from the water table and are constituted of humid soils in the mountains, rich in organic matter and traversed by moving water (Metrov 1962, 1964). He included cases where the habitat was in close proximity to caves (he gave an example from Moulis, France) and an ex ample of where the outlet of a hypotelminorheic habi tat was the beginning of a mountain stream in Risnjak, Croatia. Elaborating on Metrovs (1962) denition, Culver et al. (2006) proposed that the term hypotelminorheic be used to describe habitats with the following features (see also Culver & Pipan 2008): 1. A perched aquifer fed by subsurface water that creates a persistent wet spot; 2. Underlain by a clay or other impermeable layer typically 5 to 50 cm below the ground surface; 3. Rich in organic matter compared with other aquatic subterranean habitats. Culver et al. (2006) also indicated that the drain age area of a seepage spring is typically less than 1 ha, in a shallow depression, and that the leaves in the seepage spring are characteristically blackened and not skeleton ized. W ithout a clay layer, water should tend to move vertically, and there would be no persistent water. Clay is a critical component of hypotelminorheic habitats, not only because it acts as a barrier to the downward movement of water, but also because during periods of drought, the water retained by the colloidal clay may serve as a refuge for invertebrates in the hypotelminor heic, into which they can burrow (Holsinger & Dickson 1977). According to Ginet and Decu (1977), clay and its associated bacteria and organic matter may also have nu tritional value for subterranean crustaceans. e water exits at a seepage spring, although there may not be ow at all times during the year. For most species associated with the hypotelminor heic, it is the subterranean water of the hypotelminorheic and not the seepage spring, i.e., the groundwater/surface water ecotone (see Gibert 1991), that is their primary habitat. e seepage spring is the point of collecting most of the hypotelminorheic fauna, but is not the shallow groundwater habitat itself, but a few species are primar T ANJA PIPAN & D AVID C. C ULVER

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ACTA CARSOLOGICA 42/2-3 2013 293 ily inhabitants of the ecotone itself (see below). Many of the species found in the hypotelminorheic are hypotelmi norheic specialists with the morphology typical of sub terranean species, including loss of eyes and pigment and elongation of appendages (Culver & Pipan 2011; Pipan & Culver 2012). Micro-crustaceans are little studied in these habitats, but the macrofauna is dominated by amphipods (e.g. Culver, Holsinger & Feller 2012). e hypotelminor heic and the associated seepage spring ecotone are clearly an example of a groundwater dependent ecosystem (Ea mus & Froend 2006). Seepage springs are also isolated wetlands, although a highly miniaturized ones. Photo graphs of several seepage springs are shown in Fig. 1. E PIKARST Epikarst is the uppermost part of karst, where stress re lease, climate, tree roots, and karst processes fracture and enlarge rock joints and cracks, creating a more porous zone over the carbonate rock in which only a few verti cal joints and cracks occur (Bakalowicz 2012). is zone, the epikarst, is important as a site of cave formation, water storage, and a habitat for many species. Typically 3 to 10 m thick, epikarst overlies the water inltration zone, which is itself intersected by occasionally enlarged vertical fractures and conduits (Fig. 2). Because of this, the base of the epikarst acts as an aquitard, a layer of low permeability, resulting in a local perched water table and a perched aquifer. According to W illiams (2008) the typi cal porosity (per cent open space) of unweathered lime stone is 2 percent while that of epikarst typically exceeds 20 percent. More generally, water storage in epikarst is the reason why cave streams typically have water for long periods of drought. e connection of outow from epikarst, which is measured by the output from drips in caves coming from epikarst, has a complex connection with precipitation Kogovek (2010). Typically, output spikes aer several precipitation events, which cumulatively ll the cavities in epikarst. Based on continuous monitoring of three epikarst drips for three years in Postojnska jama (Slove nia), Kogovek (2010) was able to estimate total surface catchment area of an individual drip using precipitation and drip rate data. e largest catchment area of a drip was quite small, approximately 200 m 2 Epikarst harbors a rich aquatic fauna (Pipan 2005; Culver et al. 2012), much of it limited to subterranean Fig. 1: P hotographs of seepage springs: leseepage spring at Medvenica Mountain, the site of the rst description of the habi tat by Metrov (1962), photograph by D. Culver; rightseepage spring on Nanos Mountain, Slovenia, photograph by J Mulec, used with permission; rightseepage spring in Scott Run Re gional P ark, V irginia, USA, photograph by W.K. J ones, used with permission. O RGANIC CARBON IN SHALLO W SUBTERRANEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 294 Fig. 2: Upper panel. Sketch of epikarst, from P ipan (2005), used with permission of ZRC P ublishing. Lower panel. P hotograph of exposed epikarst on the sawed face of a quarry in the Salem Limestone, Lawrence County, Indiana, USA. From P almer (2004) and used with permission of the Karst Waters Institute. Fig. 3: e River Ain near its mouth on the Rhne River. It is a very species rich site and has numerous gravel bars and banks. P hotograph by T. P ipan Fig. 4: Conceptual cross-sectional models of surface channels and beds showing relationships of channel water to hyporheic, ground water, and impermeable zones. ick arrows indicate direction of water uxes. (a) No hyporheic zone. (b) A hyporheic zone created only by advected channel water. (c) A hyporheic zone created by advection from both channel water and ground water. (d) A hy porheic zone created by inltration of channel water beneath the stream bed (no parauvial ow). (e) A perched hyporheic zone created only by inltration of channel water beneath the stream bed. From Malard et al. (2002). T ANJA PIPAN & D AVID C. C ULVER

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ACTA CARSOLOGICA 42/2-3 2013 295 METHODS AND MATERIALS STUDY S ITES Nanos is largely ysch mountain in the Kras (Karst in English) region of Slovenia. It ranges in elevation from 1262 m at the summit to approximately 550 m at its base near the village of Razdrto. On the southern ank of the mountain, near Razdrto, we located a series of three seepage springs, one small stream with a hypor heic zone, two small springs (both gushets sensu Krei 2010) near the headwaters of a stream, and a small sur face stream. Seepage springs were separated into two Niphargus a common inhabitant of European ground water and those that lack this amphipod genus. ese sites were within 600 m of each other and ranged in ele vation from 600 to 750 m. Samples were taken monthly during 2011. SAMPLE ANALYSIS W ater was collected in a 60 cc syringe and passed through a 0.45 mm glass ber lter (Gelman GF/F). W ater sam ples were ltered into an acid washed 50 mL HDPE sample bottle and preserved to pH 2 with two drops of concentrated HCl. e samples were analyzed for dis solved organic carbon (DOC) by concentration using the persulfate digestion method on a ermo Scientic Hi PerTOC Analyzer. Each sample was analyzed ve times, and the mean value of these is what is reported. Statistical dierences were analyzed using ANOVA. habitats, and oen limited just to epikarst habitats (Pipan & Culver 2007). Numerically, the fauna is dominated by copepods and the macroscopic fauna is dominated by amphipods and isopods. HYPORHEIC e hyporheic zone, the best studied by far of all inter stitial habitats, is the surface-subsurface hydrological ex change zone beneath and alongside the channels of rivers and streams. A more technical denition is provided by Krause et al. (2011): A temporally and spatially dynamic saturated tran sition zone between surface and ground water bodies that derives its specic physical (e.g., water temperature) and biogeochemical (e.g., steep chemical gradients) characteristics from mixing of surfaceand ground wa ter to provide a dynamic habitat and potential refugia for obligate and facultative species. e hyporheic of rivers is an ecotone between sur face and groundwater (Fig. 3). e hyporheic zone may or may not have permanent groundwater (phreatic wa ter) connections (Fig. 4). In rare cases, hyporheic and groundwater are absent, and the stream ows on imper meable rock (Fig. 4a). In all other cases, there is water ow vertically and laterally from channel water and/or groundwater. W hen there are unconsolidated sediments along the stream bank, the hyporheic can extend tens of meters from the stream bank. One hyporheic type (Fig. 4e) is a rather curious one isolated from groundwa ter. On a small scale, it closely resembles the hypotelmi norheic. Transfers of matter and organisms through and into the hyporheic depend on the permeability of the sub strate. If sediment load increases, ne particles can settle in the gravel bed reducing permeability, or slightly large particles can form a subsurface seal. In urban streams, such as Rock Creek in W ashington, D.C., sediment load from storm water runo is so heavy that nearly all pores of the hyporheic zone of some tributary streams are com pletely clogged and sealed (Culver & ereg 2004). e hyporheic harbors a diverse and rich fau na, including subterranean species, species (typically Ephemeroptera and Plecoptera) that spend most of their lives in the hyportheic except for a brief period as winged adults (Gibert et al. 1994), and species also found in streams and rivers. RESULTS Estimates of dissolved organic carbon are shown in detail in Tab. 1 and summarized in Fig. 5. ere were no statisti cally signicant dierences among the sites (ANOVA), in part because of small sample sizes, but especially because of the large range of values for DOC found at each site. Except for one seepage spring, the range of values was more than 2.5 times the median. Activities of medium to large mammals (e.g. feeding and defecation) is a possible O RGANIC CARBON IN SHALLO W SUBTERRANEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 296 Tab. 1: Mean values ( Standard Error) of dissolved organic carbon from a series of habitats on Nanos Mountain, Slovenia. Each value is based on ve analyses of a single water sample. Hyporheic Niphargus seepage spring Seepage spring Seepage spring Gushet Gushet Stream January 14, 2011 1.23.24 2.67.54 ** 1.91.56 2.74.90 5.48.24 12.66.56 February 27, 2011 ** 1.74 0.60 0.40 7.07 3.23 4.05 March 25, 2011 1.38 0.83 1.11 0.96 3.31 6.53 **** May 5, 2011 1.21 9.89 5.53 1.69 1.58 0.63 **** May 27, 2011 10.44+1.05 9.48+1.73 1.34+2.59 1.43+0.69 1.02+0.93 0.45+0.33 1.30+0.47 June 16, 2011 0.41+0.22 4.66+0.44 0.84+0.48 2.17+0.71 0.13+0.56 0.41+0.56 ** August 26, 2011 1.86.28 2.34.19 * *** *** *** September 30, 2011 * 5.37.28 *** *** *** December 18, 2011 2.06.25 4.22.11 0.75.50 *** *** *** Median 1.38 3.44 1.11 1.56 2.14 2.03 4.05 Range 10.03 9.06 4.93 1.77 6.94 6.12 11.36 Dry ** Sample lost *** Habitat destruction by cattle ****Not sampled explanation for this high variability. e non-signicant dierences in median values showed some interesting patterns. First, typical values for DOC in seepage springs range between 1 and 3.5 mg L (Tab. 1). Second, DOC was higher in the seepage spring with Niphargus amphi pods than in seepage springs without macroscopic stygo bionts such as Niphargus Fig. 5: Medians, maxima, and minima of dissolved organic car bon in mg L for the ve habitats measured on Nanos Moun tain, Slovenia, listed in order of decreasing median DOC: seep age springs with the stygobiotic amphipod genus Niphargus, surface stream, gushets, hyporheic, and seepage springs without Niphargus. D ISCUSSION Fong and Kavanaugh (2010) suggest that the hypotelmi norheic amphipod Stygobromus tenuis potomacus active ly foraged in a seepage spring, and our results support his conjecture. e higher DOC levels in one seepage spring may tend to concentrate Niphargus near the seep age spring. ird, the absence of signicant dierences among sites indicates the supercal nature all of the subterranean habitats have in common. e absence of dierences also suggests that there is not signicant re duction of organic carbon levels by the consumer com munities in the dierent habitats. Overall, dissolved organic carbon at the boundary of the hypotelminorheic (i.e. the seepage spring) had a median value of approximately 1.4 mg C L when no macroscopic stygobionts were present and approximate ly 3.4 mg C L when they were (Tab. 1). Furthermore, these values were not very dierent from that of a nearby Average values are not very dierent from those observed for the hyporheic, an example of which is shown in Tab. 2. In this study, Marmonier et al. (2000) found mean values for DOC ranging from 1.9 to 3.5 mg C L very similar T ANJA PIPAN & D AVID C. C ULVER

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ACTA CARSOLOGICA 42/2-3 2013 297 to the median values found in the Slovenian hypotelmi norheic sites. Other studies, such as that of Datry et al. (2005) found that there was a sharp decline in organic carbon with depth, if measurements were made over a greater depth range. In general, it would appear that or ganic carbon values within 50 cm of the surface are usu ally not very dierent. W hat is dierent between hyporheic and hypotelm inorheic sites is their temporal and spatial variability. DOC values in the study of Marmonier et al. (2000) had small standard errors, while DOC values in the hy potelminorheic varied widely (Fig. 5). For carbon enter ing the hyporheic at sink points, there is a dilution and mixing eect that would act to reduce local dierences. e same is not the case for the hypotelminorheic, where inputs such as animal waste or decaying biomass, would not be diluted or homogenized because there is very lit tle surface water, and the vertical and horizontal distance the water moves into the ground is very small. presents a contrast with the hyporheic and hypotelmi norheic with respect to the pattern of dissolved organic carbon. ere are also more extensive data on DOC in epikarst than for the other two SSHs. All of these stud ies measure water dripping out of epikarst habitats into caves since this is the only point of access to the water. us, DOC is measured at the exit of water from the epikarst. e most extensive study is that of Ban et al. (2008). ey did an extensive series of temporal and spatial mea surements of organic carbon in three drips between 70 m and 100 m apart in Shihua Cave, China. e temporal pattern of dissolved organic carbon (DOC) is shown in Fig. 6. e general shape of the pattern is the same for all three drips, but the amount of DOC also varies from drip to drip. e highest DOC value was observed in drip JG, with a value of 2.76 mg L Mean values ranged from 1.06 mg L at JG to 0.73 mg L at PL, generally lower than values found in the hypotelminorheic (Fig. 5) and hyporheic (Tab. 2). Spikes in DOC, which were some what less extreme than those observed for the hypotelm inorheic, corresponded to precipitation events (e.g. early August 2004). Ban et al. (2008) suggest that more DOC is lost from either longer ow paths or longer times. DOC analyses by Simon, Pipan, and Culver (2007) of Organ Cave, W est Virginia, USA, and Postojna Plani na Cave System in Slovenia allow comparison of DOC levels in epikarst with other components of cave and karst system. As was the case with Shihua Cave (Ban et al. 2008), DOC levels in epikarst waters were typically around 1.0 mg L or even lower, in the case of Postojna Planina Cave System (Tab. 3). In Postojna Planina Cave System, but not in Organ Cave, cave streams had more organic matter than epikarst drips. However, DOC from epikarst water is likely very important in both caves, both because it is more ubiquitous than cave streams, and be cause the organic carbon in epikarst water is more meta bolically accessible than that in cave streams. An exam ple is shown in Fig. 7, where DOC quantity and quality are compared for Postojna Planina Cave System. Simon et al. (2010) measured specic UV absorbance (SUVA) at 254 nm, a standard measure of the frequency of aromatic compounds. Higher SUVA values tend to mean the com pounds are less reactive and less easy to metabolize but there are numerous caveats (W esihaar et al. 2003). SUVA Tab. 2: Dissolved Organic carbon (DOC) component of inter stitial water at two hyporheic sties in the Rhne River basin in France. At the Grand Gravier section of the Rhne River, 12 samples were taken from depths of 20, 50, and 100 cm. Modied from Marmonier et al. (2000). Habitat mg C L + S.E. n Vanoise Brook Surface 1.5 + 0.2 6 40 cm deep in channel 1.9 + 0.1 6 40 cm deep in gravel bar 1.8 + 0.3 6 Grand Gravier 1.5 m in river channel 3.5 + 0.3 12 shore 3.2 + 0.3 12 1.5 m on bank 2.8 + 0.2 12 Fig. 6: Rainfall amount and intraand inter-annual variation of DOC in three drip sites at Shihua Cave during the period be tween April 2003 and December 2006. e important rainfall events causing the increase in DOC concentration in drip water are marked with hollow stars. From Ban et al. (2008). O RGANIC CARBON IN SHALLO W SUBTERRANEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 298 C ONCLUSIONS W ater in the three major shallow subterranean habitats diers in the amount and variability of dissolved organic carbon. e hypotelminorheic, the most supercial of the three and one with close surface connections, has levels of dissolved organic carbon typical of surface waters in the immediate vicinity, between 1 and 4 mg L in the Na nos Mountain study area. ere was very high temporal variability with occasional spikes of 10 mg L or more, reecting the poorly integrated nature of the system. e hyporheic, which is also closely connected with the sur face, had similar values of DOC, but with less variability, reecting the well connected and integrated nature of this habitat. Finally, epikarst showed the lowest average values, with spikes associated with precipitation events. Preliminary analysis of organic carbon quality suggests that epikarst water is more accessible to the subterranean community than cave stream water. values, as well as several oth er measures of reactivity, all suggest that organic carbon in epikarst water is more ac cessible to the subterranean community than cave stream water. Clearly this is a topic worthy of more research. REFERENCES Bakalowicz, M., 2012: Epikarst.In: W hite, W .B. & D.C. Culver (eds.) Encyclopedia of caves, second edition. Elsevier Press, pp. 284, Amsterdam, e Neth erlands. Ban, R., Pan, G., Zhu, J., Cai, B. & M. Tan, 2008: Tem poral and spatial variations in the discharge and dissolved organic carbon of drip waters in Beijing Shihua Cave, China.Hydrological Processes, 22, 3749. Tab. 3: Dissolved organic carbon ( + Standard Error) for Organ Cave, West V irginia, USA, and P ostojna P lanina Cave System, Slovenia. From Simon et al. (2007). Organ Cave Postojna Planina Cave System Habitat DOC (mg L ) n DOC (mg L ) n Sinking Streams 7.67 + 1.03 3 4.36 + 0.46 2 Epikarst Drips 1.10 + 0.15 20 0.70 + 0.04 99 Cave Streams 1.08 + 0.32 6 4.75 + 1.57 3 Resurgence 0.90 + 0.17 3 2.67 + 0.80 2 Fig. 7: Comparison of amounts of DOC (le panel) and specic UV absorbance at 254 nm (right panel) for a series of habitats in P os tojna P lanina Cave System, Slovenia. Modied from Simon et al. (2010). T ANJA PIPAN & D AVID C. C ULVER

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ACTA CARSOLOGICA 42/2-3 2013 299 Culver, D.C. & T. Pipan, 2008: Supercial subterranean habitats gateway to the subterranean realm?.Cave and Karst Science, 35, 5. Culver, D.C. & T. Pipan, 2011: Redening the extent of the aquatic subterranean biotope shallow subter ranean habitats.Ecohydrology, 4, 721. Culver, D.C., Pipan, T. & S. Gottstein, 2006: Hypotelm nean Biology, 4, 1. Culver, D.C., Holsinger, J.R. & D.J. Feller, 2012: e fauna of seepage springs and other shallow subter ranean habitats in the mid-Atlantic Piedmont and Coastal Plain.Northeastern Naturalist Memoirs, 19 (Monograph 9), 1. Culver, D.C. & I. ereg, 2004: Kenks Amphipod (Stygo bromus kenki) and other amphipods in Rock Creek P ark, W ashington, D.C.Report to Rock Creek Park, National Park Service, W ashington, DC. USA. Datry, T., Malard, F. & J. Gibert, 2005: Response of in vertebrate assemblages to increased groundwa ter recharge rates in a phreatic aquifer.Journal of the North American Benthological Society, 24, 461. Danielopol, D.L., Pospisil, P., Dreher, J., Msslacher, F., Torreiter, P., Geiger-Kaiser, M. & A. Gunatilaka, 2000: A groundwater ecosystem in the Danube wet lands at W ien (Austria).In: W ilkens, H., Culver, D.C. & W .F. Humphreys (eds.) Subterranean ecosys tems. Elsevier Press, pp. 481, Amsterdam, e Netherlands. Eamus, D. & R. Froend, 2006: Groundwater-dependent ecosystems.Australian Journal of Botany, 54, 91. Fong, D.W & K.E. Kavanaugh, 2010: Population dy namics of the stygobiotic amphipod crustacean Stygobromus tenuis potomacus and isopod crusta cean Caecidotea kenki at a single hypotelminorheic habitat over a two-year span.In: Mokri, A. & P. Trontelj (eds.) ICSB International Conference on Subterranean Biology, 2010 Abstract Book, 22, Postojna, Slovenia. Gibert, J., 1991: Groundwater systems and their bound aries: conceptual framework and prospects in groundwater ecology.Verhaltlungen der Interna tionalen Vereinigung fr eoretische und Ang ewandte Limnologie, 24, 1605. Gibert, J., Stanford, J.A., Dole-Olivier, M.-J. & J.V. W ard, 1994: Basic attributes of groundwater ecosystems and prospects for research.In: Gibert, J., DanieloIn: Gibert, J., Danielo pol, D.L. & J.A. Stanford (eds.) Groundwater ecol ogy. Academic Press, pp. 7, San Diego, CA. Ginet, R. & V. Decu, 1977: Initiation la biologie a lcologie souterraines.J-P Delarge, Paris. Holsinger, J.R. & G.W Dickson, 1977: Burrowing as a means of survival in the troglobitic amphipod crus tacean Crangonyx antennatus Packard (Crangonyc tidae).Hydrobiologia, 54, 195. Huntsman, B.M., Venarsky, M.D. & J.P. Bernstead, 2011: Relating carrion breakdown rates to ambient re source level and community structure in four cave stream ecosystems.Journal of the North American Benthological Society, 30, 882. Kogovek, J., 2010: Characteristics of percolation through the karst vadose zone. ZRC Publishing, Ljubljana, Slovenia. Krause, S., Hannah, D.M., Fleckenstein, J.H., Heppell, C.M., Kaeser, D., Pickup, R., Pinay, G., Robertson, A.L. & P.J. W ood, 2011: Inter-disciplinary perspec tives on processes in the hyporheic zone.Ecohy drology, 4, 481. Krei, N., 2010: Types and classications of springs.In: Kresic, N. & Z Stevanovic (eds.) Groundwater hy drology of springs. Engineering, theory, management, and sustainability. Elsevier Press, pp. 31, Am sterdam, e Netherlands. Malard, F., Tockner, K., Dole-Oliver, M.-J. & J.V. W ard, 2002: A landscape perspective of surface-subsurface hydrological exchange in river corridors.Freshwa ter Biology, 47, 621. Marmonier, P., Creuz des Chtelliers, M., Dole-Olivier, M.-J., Plnet, S. & J. Gibert, 2000: Rhne ground water systems.In: W ilkens, H., Culver, D.C. & W .F. Humphreys (eds.) Subterranean ecosystems Elsevier Press, pp. 513, Amsterdam, e Netherlands. Metrov, M., 1962: Un nouveau milieu aquatique sou terrain: le biotope hypotelminorheique.Compte Rendus de lAcadmie des Sciences, Paris, 254, 2677. Metrov, M., 1964: Dirences et relations faunistiques et cologiques entre les milieu souterrains aqua tiques.Spelunca Mmoires, 4, 185. Palmer, A.N., 2007: Cave geology. Cave Books, Dayton, Ohio. Pipan, T., 2005: Epikarst a promising habitat .Zaloba ZRC, Ljubljana, Slovenia. Pipan, T. & D.C. Culver, 2007: Regional species rich epikarst copepods.Journal of Biogeography, 34, 854. Pipan, T., Lpez, H., Orom, P., Polak, S. & D.C. Cul ver, 2011: Temperature variation and the presence of troglobionts in terrestrial shallow subterranean habitats.Journal of Natural History, 45, 253. O RGANIC CARBON IN SHALLO W SUBTERRANEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 300 Pipan, T. & D.C. Culver, 2012: Convergence and diver gence in the subterranean realm: a reassessment.Biological Journal of the Linnean Society, 107, 1. Simon, K.S. & E.F. Beneld, 2002: Ammonium retention and whole-stream metabolism in cave streams.Hydrobiologia, 482, 31. Simon, K.S., Pipan, T. & D.C. Culver, 2007: A concep tual model of the ow and distribution of organic carbon in caves.Journal of Cave and Karst Studies, 69, 279. Simon, K.S., Pipan, T., Ohno, T. & D.C. Culver, 2010: Spatial and temporal patterns in abundance and character of dissolved organic matter in two karst aquifers.Fundamental and Applied Limnology, 177, 81. Venarsky, M.D., Bernstead, J.P. & A.D. Huryn, 2012: Eects of organic matter and season on leaf litter colonization and breakdown in cave streams.Fresh W ater Biology, 57, 773. Vervier, P., Gibert, J., Marmonier, P. & M.-J. Dole-Olivi er, 1992: A perspective on the permeability of the surface freshwater-groundwater ecotone.Journal of the North American Benthological Society, 11, 93. W eishaar, J. L., Aiken, G. R., Bergamaschi, B. A., Fram, M. S., Fujii, R. & K. Mopper, 2003: Evaluation of specic ultraviolet absorbance as an indicator of the chemical composition and reactivity of dissolved organic carbon.Environmental Science and Tech nology, 37, 4702. W illiams, P.W ., 2008: e role of the epikarst in karst and cave hydrogeology: a review.International Journal of Speleology, 37, 1. T ANJA PIPAN & D AVID C. C ULVER



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V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC CONDITIONS S PREMENLJIVOST PRETAKANJA VODA IN PRENOSA SNOVI V KRASU OB RAZLI NIH HIDROLO KIH POGOJIH Nataa R AVBAR 1 Izvleek UDK 911.2:551.444 556.33:551.44 Nataa Ravbar: Spremenljivost pretakanja voda in prenosa snovi v krasu ob razlinih hidrolokih pogojih Predstavljeni so pomen hidroloke spremenljivosti v krasu, dejavniki, ki sproajo tovrstno spremeljivost in posledice. Za pretakanje podzemnih krakih voda je znailna velika spre menljivost v odvisnosti od razlinih hidrolokih razmer. Nihanja podzemne vode se lahko spreminjajo za ve deset metrov, razlike v hitrostih pretakanja voda ob nizkem ali vi sokem vodostaju so lahko deset ali vekratne. Glede na tre nutne hidroloke razmere pogosto prihaja do spreminjanja smeri podzemskega toka, kar lahko povzroi razlino prispe vnost doloenih delov vodonosnika k individualnemu izviru. Opisana hidroloka spremenljivost pa lahko izrazito vpliva na transport onesnaeval, na razpololjivost podzemne vode ter na njeno obutljivost na onesnaenje. Dvig gladine podzemne vode povzroa zmanjanje debeline nezasiene cone in s tem nijo samoistilno sposobnost prenikajoih vod. Vije hitrosti pretakanja voda vplivajo na kraje zadrevalne ase v podzem lju, podzemni tok je bolj turbulenten in zato transport in mobi lizacija topnih in netopnih snovi bolj efektivna. Zato je poseb no na krakih obmojih, za katere je znailna velika hidroloka spremeljivost, to lastnost potrebno upotevati pri prouevanju, razumevanju ali napovedovanju hidrolokega obnaanja vodonosnika, ali pri pripravi ustreznih postopkov varovanja. Kljune besede: kraki vodonosnik, asovna hidroloka spre menljivost, transportni procesi, kraki izvir, zaita vodnega vira. 1 Karst Research Institute ZRC SAZU, Postojna, Slovenia, e-mail: natasa.ravbar@zrc-sazu.si Received/Prejeto: 1.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 327, POSTOJNA 2013 Abstract UDC 911.2:551.444 556.33:551.44 Nataa Ravbar: Variability of groundwater ow and trans port processes in karst under dierent hydrologic conditions Signicance of hydrological variability in karst is presented, which also discusses factors inducing such variability and con sequences it may cause. Groundwater ow in karst aquifers is oen characterized by strong variability of ow dynamics in response to dierent hydrologic conditions within a short time period. Consequently, water table uctuations are oen in the order of tens of meters, dierences in ow velocities between lowand high-ow conditions can reach ten or even more times. In dependence to respective hydrologic conditions groundwater ow also results in variations of ow directions, and thus in contribution of dierent parts of the aquifer to a particular spring. e described hydrological variability has many implications for contaminant transport, groundwater availability and vulnerability. Groundwater level rising reduces thickness of the unsaturated zone and decreases protective ness of the overlying layers. Higher water ow velocities reduce underground retention. Due to more turbulent ow, transport and remobilization of solute and insoluble matter is more eec tive. During high-ow conditions there is usually more surface ow and hence more concentrated inltration underground. Particularly in karst systems that show very high hydrologic variability, this should be considered to correctly characterize, understand or predict the aquifers hydrological behaviour and to prepare proper protection strategies. Keywords: karst aquifer, temporal hydrological variabil ity, groundwater ow, transport processes, karst spring, water source protection.

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ACTA CARSOLOGICA 42/2-3 2013 328 Since recognition of the signicance of karst aquifers as important water resource and valuable ecosystems is growing worldwide, these hydrological systems are re ceiving rapidly increasing attention from the scientic, engineering and regulatory communities. Due to the many challenges related to their characterization and management, such aquifers require good knowledge and comprehension of groundwater ow characteristics (Bakalowicz 2005; Goldscheider & Drew 2007; Bonacci et al. 2009; Kresic & Stevanovic 2010). Karst hydrogeologic systems are associated with a high level of heterogeneity encountered in karstied rocks that is manifested in a duality of fundamental hy draulic processes occurring in the aquifer (Kirly et al. 1995; W orthington 1991). e hydrogeological hetero geneity concern aquifer recharge (diuse/concentrat ed), storage and porosity (pores and micro-fractures/ conduits), and discharge (diuse/concentrated). Unlike all other aquifer types, ow behavior in karst aquifers exhibits rapid or immediate inltration, close interrela tionship between surface and groundwater, high under ground ow velocities (reaching up to several hundred meters per hour) and high conductivity in the prevailing conduit systems. ere are frequently connections and intersections of water paths over large distances (up to many tens of kilometers). Due to the specic nature of karst aquifers, the system response to recharge processes is also controlled by the manner in which inltrating water is transmitted through the aquifer. Consequent ly, karst hydrological systems may be very sensitive to changes in the hydrometeorological conditions of re charge. Hydraulic reactions to dierent hydrologic con ditions may result in dierences of ow which is relevant in many respects (W hite 1988; Bakalowicz 2005; Ford & W illiams 2007). Accordingly, behavior of karst aquifers is oen un predictable and their water resources are extremely vul nerable to contamination, over-exploitation and climate change (Drew & Htzl 1999). Ignoring these described characteristics, when carrying out hydrogeological in vestigations in carbonate aquifers, when confronting specic environmental and engineering problems or when planning management of groundwater resources, is potentially erroneous. Moreover, investigating and planning in karst requires special adaptations of investi gation techniques, avoiding generalizations and/or inter polations (Goldscheider & Drew 2007; Milanovi 2006; Sass & Burbaum 2010). Holistically taking into account the karstic nature of carbonate aquifers therefore rep resents a key step toward appropriate investigation and planning in karst. Many excellent publications prove that the unique ness of water ow in karst is already very well known and appreciated, at least among those academicians, practicing hydrogeologists, and water resources profes sionals who regularly deal with karst and karst related problems. General specics of water ow in karst are also adequately abided in various studies, expertise, when choosing research methods, solving problems, etc. However, temporal hydrological variability is considered in a minor degree, although some karst systems (e.g. Di naric karst, karst under sub-tropical/monsoon climates) exhibit considerable hydrologic variability. ere is also a lack of readily available research systematically study ing the role of temporal variability in karst. erefore the aim of this contribution is to list the reasons for temporal hydrologic variability, and to dis cuss its relevance to questions of karst aquifer behavior, as well as groundwater availability and protection. Some case studies that illustrate variability of groundwater ow and transport processes in karst under dierent hydro logic conditions are presented, and stress the importance of considering hydrologic variability. I NTRODUCTION K ARST A QUIFER HETEROGENEITY AND TEMPORAL HYDROLOGICAL VARIABILITY From a hydrogeological perspective, the most distinctive property of karst aquifers that dierentiate them from other hydrogeological systems is the high solubility of the rock medium determining the high heterogeneity of hy draulic aquifer properties (W hite 1988; Ford & W illiams 2007; Klimchouk & Ford 2000; Kirly 2002). e duality and heterogeneity of karst is reected in a fast water com ponent with a low storage capacity in conduits (preferen tial ow) on one hand, and the slow water ow and high storage capacity component in the ssured-porous karst system (diuse ow) on the other hand. Solutional enlarging of ssures being the unique hydrogeological characteristic of karst rocks makes them highly permeable and enables immediate inltration of N ATA A R AVBAR

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ACTA CARSOLOGICA 42/2-3 2013 329 Fig. 1: Conceptual three-dimensional model of karst aquifer and groundwater ow, illustrating the heterogeneity of karst: allogenic and autogenic type of recharge, point and diuse inltration, ow through conduits and low permeability matrix (modied aer Ravbar 2007). water into the subsurface. In the underground it creates cavities organized in a ow net in a hierarchical manner (Bakalowicz et al. 1994; Gabrovek 2000). e under ground drainage system is then integrated into ecient, mainly sub-horizontally oriented conduits for the collec tion, transport and ultimately discharge of recharge wa ters (Fig. 1). Although the karst conduit system occupies only a small portion of the total aquifer porosity, it may have a major impact on the specic hydraulic behavior of the karst system. Among the special properties of water ow in some karst aquifers, fast and strong hydrologic variations in response to precipitation events or snowmelt are oen shown. ese strongly depend on hydrometeorological and hydrogeological factors. e rst group of factors includes the type, amount, intensity and distribution of precipitation, and factors governing snowmelt, such as temperature and wind. e second group comprises catchment size, aquifer geometry, eective porosity, the dimensions and connectivity of the karst conduits and the antecedent soil moisture. ese variations may include: a) signicant groundwater level uctuations; b) changing ow directions; c) changing ow velocities; d) change from open-channel to pressurized ow; e) change from laminar to turbulent ow. Consequently, shiing groundwater divides and dierent types of surface-groundwater interaction, such as activating temporarily active streams, swallow holes, etc. can be observed. Likewise, the hydrologic state of estavelles changes. Spring water physical and chemical properties may be subject to signicant variability as well. ese variations are relevant with respect to con taminant transport and groundwater availability and vulnerability V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC ...

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ACTA CARSOLOGICA 42/2-3 2013 330 WATER TABLE OSCILLATIONS AND DIFFERENCES IN GROUND W ATER FLO W Groundwater level variations in karst aquifers are oen on the order of tens of meters within a short time period. ey may result in variations of ow velocities and di rections, in divergent ow and consequently in shiing catchment boundaries, and surface-groundwater inter actions. Strong water table oscillations are particularly pronounced in the Dinaric karst that stretches along the Adriatic coast, characterized by vast karst plateaus in tersected by karst poljes. In the famous cave system of kocjanske Jame, water levels can rise up to about 100 m (Habe 1966). In Popovo Polje in Bosnia and Herzegovi na, the largest karst polje in the Dinaric karst, the highest variations are recorded at more than 200 m (Milanovi 2006). Particularly at the intermediate karst poljes, shallow karst areas or in the contact karst areas, various forms of interaction between groundwater and surface water can be observed as well. Examples from the Slovene Classical karst landscapes that show important hydrologic varia tions are presented below. At Cerkniko Polje and Pivka valley, which are located on the boundary of AdriaticBlack Sea watershed, ow bifurcations can be observed (Fig. 2): During low-water conditions, groundwater from the Javorniki Mountains and Pivka Valley drain toward Planinsko Polje to the northeast. In high-water condi tions, water levels rise and a groundwater divide forms below the Javorniki Mountains so that a part of the area drains towards Pivka Valley to the southwest. Furthermore, due to groundwater uctuations and weak connections between dierent karst conduits, sev eral intermittent lakes of dierent size, temporal dura tion and frequency occur in this region. e largest one is the lake of Cerkniko Jezero, which can extend over 26 km 2 and contain more than 82 million m 3 of water. C ASE STUDIES Fig. 2: P hysical map of the Slovene Classical karst with the underground connections proven by tracer tests, a schematic cross-section of the area showing ow bifurcation, and the conceptual model of the aquifer system functioning during lowand high-water condi tions with wider arrows indicating proportionately great ow volume (modied aer Gospodari & H abi 1976; Kogovek et al. 1999; Gabrovek et al. 2010; P etri 2010; Ravbar et al. 2012). N ATA A R AVBAR

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ACTA CARSOLOGICA 42/2-3 2013 331 e lake usually lls twice per year via springs and es tavelles, but remains empty in dry years; in wet years, it occurs several times per year and/or does not dry up en tirely (Kovai 2010). Malenica Spring is the principal spring of the area and is a regionally important drinking water supply that outows into Planinsko Polje. Nearby Unica Spring out ows from the well known Planinska Jama, where two subsurface river channels (Rak and Pivka branches) converge. In recent years, several tracer tests have been done in dierent hydrological conditions (Gabrovek et al. 2010; Petri 2010; Ravbar et al. 2012). e results revealed that the relations between various contribution areas to the springs are strongly dependant on temporal hydrologic conditions (Fig. 2). In general, Malenica Spring is recharged mainly from the Cerknica direction and Javorniki karst aquifer, and there is no direct connection with the ponor of the Pivka River sinking in Postojnska Jama. At high water levels, inows from the Cerknica part dominate. e outow from Malenica Spring is thus bounded and the Rak branch in Planinska Jama acts as an overow. e Rak branch is recharged from both the Cerknica and Javorniki parts, and the Pivka branch from Javorniki and the Pivka parts. At low waters, aer the emptying of the intermittent Cerkniko Jezero, the proportion of inow from the Javorniki part to the Malenica Spring is more important. e Rak branch progressively drains and Unica Spring is principally fed by the Pivka branch. Studies made at higher groundwater level showed that the apparent dominant ow velocities were for two to four times higher than in conditions of constant water level recession. In well developed conduit networks (e.g. cave system of Postojnska Jama) ow velocities during medium water levels were at least seventeen times higher in comparison to the velocities observed at low water levels. Several other studies worldwide also showed con siderable variations of groundwater ow velocities and directions as a function of respective hydrologic condi tions (e.g., Kogovek & Liu 2000; Gppert & Goldsc heider 2008; Pronk et al. 2007). V ARIABILITY OF TRANSPORT PROCESSES e extent to which ssures or conduits are lled by a re charge pulse will determine the nature of the water ow and transport processes, and there may be variable lags between the input pulse and the response at the spring. Changes from laminar to turbulent ow may occur, resulting in higher transport velocities, shorter transit times, more eective transport of sediments and bacte ria, and mobilization of DNAPLs (Dense Non-Aqueous Phase Liquids). Raising water table above the conduit ceiling induces changes from open-channel to pressu rized ow. Storm water, together with contaminants, thus penetrate and are temporarily stored in the adja cent less karstied zones of the aquifer and the overlying unsaturated zone. Vadose conduits then become tempo rarily phreatic and change the nature of the karst drain age system. Groundwater ow velocities decrease in the transition from open channel to pressurized ow due to the increased head loss of pressurized ow compared to open channel ow. Later, when the recharge decreases, the ooded zones are drained back again into the conduits and the temporarily stored water leaves the system slowly (Fig. 3). Solute matter may therefore persistent in the outow for Fig. 3: H ydraulic interaction between karst conduits and the surrounding aquifer during dierent hydrologic conditions. V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC ...

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ACTA CARSOLOGICA 42/2-3 2013 332 longer periods, whereas particles may be attenuated in the aquifer (Cornaton & Perrochet 2002). In a study by Raeisi et al. (2007), changes in hy draulic ow of a conduit within the Mammoth Cave System were compared between partially-full pipe and full pipe conditions of the Logsdon River. Analysis of temperature, electrical conductivity, water level and velocity demonstrated relationships to the geometry to conduit and uid transport behavior. W hen Logsdon River exceeds the capacity of its conduit, open channel ow changes into pressurized ow. Meanwhile, initial minimums in electrical conductivity represented the early movement of storm water through the conduit, while the subsequent second minimum was interpreted as storm water temporarily stored in adjacent areas that drained back into the conduit when the water level was decreasing. Changes in electrical conductivity during partially-full pipe conditions were mainly controlled by external recharge conditions, such as the behavior of sinking streams. Similarly Goldscheider (2005) observed dierent breakthrough curves when tracing in the Hochifen-Got tesacker area in the Austro-German Alps during lowand variable high-ow conditions. During the experi ment under low-ow conditions, the outow of tracer was concentrated, continuous with a uniformly shaped breakthrough curve inferring to tracer transport that was restricted to a single ssure or a series of parallel ssures. On the contrary, due to the sudden increase of hydraulic pressure in the conduits and the rise of the water table during the experiment under variable ow conditions, the tracer breakthrough curves began with extremely steep increases in tracer concentration. e rst appear ance of the tracer coincided with its maximum concen tration and aer the steep decrease in the concentrations the outow of the tracer persisted for several weeks. Such behavior indicates quick transfer of the tracer through preferential ow paths, followed by slow depletion from the low permeability volumes. V ARIATIONS OF SPRING W ATER PROPERTIES Karst springs represent the recharge from their entire aquifer systems. ey reect water ow characteristics within the systems and may reveal possible contamina tion in their catchments. Owing to fast and strong reac tions of karst systems to variable recharge conditions, karst springs generally result in sudden variations of discharge, physical, chemical, isotopic and microbiologi cal water composition (Ryan & Meiman 1996; Katz et al. 1998; Auckenthaler et al. 2002; Vesper & W hite 2003). ere is a lot of literature on the variability of spring wa ter characteristics, particularly the interpretations of dif ferent time series variables (hydrographs, chemographs, turbidigraphs, etc.) on both seasonal time scales and in dividual storm events. Generally, discharge rates may vary by many orders of magnitude and may include abrupt changes in water quality. is is a key problem for water suppliers, because otherwise safe water sources may suddenly be charged with high levels of contaminants, such as DNAPLs (Loop & W hite 2001), toxic metals (Vesper & W hite 2003) or fecal and pathogenic bacteria (Pronk et al. 2007). An example of the variable properties of spring wa ter is Hannett Spring, an important karst water source in Normandy, France (Massei et al. 2003; Fournier et al. 2007). e spring drains a chalk binary karst system. It is oen characterized by high turbidity and phosphate con centrations, which poses potential health problems be cause of the great ability of bacteria to sorb onto particu lates. erefore the analyses of turbidity dynamics was used to characterize the direct transfer and re-suspension of components induced by the change of recharge condi tions. e results reveal that aer a recharge increase, a primary turbidity peak that coincides with a decrease in electrical conductivity and increase of phosphates indi cates direct transfer of surface water inltrated at a swal low hole. A secondary turbidity peak coincides with an increase in discharge and nitrate concentrations which indicate groundwater level rising and diuse inltration. is peak corresponds to re-suspension of intraclastic sediments induced by pressure transfer, which allows the increase of velocity and change from laminar to turbu lent ow inside the karst conduits. is information can be used to optimize water treatment or disconnect the spring from the distribution network. However, water quality variability includes much valuable hydrogeologic information. An example is the anomalous behavior of electrical conductivity observed at a typical karst spring in Slovenia (Ravbar et al. 2011). e spring is mainly recharged by diuse inltration. As a function of the hydrologic conditions, the groundwa ter table in its catchment uctuates several tens of meters which can be observed in a cave near the spring and in the two intermittent lakes in its catchment. Aer typical response of the spring during several high-ow events, electrical conductivity rises again and remains elevated during the entire high-ow period, typically 2040 S/cm above the baseow value. Based on the tracer tests results and other considerations (water balance, topography, geologic structure) this behavior is explained by variable catchment boundaries and variations of ow directions (Fig. 4); when the water level in the aquifer rises, the catchment expands, incorporating zones of groundwater with higher electrical conductivity, caused by higher un saturated zone thickness and subtle lithologic changes. N ATA A R AVBAR

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ACTA CARSOLOGICA 42/2-3 2013 333 Fig. 4: A) Anomalous behavior of electrical conductivity observed at the P odstenjek spring following short but intense rainfall. B) Con ceptual model for the anomalous electrical conductivity variability induced by variable catchment boundaries and variations of ow directions (modied aer Ravbar et al. 2011). V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC ...

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ACTA CARSOLOGICA 42/2-3 2013 334 In karst landscapes, rivers and streams only occasion ally ow on the surface. Subterranean drainage through mostly unknown and unpredictable ow paths prevail. In high karst plateaus and mountainous karst regions, the water table is oen very deep below the surface, some times several hundreds of meters deep and inaccessible or too expensive for drilling water wells (Audra et al. 2003; Zhu et al. 2007; Plan et al 2009). erefore, in many situ ations, karst groundwater can only be observed in karst springs and some caves. Caves that are large enough for humans to enter oer the unique opportunity for direct observation of water seepage through the unsaturated zone (Emblanch et al. 2003; W illiams 2008; Kogovek 2010) or consoli dated water ow as cave streams in karst conduits and channels (Perrin et al. 2003; Groves & Meiman 2005; Ravbar et al. 2012). Karst springs represent natural out ows of groundwater from the aquifer. Due to poor sur face access, they represent ideal spots for the insight into the underground. However, they are typically sparsely distributed, but may be very large, achieving discharge values of some tens of m 3 /s. Nevertheless, karst springs are regarded as important monitoring points and are commonly used as monitoring locations and to collect data of karst hydrogeological regimes (Shuster & W hite 1971; Bonacci 2001; Bakalowicz 2005). Due to the inherent characteristics of karst de scribed above, conventional hydrologic and hydrogeo logic investigation methods are oen of limited value or fail when applied to karst. For example, hydraulic mea surements in boreholes provide useful information about subsurface aquifers, but the heterogeneity of karst aqui fers poses a real challenge in conventional well logging since it is oen dicult to drill a successful borehole in tersecting preferential ow paths within the surrounding rock. Boreholes are also oen not really representative of the organization and functioning of the karst aquifer (Drogue 1980; Jeannin & Sauter 1998; W orthington & Ford 2009). Characterizing and quantifying the eects of temporal hydrologic variability thus requires special adaptations of the classical methodological approaches of sampling and monitoring strategies, and karst-specic methods as already emphasized by Bakalowicz (2005) and Goldscheider and Drew (2007). Analyses of groundwater physiochemical param eters combined with karst spring hydrographs are oen used to obtain information on the behavior and struc ture of karst aquifers, on dierent recharge mechanisms and on contamination problems. Event-based, quarterly or semi-annual sampling of groundwater physical and chemical properties and the combined interpretation of karst spring hydrographs and chemographs is the most widespread technique used for indicating aquifer char acteristics (Shuster & W hite 1971). Recently, many stud ies were based on monitoring and analyzing parameters such as detailed water chemistry (e.g., Ryan & Meiman 1996; Grasso et al. 2003; Birk et al. 2004; Liu et al. 2004; Mudarra & Andreo 2010), dissolved or total organic car bon (DOC/TOC) and turbidity (e.g., Emblanch et al. 1998; Batiot et al. 2003; Massei et al. 2006; Pronk et al. 2006), as well as isotopes (e.g., Lee & Krothe 2003; Gai non et al. 2007). Due to abrupt changes in groundwater level, high and rapid variations of discharge, chemical and microbial constituents, continuous high-resolution records of karst water quality and aquifer behavior have proven very use I NVESTIGATION TECHNI QUES Fig. 5: A) Cave stream monitoring of natural parameters (water level, temperature and electrical conductivity), P ivka branch of P lanin ska J ama, Slovenia. B) Automatic sampler at Malenica Spring, Slovenia (P hotos: N. Ravbar). N ATA A R AVBAR

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ACTA CARSOLOGICA 42/2-3 2013 335 C ONCLUSION e issues and examples presented here on the variabil ity of groundwater ow and transport processes in karst, under dierent hydrologic conditions, illustrate the im portance of temporal hydrologic variations in some karst aquifers on their behavior that may consequently have many implications particularly for changes in groundwa ter availability, contaminant transport and groundwater vulnerability. Several tens of meters large water level uctuations may result in a considerably variable thickness of the unsaturated zone, dierent surface-groundwater interac tions and in divergent ow. is may cause decreasing unsaturated zone thickness, decreasing protectiveness of the overlying layers, and increasing vulnerability. Tem poral variations may induce preferential ow to occur, which permit contaminants to bypass overlying layers that may otherwise attenuate them. Variations of ow di rections can result in contributions from dierent parts of the aquifer to a particular spring, and thus in variable catchment boundaries which are crucial for source pro tection. Variations in groundwater ow velocities and variability of transport processes may drastically change the properties of water at the observation point and in uence the derived results and interpretation. Moreover, higher water ow velocities reduce underground reten tion. Some contaminants may be largely immobile dur ing low-water conditions, as they might be stored in the unsaturated zone, in cave sediments or at the bottom of water-lled cavities. During high-ow events, these con taminants may be remobilized and cause sanitary prob lems. erefore, in areas of considerable hydrological variability, this specic need to be taken into account when executing dierent aquifer studies, modeling and interpretations, and when preparing various utilization and management strategies. An example of how tempo ral variability can be considered in water resource pro tection and management practices is its integration into vulnerability assessment (Ravbar & Goldscheider 2007, 2009). us, a thorough understanding of water and con taminant transfer through the soil, epikarst, unsaturated zone and active conduit network toward springs or other drinking water withdrawal points are required in order to assure appropriate investigation and planning in karst or to avoid eventual environmental and socio-economic consequences of groundwater contamination. ful for the interpretation of temporal variability and of the functioning of karst aquifer systems (Fig. 5). Due to rapid aquifer responses, signicantly shorter sampling intervals during high-ow events need to be employed. e study of Liu et al. (2007) stressed the need for con tinuous hydrogeochemical monitoring for at least one Fig. 6: Injection of 50 g of uranine into a stream sinking at Qiaotou P onor in central Y unnan P rovince, China (P hoto: N. Ravbar). hydrological year, including detailed seasonal, diurnal and storm-scale patterns in order to reveal highly accu rate dynamics and variability of the systems behavior. In addition, tracer techniques using articial trac ers can be complementary to dening contribution/ catchment areas, investigating underground water ow characteristics and transport of matter in dierent hy drological conditions (Kss 1998; Benischke et al. 2007). Detailed breakthrough curves (BTC) and tracer recovery provide data on hydraulic connections, dispersivity, and underground ow rates in the system. e advantage of the tracer method is generally controlled injection con ditions (selection of injection points, injection mode; Fig. 6), however, cost eective articial tracing can most oen only be applied in limited areas or over a small sur face (Andreo et al. 2006). e combined application of articial and environ mental tracers can be used to assemble the information of temporal variability of karst aquifers. However, to date only a few studies combined the study of natural and articial tracers (e.g., Kogovek 2001; Auckenthaler et al. 2002; Einsiedl 2005; Savoy 2007; Pronk et al. 2008; Ravbar et al. 2012). V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC ...

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ACTA CARSOLOGICA 42/2-3 2013 336 Andreo, B., Goldscheider, N., Vadillo, I., Vas, J.M., Neu kum, C., Sinreich, M., Jimnez, P., Brechenmacher, J., Carrasco, F., Htzl, H., Perles, J.M. & F. Zwahlen, 2006: Karst groundwater protection: First applica tion of a Pan-European Approach to vulnerabil ity, hazard and risk mapping in the Sierra de Lbar (Southern Spain).Science of the Total Environ ment, 357, 54. Auckenthaler, A., Raso, G. & P. Huggenberger, 2002: Particle transport in a karst aquifer: natural and articial tracer experiments with bacteria, bacterio phages and microspheres.W ater Sci Technol 46, 3, 131. Audra, P., Q uinif, Y. & P. Rochette, 2003: e genesis of the Tennengebirge karst and caves (Salzburg, Austria).Journal of Cave and Karst Studies 64, 3, 153. Bakalowicz, M., 2005: Karst groundwater: a challenge for new resources.Hydrogeology Journal, 13,1, 148. Bakalowicz, M., Crochet, P., DHulst, D., Mangin, A., Marsaud, B., Ricard, J. & R. Rouch, 1994: High dis charge pumping in a vertical cave: fundamental and applied results.In: Crampon, N. & M. Bakalowicz (eds.). Basic and applied hydrogeological research in French karstic areas. COST Action 65. pp. 93, Montepellier, Millau. Batiot, C., Emblanch, C. & B. Blavoux, 2003: Carbone organique total (COT) et magnsium (Mg2+):deux traceurs complmentaires du temps de sjour dans laquifre karstique.Geoscience, 335, 205. Benischke, R., Goldscheider, N. & C. Smart, 2007: Tracer techniques.In: Goldscheider, N. & D. Drew (eds.) Methods in Karst H ydrogeology. Taylor & Francis, pp. 147, London. Birk, S., Liedl, R. & M. Sauter, 2004: Identication of localised recharge and conduit ow by combined analysis of hydraulic and physico-chemical spring responses (Urenbrunnen, SW -Germany).Journal of Hydrology, 286, 1, 179. Bonacci, O., 2001: Analysis of the maximum discharge of karst springs.Hydrogeology Journal, 9, 328. Bonacci, O., Pipan, T., Culver, D.C., 2009: A framework for karst ecohydrology.Environ. Geol., 56, 891 900. Cornaton, F. & P. Perrochet, 2002: Analytical 1-D dualporosity equivalent solutions to 3-D discrete singlecontinuum models. Application to karstic spring hydrograph modeling.Journal of Hydrology, 262, 165. Drew D. & H. Htzl (eds.), 1999: Karst H ydrology and H uman Activities.International Contributions to Hydrogeology. A. A. Balkema, International Asso ciation of Hydrologists, pp. 322, Rotterdam. Drogue, C., 1980: Essai didentication dun type de structure de magasine carbonates, ssurs [An at tempt at identication of types of water storage in ssured carbonate rocks].Mem Hydrogeol Ser Soc Geol Fr, 2, 101. Einsiedl, F., 2005: Flow system dynamics and water stor age of a ssured-porous karst aquifer characterized by articial and environmental tracers.Journal of Hydrology, 312, 312. Emblanch, C., Blavoux, B., Puig, J.M. & J. Mudry, 1998: Dissolved organic carbon of inltration within the autogenic karst hydrosystem.Geophysical Re search Letters, 25, 9, 1459. Emblanch, C., Zuppi, G.M., Mudry, J., Blavoux, B. & C. Batiot, 2003: Carbon 13 of TDIC to quantify the role of the unsaturated zone: the example of the Vau cluse karst systems (Southeastern France).Journal of Hydrology, 279, 1, 262. Ford, D. & P. W illiams, 2007: Karst hydrogeology and geo morphology.Academic Division of Unwin Hyman Ltd, pp. 601, London. Fournier, M., Massei, N., Bakalowicz, M., DussartBaptista, L., Rodet, J. & J. P. Dupont, 2007: Using turbidity dynamics and geochemical variability as a tool for understanding the behavior and vulner ability of a karst aquifer.Hydrogeology Journal, 15, 4, 689. Gabrovek, F., 2000 (Ed.): Evolution of karst: from preka rst to cessation. ZRC Publishing, ZRC SAZU, pp. 150, Ljubljana. Gabrovek, F., Kogovek, J., Kovai, G., Petri, M., Ravbar, N. & J. Turk, 2010: Recent results of tracer tests in the catchment of the Unica River (SW Slo venia).Acta Carsologica 39, 1, 27. Gainon, F., Goldscheider, N. & H. Surbeck, 2007: Con ceptual model for the origin of high radon levels in spring waters the example of the St. Placidus spring, Grisons, Swiss Alps.Swiss Journal of Geo sciences, 100, 2, 251. Goldscheider, N., 2005: Fold structure and under ground drainage pattern in the alpine karst system Hochifen-Gottesacker.Ecologae Geologicae Hel vetiae, 98, 1, 1. Goldscheider, N. & D. Drew, 2007: Methods in Karst H y drogeology. Taylor & Francis, pp. 264, London. REFERENCES N ATA A R AVBAR

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ACTA CARSOLOGICA 42/2-3 2013 337 Gppert, N. & N. Goldscheider, 2008: Solute and colloid transport in karst conduits under lowand highow conditions.Ground W ater, 46, 1, 61. Gospodari, R. & P. Habi, (ed.), 1976: Underground wa ter tracing: Investigations in Slovenia 1972.ird International Symposium of Underground W ater Tracing (3. SUW T), pp. 312, Ljubljana, Bled. Grasso, D.A., Jeannin, P.Y. & F. Zwahlen, 2003: A deter ministic approach to the coupled analysis of karst springs hydrographs and chemographs.Journal of Hydrology, 271, 65. Groves, C. & J. Meiman, 2005: W eathering, geomorphic work, and karst landscape evolution in the Cave City groundwater basin, Mammoth Cave, Ken tucky.Geomorphology, 67, 1, 115. Habe, F., 1966. Katastrofalne poplave pred naimi turistinimi jamami.Nae jame, 8, 45. Jeannin, P-Y. & M. Sauter, 1998: Analysis of karst hy drodynamic behaviour using global approach: a re view.Bulletin dHydrogeologie, 16, 31. Kss, W ., 1998: Tracing technique in Geohydrology.A. A. Balkema, pp. 581, Rotterdam, Brookeld. Katz, B.G., Catches, J.S., Bullen, T.D. & R.L. Michel, 1998: Changes in the isotopic and chemical com position of ground water resulting from a recharge pulse from a sinking stream.Journal of Hydrology, 211, 1, 178. Kirly, L., 2002: Karstication and groundwater ow. In: Gabrovek, F. (ed.) Evolution of karst: from prekarst to cessation. ZRC Publishing, ZRC SAZU, pp. 155 190, Ljubljana. Kirly, L., Perrochet, P. & Y. Rossier, 1995: Eect of the epikarst on the hydrograph of karst springs: a nu merical approach.Bulletin d'Hydrogeologie, 14, 199. Klimchouk, A.B. & D.C. Ford, 2000: Types of karst and evolution of hydrogeologic settings.In: Klim chouk, A.B., Ford, D.C., Palmer, A.N. & W Drey brodt (eds.) Speleogenesis: Evolution of karst aqui fers. National Speleological Society, pp. 45, Huntsville, Alabama. Kogovek, J. & H. Liu, 2000: W ater tracing test in the Tianshengan region, Yunnan China at high water level.Acta Carsologica, 29, 2, 249. Kogovek, J., 2001: Monitoring the Malenica water pulse by several parameters in November 1997.Acta Carsologica, 30, 1, 39. Kogovek, J., 2010: Characteristics of percolation through the karst vadose zone.ZRC Publishing, ZRC SAZU, pp.168, Ljubljana. Kogovek, J., Knez, M., Mihevc, A., Petri, M., Slabe, T. & S. ebela, 1999: Military training area in Kras (Slo venia).Environmental Geology, 38, 1, 69. Kovai, G., 2010: An attempt towards an assessment of the Cerknica Polje water balance. ).Acta Carso logica 39, 1, 39. Kresic, N. & Z. Stevanovic (eds.), 2010: Groundwater hy drology of springs: engineering, theory, management, and sustainability.Butterworth-Heinemann, pp. 573, Burlington. Lee, E.S. & N.C. Krothe, 2003: Delineating the karstic ow system in the upper Lost River drainage basin, south central Indiana: using sulphate and t34SSO4 as tracers.Appl Geochem, 18, 145. Liu, Z.H., Groves, C., Yuan, D.X., Meiman, J., Jiang, G.H., He, S.Y. & Q .A. Li, 2004: Hydrochemical variations during ood pulses in the south-west China peak cluster karst: impacts of CaCO 3 -H 2 O-CO 2 interac tions.Hydrological Processes, 18, 13, 2423. Liu, Z.H., Li, Q ., Sun, H.L. & J.L. W ang, 2007: Seasonal, diurnal and storm-scale hydrochemical variations of typical epikarst springs in subtropical karst areas of SW China: Soil CO2 and dilution eects.Jour nal of Hydrology, 337, 207. Loop, C.M. & W .B. W hite, 2001: A conceptual model for DNAPL transport in karst ground water basins.Ground W ater, 39, 1, 119. Massei, N., W ang, H.Q ., Dupont, J.P., Rodet, J. & B. Laignel, 2003: Assessment of direct transfer and resuspension of particles during turbid oods at a karstic spring.Journal of Hydrol, 275, 1, 109 121. Massei, N., Dupont, J.P., Mahler, B.J., Laignel, B., Fourni er, M., Valdes, D. & S. Ogier, 2006: Investigating transport properties and turbidity dynamics of a karst aquifer using correlation, spectral, and wavelet analyses.Journal of Hydrology, 329, 1, 244. Milanovi, P.T., 2006: Karst of Eastern Herzegovina and Dubrovnik littoral.ZUHRA, pp. 362, Beograd. Mudarra, M. & B. Andreo, 2010: Hydrogeological func tioning of a karst aquifer deduced from hydrochem ical components and natural organic tracers present in spring waters. e case of Yedra Spring (South ern Spain).Acta Carsologica, 39, 2, 261. Perrin, J., Jeannin, P.Y. & F. Zwahlen, 2003: Implications of the spatial variability of inltration-water chem istry for the investigation of a karst aquifer: a eld study at Milandre test site, Swiss Jura.Hydrogeol ogy Journal, 11, 673. Petri, M., 2010: Characterization, exploitation, and pro tection of the Malenica karst spring, Slovenia.In: Kresic, N & Z. Stevanovic (eds.) Groundwater hy drology of springs. Engineering, eory, Manage ment, and Sustainability. Butterworth-Heinemann, pp. 428, Amsterdam. V ARIABILITY OF GROUND W ATER FLO W AND TRANSPORT PROCESSES IN KARST UNDER DIFFERENT HYDROLOGIC ...

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ACTA CARSOLOGICA 42/2-3 2013 338 Plan, L., Decker, K., Faber, R., W agreich, M. & B. Grase mann, 2009: Karst morphology and groundwater vulnerability of high alpine karst plateaus.Environ Geol, 58, 285. Pronk, M., Goldscheider, N. & J. Zop, 2006: Dynam ics and interaction of organic carbon, turbidity and bacteria in a karst aquifer system.Hydrogeology Journal, 14, 4, 473. Pronk, M., Goldscheider, N. & J. Zop, 2007: Particle-size distribution as indicator for fecal bacteria contami nation of drinking water from karst springs.Envi ronmental Science and Technology, 41, 8400-8405. Pronk, M., Goldscheider, N., Zop J. & F. Zwahlen, 2008: Percolation and Particle Transport in the Unsatu rated Zone of a Karst Aquifer.Ground W ater, 47, 3, 361. Raeisi, E., Groves, C. & J. Meiman, 2007: Eects of partial and full pipe ow on hydrochemographs of Logs don river, Mammoth Cave Kentucky USA.Journal of Hydrology, 337, 1. Ravbar, N., 2007: e protection of karst waters. ZRC Publishing, ZRC SAZU, pp. 256, Ljubljana. Ravbar, N. & N. Goldscheider, 2007: Proposed method ology of vulnerability and contamination risk map ping for the protection of karst aquifers in Slove nia.Acta Carsologica, 36, 3, 461. Ravbar, N. & N. Goldscheider, 2009: Comparative appli cation of four methods of groundwater vulnerabil ity mapping in a Slovene karst catchment.Hydro geology Journal, 17, 3, 725. Ravbar, N., Barber J.A., Petric, M., Kogovsek, J. & B. Andreo, 2012: Study of hydrodynamic behaviour of a complex karst system under low-ow conditions using natural and articial racers (springs of the Unica River, SW Slovenia).Environmental Earth Sciences, 65, 8, 2259. Ravbar, N., Engelhardt, I. & N. Goldscheider, 2011: Anomalous behaviour of specic electrical con ductivity at a karst spring induced by variable catchment boundaries: the case of the Podstenjek spring, Slovenia.Hydrological Processes, 25, 13, 2130. Ryan, M. & J. Meiman, 1996: An Examination of ShortTerm Variations in W ater Q uality at a Karst Spring in Kentucky.Ground W ater, 34, 23. Sass, I. & U. Burbaum, 2010: Damage to the historic town of staufen (Germany) caused by geothermal drillings through anhydrite-bearing formations.Acta Carsologica, 39, 2, 233. Savoy, L., 2007: Use of natural and articial reactive tracers to investigate the transfer of solutes in karst systems. PhD esis, University of Neuchtel, pp. 194. Shuster, E.T. & W .B. W hite, 1971: Seasonal uctuations in the chemistry of limestone springs: A possible means for characterizing carbonate aquifers.Jour nal of Hydrology, 14, 93. Vesper, D.J. & W .B. W hite, 2003: Metal transport to karst springs during storm ow: an example from Fort Campbell, Kentucky/Tennessee, USA.Journal of Hydrology, 276, 1, 20. W hite, W .B., 1988: Geomorphology and hydrology of karst terrains.University Press, pp. 464, New York, Oxford. W illiams, P.W ., 2008: e role of the epikarst in karst and cave hydrogeology: a review.International Journal of Speleology, 37, 1, 1. W orthington, S.R.H., 1991: Karst hydrogeology of the Ca nadian Rocky Mountains. PhD esis, McMaster University, 380 p. W orthington, S.R.H. & D. Ford, 2009: Self-organized permeability in carbonate aquifers.Ground W ater, 47, 326. Zhu, X.W ., Chen, W .H. & E. Lynch, 2007: W ulong karst systems and as an indicator of local tectonic upli.Carsologica Sinica, 2007. N ATA A R AVBAR



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From January 7 to January 11, 2013, the Karst W aters In stitute (KWI) and the National Cave and Karst Research Institute (NCRKI) held an international and multidisci plinary symposium on Carbon and Boundaries in Karst at NCKRI headquarters in Carlsbad, New Mexico. ere is growing interest in the dynamics of both inorganic and organic carbon in karst systems, and espe cially in the ux of carbon and nutrients between the sur face and subsurface, and between dierent components (e.g. epikarst and vadose zone) in the karst subsurface. is symposium was about these and other questions connected to carbon in karst and boundaries in karst. It was especially timely both because of rapid advances in the eld and the importance of carbon sequestration in global climate change e symposium highlighted re cent advances in biology, geology, and hydrology that are helping us understand the dynamics of karst ecosystems, especially with respect to carbon. e talks were orga nized around seven main themes: 2 Sixty participants from seven countries attended the week-long meeting which included an excursion to Carlsbad Caverns National Park. For the rst time at a KWI meeting, several participants, who were unable to attend in person, gave their presentations via Skype. e meeting was highlighted by two keynote presentations: Groundwater Ecology of Alluvial River Flood P lains, Jack Stanford, Flathead Lake Biological Station, Polson, Montana Karst Conduit Matrix Exchange and the Karst H yporheic Z one, John W ilson, New Mexico Institute of Mining and Technoloogy, Socorro, New Mexico. Two most distinguished karst scientists, W illiam B. W hite of Pennsylvania State University and Derek Ford of McMaster University jointly summed up the meeting. e following is a list of oral and poster presentations given at the meeting. Participants were invited to submit articles that elaborated their meeting presentations to Acta Carsologica. INTRODUCTION TO THE SYMPOSIUM David C. Culver, Guest Editor MEETING P RESENTATIONS Chemotrophy meets heterotrophy: the inverted 'critical zone' of the subsurface Penny J. Boaston Microbial controls on in situ production of dissolved organic matter Kathleen Brannen*, Annette Engel, and Ross Larson Redox state in karst aquifers: Impacts of DOCand DO-rich river water intrusion into Floridan aquifer springs Amy L. Brown*, Jonathan B. Martin, Elizabeth Screaton, John Ezell, James Sutton and Patricia Spellman Component isolation and lipid proling to characterize dissolved organic matter transformations along a groundwater ow path Terri Brown*, Susan M. Pner, and Annette S. Engel Using biominerals to assess anthropogenic inpact: a case study in Carter Salt Peter Cave, Carter County, TN Sarah K. Carmichael*, Mary J. Carmichael, Amanda Strom, Krissy W Johnson, Leigh Anne Roble, Yongli Gao, Cara M. Santelli, and Suzanna L. Bruer A simple theoretical framework to interpret spring variations and constrain mechanistic models of karst processes Matthew D. Covington Convergence and Divergence in Caves and Shallow Subterranean Habitats David C. Culver* and Tanja Pipan Microbial activities at geochemical interfaces in cave and karst environments Annette Summers Engel ACTA CARSOLOGICA 42/2-3, 173, POSTOJNA 2013

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ACTA CARSOLOGICA 42/2-3 2013 174 Interactions between surface and subterranean amphipods in springs Cene Fier Preliminary carbon sequestration and denudation rates within the karst of the Cumberland Plateau, USA Lee J. Florea Determinants of macroinvertebrate diversity in karst springs of the Mid-Atlantic region, USA Daniel W Fong*, Christopher Seabolt, and Kaitlin C. Esson Bicarbonate water chemistry of Little Limestone Lake, a beautiful marl lake in Manitoba, Canada Derek Ford e relative importance of speleogenetic phases as revealed by numerical models Franci Gabrovek Dynamics and limitations of organic carbon turnover in porous aquifers Christian Griebler e longitudinal response of benthic invertebrate communities to caves Jonathan S. Harding* and Troy W atson Experimental design and instrumentation to observe karst conduit hyporhiec ow Katrina K. Henry*, Kenneth A. Salaz, and John L. Wilson Biological control on acid generation at the conduitbedrock boundary in submerged caves Janet S. Herman*, Alexandria G. Hounshell, Rima B. Franklin, and Aaron L. Mills Environmental controls on organic matter production and transport across surfacesubsurface and geochemical boundaries in the Edwards Aquifer, Texas, USA Benjamin T. Hutchins*, Benjamin F. Schwartz, and Annette S. Engel Subaerial microbial life in the suldic Frasassi Cave System, Italy Daniel S. Jones*, Irene Schaperdoth, and Jennifer L. Macalady Physical Structure of the epikarst W illiam K. Jones Stratigraphic control on conduit development in the Ozark Karst, Missouri, USA James E. Kaufmann* and Jeery Crews Using isotopes of dissolved inorganic carbon species and water to separate sources of recharge in a cave spring, northwestern Arkansas Katherine J. Knierim*, Erik Pollock, and Phillip D. Hays Quantitatively modeling source inuences on cave air carbon dioxide chemistry Andrew J. Kowalczk Quaternary glacial cycles: karst processes and the global CO 2 budget Erik B. Larson* and John E. Mylroie Karst in the global carbon cycle Jonathan B. Martin Mitra Khadka, Marie Kurz, John Ezell, Amy Brown Spatio-temporal trends in diversity of subsurface assemblages from the vadose zone of the Carpathian karst in Romania Ioana N. Meleg Comparison of water quality in submerged caves with that of diuse groundwater immediately proximal to the conduit Aaron L. Mills*, Janet S. Herman, and Terrence N. Tysall Carbon cycling in arid land caves: implications for microbial processes Diana E. Northup*, Noelle G. Martnez, Lory O. Henderson and Elizabeth T. Montano Shallow Subterranean Habitats in Volcanic Terrain Pedro Orom* 1 and Heriberto D. Lpez 1,2 Particulate inorganic carbon ux in karst and its signicance to karst development and the carbon cycle Randall L. Paylor* and Carol M. W icks Patterns of organic carbon in shallow subterranean habitats (SSHs) Tanja Pipan* and David C. Culver

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ACTA CARSOLOGICA 42/2-3 2013 175 Seasonal, diurnal and storm-scale PCO 2 variations of cave stream in subtropical karst area, Chongqing, SW China Junbing Pu*, Daoxian Yuan, Licheng Shen and Heping Zha Variability of groundwater ow and transport processes in karst under dierent hydrologic conditions Nataa Ravbar Wheres the re? An analysis of carbon precipitates in Black and other caves of the Upper Guadalupe Mountains, New Mexico Sam Rochelle*, Michael N. Spilde, and Penny J. Boston Using hydrogeochemical and ecohydrologic responses to understand epikarst processes in semi-arid systems, Edwards Plateau, Texas, USA Benjamin F. Schwartz*, Susanne Schwinning, Brett Gerard, Kelly R. Kukowski, Chasity L. Stinson, and Heather C. Dammeye Carbon ux in the Dorvan-Cleyzieu karst: lessons from the past to guide future research Kevin S. Simon Groundwater ecology of alluvial river ood plains Jack A. Stanford Seasonal inux of organic carbon into Marengo Cave, Indiana, USA Philip van Beynen*, Derek Ford and Henry Schwarcz Testing carbon limitation of a cave stream ecosystem using a whole-reach detritus amendment Michael P Venarsky*, Brock M Huntsman, Jonathan P Benstead, Alexander D Huryn e role of karst conduit morphology, hydrology, and evolution in the transport, storage, and discharge of carbon and associated sediments George Veni Carbon uxes in karst aquifers: sources, sinks, and the eect of storm flows W illiam B. W hite Hydrograph interpretation changes in time Carol W icks Karst conduit-matrix exchange and the karst hyporheic zone John L. W ilson e role of geological processes in global carbon cycle: a review Yuan Daoxian e stability of carbon sink eect related to carbonate rock dissolution: a case study of the Caohai Lake geological carbon sink Zhang Qjang

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e initiator of this issue was professor David Culver, who suggested the publication of papers presented at the multidisciplinary symposium on Carbon and Boundaries in Karst in our journal. Being familiar with the high qual ity of past meetings organized by the Karst W aters Insti tute, the editorial board of AC agreed with the proposal and invited Dave as a guest editor. His editorial work was highly ecient and thorough; he found relevant reviewers and provided a high quality reviews for each manuscript. e issue contains high quality review and original research paper, presenting a comprehensive cov erage of the role of karst in the global carbon cycle. is issue would not be possible without a wide cooperation of reviewers who provided thorough and thoughtful re views of all manuscripts. Several manuscripts have been rejected for dierent reasons, and most of the others were considerably improved aer the review. Even though the review process was anonymous, we present the list of those reviewers that have agreed to be acknowledged in alphabetical order: Pavel Bosk, Annette S Engel, Derek Ford, Christian Griebler, Ellen Herman, Janet S Herman, W illiam K Jones, Alexander Klimchouk, Florian Malard, Pierre Marmonier, Jonathan B Martin, MaryLynn Mus grove, John E Mylroie, Diana Northup, Metka Petri, Tanja Pipan, Nataa Ravbar, Benjamin Schwartz, Kevin Simon, Branka Trek, Michael Venarsky, George Veni and W illiam B W hite. W e hope that the readers will enjoy reading the is sue. Franci Gabrovek, the editor EDITORIAL ACTA CARSOLOGICA 42/2-3, 176, POSTOJNA 2013



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C ARBON FLU X ES IN K ARST A QUIFERS : S OURCES SINKS AND THE EFFECT OF STORM FLO W T OK OGLJIKA V KRA KIH VODONOSNIKIH: IZVORI, PONORI IN UINKI POPLAVNIH TOKOV W illiam B. WHITE 1 Izvleek UDK 546.26:551.444 William B. White: Tok ogljika v krakih vodonosnikih: izvori, ponori in uinki poplavnih tokov Vsebnost ogljika (carbon loading [ g/l]) v krakih vodah lahko doloimo z meritvijo alkalnosti in pH. Vsebnost je neodvisna od stopnje nasienja in stanja ravnoteja vode glede na karbonatno matino kamnino. Za vrednotenje izvorov, prenosa in skladienja CO 2 smo uporabili obseno bazo geokemijskih podatkov iz ve krakih sistemov v ZDA. Ti podatki so bili pri dobljeni v okviru razlinih tudentskih del v zadnjih 40 letih. Vsebnost ogljika v krakih vodah je odvisna od intenzivnosti produkcije CO 2 v prsti krakega zaledja in razmerja med delom CO 2 ki se porabi za raztapljanje karbonatov in delom CO 2 ki se z difuzijo vrne nazaj v atmosfero. Pri tem so pomembni vpli vi vegetacije, tipa prsti in dele vode, ki jo v zaledje prinesejo ponornice. Izgube CO 2 v atmosfero nastajajo z izloanjem sige, z razplinjanjem CO 2 na izvirih in izloanjem lehnjaka. Izvori, transport in ponori CO 2 so tudi sezonsko pogojeni. e povzamemo, je kraki vodonosnik neto, vendar puajoi, ponor CO 2 Kljune besede: CO 2 kraki vodonosnik, izviri, vsebnost oglji ka. 1 Department of Geosciences, e Pennsylvania State University, Deike Building, University Park, PA 16802 USA, e-mail: wbw2@psu.edu Received/Prejeto: 13.2.2013 COBISS: 1.02 ACTA CARSOLOGICA 42/2-3, 177, POSTOJNA 2013 Abstract UDC 546.26:551.444 William B. White: Carbon uxes in Karst aquifers: Sources, sinks, and the eect of storm ow An eective carbon loading can be calculated from measured alkalinity and pH of karst waters. e carbon loading is inde pendent of the degree of saturation of the water and does not depend on the water being in equilibrium with the carbonate wall rock. A substantial data base of spring water analyses ac cumulated by students over the past 40 years has been used to probe the CO 2 generation, transport, and storage in a variety of drainage basins that feed karst springs. Carbon loading in the water exiting karst drainage basins depends on the rate of CO 2 generation in the soils of the catchment areas and on the partitioning between CO 2 dissolved in inltration water and CO 2 lost by diusion upward to the atmosphere. For any given drainage basin there are also inuences due to vegetative cover, soil type, and the fraction of the water provided by sink ing stream recharge. Losses of CO 2 back to the atmosphere oc cur by speleothem deposition in air-lled caves, by degassing of CO 2 in spring runs, and by tufa deposition in spring runs. ere are seasonal cycles of CO 2 generation that relate growing season and contrasts in winter/summer rates of CO 2 genera tion. Overall, it appears that karst aquifers are a net, but leaky, sink for atmospheric CO 2 Keywords: CO 2 karst aquifers, springs, carbon loading. I NTRODUCTION Concern over rising CO 2 levels in the atmosphere and their probable inuence on climate has prompted an in tense interest in the global cycle of CO 2 including both natural and anthropogenic sources and sinks. Rock weathering by chemical reactions of silicate minerals with CO 2 -bearing solutions has been identied as a CO 2

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ACTA CARSOLOGICA 42/2-3 2013 178 WILLIAM B. WHITE sink and therefore a site of carbon sequestration (Gail lardet et al. 1999; Franck et al. 1999; Kump et al. 2000; Moosdorf et al. 2011). Rock weathering in carbonate ter rain remains somewhat enigmatic. Dissolution of lime stone and dolomite by CO 2 -bearing ground waters con sumes CO 2 and has been proposed as an explanation for about one fourth of an unknown sink in the continental biosphere (Gombert 2002). However, other karstic proc esses can release the CO 2 back into the atmosphere with the net eect of transferring carbon but not sequester ing it. As a consequence, some authors discount carbon ate rock weathering as an important contributor to the carbon ux. Earlier estimates (Liu & Zhao 2000) assign about one fourth of the carbon ux in karst terrain to a carbon sink with the remaining three fourths recycled by rainfall and atmospheric circulation. Later calculations by Liu et al. (2011) argue that the CO 2 sink from carbon ate rock weathering has been underestimated. ere have been many investigations of the chem istry of karst waters but most are concerned with satu ration state and the dissolution and precipitation of carbonate minerals. e present study uses the same data recalculated as total dissolved inorganic carbon to probe the various controlling inuences and to deter mine whether it is possible to tally the karst components of the CO 2 budget. ese calculations apply to carbon uxes in individual karstic drainage basins and do not address carbon uxes on a global scale. However, when the sources and sinks are added up, should karst aquifers be included as a CO 2 sink in the global carbon budget? CARBON FLU X ES IN KARST A QUIFERS: THE CONCEPTUAL MODEL Supercially, the function of karst processes in CO 2 trans port and sequestration can be described by the reaction CaCO 3 + CO 2 + H 2 O 2HCO 3 + Ca 2+ (1) For each mole of calcite dissolved, a mole of CO 2 is extracted from the environment and a mole of CO 2 is released from long-term sequestration in the car bonate rock. The carbon is transported through the karst drainage basin mainly as dissolved CO 2 and bicarbonate ions carried in solution to a spring and then taken out of the basin by surface streams. Un fortunately, the reaction above and other reactions in carbonate chemistry are easily reversible. The actual carbon flux through karst aquifers is much more com plicated and involves both sources and sinks. Most importantly with regard to karst aquifers as overall CO 2 sinks is that the systems are leaky and there are multiple pathways by which CO 2 consumed in car bonate reactions can be returned to the atmosphere. Fig. 1 gives a schematic of the most important interac tions and flow paths. Fig. 1: Flow sheet for the move ment of carbon dioxide through karst aquifers. Green lines repre sent a ux of CO 2 into the aqui fer; red lines represent loss of CO 2 back to the atmosphere. Green lines with a parallel dashed blue line are pathways in which the carbon is transported in solution.

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ACTA CARSOLOGICA 42/2-3 2013 179 C ARBON FLU X ES IN K ARST A QUIFERS : S OURCES SINKS AND THE EFFECT OF STORM FLO W e background reservoir for CO 2 budgets is the atmosphere. e CO 2 concentration in the atmosphere as of 2012 is 0.0395 volume percent. is value is drawn from the CO 2 records collected at Mauna Loa observa tory, Hawaii. e Mauna Loa data are available on the National Oceanographic and Atmospheric Administra tion website. Atmospheric carbon dioxide dissolves in atmo spheric moisture and is carried to the ground by rainfall. Even this seemingly simple process is more complicated than it seems. Some portion of the rainfall evaporates taking both the water and the dissolved CO 2 back into the atmosphere. Some portion is taken up by growing plants with a portion of the water returned to the atmo sphere by transpiration. e CO 2 in this portion may be retained by the plants as part of their biomass. Finally, a third portion of the rainfall inltrates deeper into the soil, carrying its load of dissolved CO 2 deeper into the epikarst. Direct extraction from the atmosphere is not the primary source of CO 2 driving carbonate rock dissolu tion. CO 2 concentrations in soils typically range from fractions of a percent to as high as ten percent due to CO 2 from plant roots, from microbial activity, and from decaying vegetation. us the carbon ux through karst aquifers is, in large part, spun o from the vegetative car bon cycle. Photosynthesizing plants draw CO 2 from the atmo sphere and sequester the carbon temporarily as cellulose and other plant components (their bulk composition represented as CH 2 O). In rough mass-balance terms: CO 2 + H 2 O CH 2 O + O 2 (2) W hen the vegetation dies, the cellulose oxidizes and decays (the reverse of equation (2)), releasing the stored carbon as CO 2 e decay of plant material in the epikarst provides the primary source of carbon that is fed into the karst system. A fraction of the CO 2 withdrawn from the atmosphere by plant growth is transferred to the karst system and not discharged back to the atmosphere thus breaking what would otherwise be a closed cycle. e main part of the dissolutional denudation of karst takes place where inltration water from the epikarst reaches the underlying carbonate bedrock. Here is where dissolved carbon dioxide extracts additional carbon from the carbonate rocks and releases the reac tion products, mostly the bicarbonate ion, into water descending through the vadose zone. If the pathways through the vadose zone consist mostly of joints and fractures, there will be little gain or loss of CO 2 as the water descends to the water table although it is possible that organic matter, carried downward through open fractures, provides a CO 2 source at depth (W ood, 1985). If the descending vadose water happens to intersect an open cave passage in which the CO 2 pressure is less than that of the vadose water, CO 2 will be degassed resulting in the re-precipitation of dissolved calcite. Half of the transported carbon will be re-sequestered as speleothem calcite but the other half will be released into the cave at mosphere. CO 2 in the cave atmosphere can be dissolved in cave stream waters if these waters are not already satu rated or may be removed to the surface atmosphere by diusion or by circulating air currents. CO 2 is also carried into karst aquifers by sink ing surface streams. Concentrations of CO 2 depend on the source of the water in the sinking surface stream whether the catchments are made up of forest, grassland, agricultural land, or urban areas, and also what fraction of the stream ow is contributed by ground water and what fracture by overland runo. Spring discharges are not the end of the story for carbon transport. Some coastal karst springs discharge directly into the ocean, some are located on the banks of major rivers, but most karst springs feed into small tributaries (spring runs) that ow as surface streams for considerable distances before linking to the main paths of the surface drainage system. e spring run must be considered part of the karst system because the water of ten undergoes chemical changes downstream from the spring. If the spring water, even if undersaturated, has a calculated CO 2 partial pressure higher than that of the atmosphere, CO 2 will be degassed, the pH and saturation index will rise, and some of the dissolved CO 2 will be returned to the atmosphere. If the saturation index be comes positive, calcium carbonate may be deposited as tufa, returning still more CO 2 to the atmosphere. e conceptual model presented in Fig. 1 ends where the spring run leaves the carbonate basin. ere are more sources and sinks for carbon along the route through the surface river system to the ocean. In the ocean, the dissolved carbonates can be precipitated by marine organisms, again releasing CO 2 that contributes to ocean acidication. Although necessary for globalscale models, downstream processes are not considered further in this paper.

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ACTA CARSOLOGICA 42/2-3 2013 180 e carbonate geochemistry needed for these calcula tions is discussed in detail in W hite (1988), Langmuir (1997), or Drever (1997). e total inorganic carbon in solution is simply the sum of all dissolved carbon species. [ Carbon ] = [ CO 2 ( aq )] + [ H 2 CO 3 ] + [ HCO 3 ] + [ CO 3 2 ] (3) where the quantities in brackets are molar concen trations. e rst two terms are taken together as H 2 CO 3 e concentration of the carbonate ion in the pH range of karst waters is usually negligible. e combined aqueous CO 2 and neutral carbonic acid can be calculated from alkalinity and pH (4) e bicarbonate ion concentration, [HCO 3 ] in molar units, and the pH were extracted from the avail able data bases. e ionization constant, K 1 of carbonic acid is dependent on temperature; a temperature-tting function provided by Langmuir (1997) was used in all calculations. In addition, the calculation requires the activity coecient, which can be calculated from the ionic strength using the Debye-Hckel relationship if an estimate can be obtained of the ionic strength Ionic strength can be easily calculated from the spe cic conductance, a parameter that is frequently mea sured as a proxy for hardness and total dissolved solids (e.g. Krawczyk & Ford 2006). A new estimate for the relationship was obtained from a data base (Jacobson, 1973) containing analyses from 60 karst springs, lime stone wells, and dolomite wells in the Ordovician car bonate rocks in central Pennsylvania, USA. Measured ion concentrations and specic conductance allowed the calculation of a linear regression (Eqn. 5) that t all of the data to an R 2 of 0.99. I = 1.697 x 10 Spc .7 x 10 (5) Ionic strength is in moles/L; Spc is in S/cm. e data points derived from springs, limestone wells, and dolomite wells all had about the same small scatter around the same linear tting line showing that the re lationship is independent of the Ca/Mg ratio and satura tion state of the waters. Substituting (3) into (4) and neglecting CO 3 2, we obtain the total concentration of dissolved carbon in molar units. (6) Multiplying [Carbon] by 12.011, the atomic weight of carbon, gives the carbon loading in units of g/L (or kg/m 3 ). e total carbon ux, F C in units of kg/day, is then given by (7) where Q is the spring discharge is m 3 /sec and the bicarbonate ion concentration, C HCO3 is in units of mg/L, the value that is typically reported. e numerical coe cient contains atomic and molecular weights and neces sary unit conversions (12.011 x 86,400)/(1000 x 61.016). It may be noted that the calculated carbon con centrations and uxes do not depend on the degree of saturation with respect to calcite or dolomite. It is thus possible to meaningfully compare seepage and drip wa ters that are oen saturated with carbonates with spring and cave stream waters which are frequently undersatu rated. CALCULATION OF CARBON LOADING AND FLUXES CARBON FLU X ES W ITHIN KARST A QUIFERS CARBON IN THE EPIKARST As a source, the epikarst is remarkably complicated. CO 2 is produced by exhalation through plant roots, by decaying organic material, and other microbial proc esses. e rate of CO 2 production depends of soil char acteristics, on the vegetative cover, and is sensitive to both moisture conditions and temperature. Only a por tion of the generated CO 2 is swept into the subsurface by inltrating precipitation; much is lost by upward diusion back into the atmosphere. e many com plications of soil CO 2 are discussed in some detail by Jakucs (1977). ere have been several eorts to model CO 2 production on both global scale (Brook et al. 1983; Gunn 1984) and site-specic scales with coupled rate WILLIAM B. WHITE

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ACTA CARSOLOGICA 42/2-3 2013 181 equations ( imnek & Suarez 1993; Suarez & imnek 1993). e seasonal variability of soil CO 2 is illustrated with the comprehensive data of De Jong and Schappert (1972) (Fig. 2). eir data were collected in a non-karstic prairie soil in Canada. Several features stand out. CO 2 concentrations range up to more than one percent dur ing the growing season but fall to only a few times at mospheric background as the growing season fades into winter. e CO 2 concentration decreases with depth and the seasonal oscillation decreases and shis to later in the year. At 90 cm depth, the seasonal variation has es sentially disappeared. A similar range was observed in a forest soil in the sub-tropical climate of Pigeon Moun tain, Georgia (Kiefer 1990) where low CO 2 concentra tions in the non-growing season were also observed (Dyer & Brook 1991). Data collected from a wide variety of locations gave similar results (Miotke 1974). Tropical soils tend to contain higher CO 2 concentrations, 5 vol ume percent or higher in some locations measured in Malaysia (Crowther, 1984). e many measurements of soil CO 2 concentration represent the steady-state result of three competing pro cesses: (1) the rate of production of CO 2 by a variety of processes, Pr, (2) the rate at which CO 2 is dissolved in inltration water and carried down to the reaction zone at the bedrock contact, F i and (3) the rate at which CO 2 diuses upward to be lost to the atmosphere, F D d Pr dt d Fidt d FDdt (8) e coecients a and b describe the relative pro portions of CO 2 lost by the two routes, parameters that are generally not known. ey would describe the frac tion of CO 2 available for dissolution of limestone and thus potentially sequestered. e ux, F D represents carbon returned to the growing plant/dead plant/decay closed cycle. SINKING STREAM INPUTS Sinking streams contain both surface runo and ground water discharging from the stream banks. If the CO 2 ux from groundwater discharging along the bank is high, the owing free surface stream may degas CO 2 before the stream reaches its swallet. Sinking streams usually input their water and dissolved CO 2 load directly into the con duit system where it moves rapidly through the subsur face to a spring Relatively few analyses of sinking stream waters are available. Data from three such streams are presented in Fig. 3. e Logan Gap stream, in central Pennsylvania, ows across Silurian sandstones and shales, then sinks into a swallet at the contact with Ordovician carbon ates. Little Sinking Creek rises on the carbonate-rich Salem and W arsaw formations on the southern edge of the Sinkhole Plain in the Central Kentucky Karst and ows on carbonates for some distance before sinking into the lower St. Louis Limestone. Rio Camuy rises on the Tertiary volcanics that make up the core of the Island of Puerto Rico. e river sinks into Tertiary carbonates where it ows through the Rio Camuy Cave System to a resurgence cave and then as a surface river to the coast. e carbon loading in sinking streams is very strongly dependent on the details of the specic catch ment. Mountain streams from sandstone and shale carry only a few mg/L of carbon. Even a tropical basin carries little carbon from a catchment located on volcanic rocks. Fig. 2: Seasonal variation in CO 2 concentration in a prairie soil measured at the designated depths. Days are counted from J anu ary 1. Raw data taken from de J ong and Schappert (1972). Fig. 3: Dissolved carbon loading for three sinking streams. All measurements made near the stream swallets. Logan Gap: 19611962; Little Sinking Creek: 1972-1973 (H ess 1974); Rio Camuy: 1980-1984 (Troester 1994). Data are arranged by J ulian day al though not all data were collected in the same year. C ARBON FLU X ES IN K ARST A QUIFERS : S OURCES SINKS AND THE EFFECT OF STORM FLO W

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ACTA CARSOLOGICA 42/2-3 2013 182 e carbon loading of both non-carbonate rock streams is essentially constant with respect to season, tempera ture, and discharge. e dissolved carbon in Little Sink ing Creek, a partly carbonate stream, exhibits a seasonal cycle with a minimum in late March and a maximum in early September. CARBON RELEASE BY SPELEOTHEM DEPOSITION e CO 2 partial pressures in air-lled caves in the vadose zone are typically ten or more times higher than the at mospheric CO 2 partial pressure and vary widely between caves and between locations within the same cave (Ek & Gewalt 1985; Kempe et al. 1998). ere is a seasonal variation in the CO 2 concentration in cave atmospheres that reects the seasonal cycle in CO 2 production in the epikarst (Troester & W hite 1984). e vadose seepage waters percolating downward from the epikarst typi cally have CO 2 partial pressures that are higher than cave atmospheres. As a result, vadose seepage waters de-gas CO 2 into the cave atmosphere with concurrent deposi tion of calcite (or aragonite) onto growing speleothems. e fraction of the CO 2 in the cave air that makes its way back to the surface atmosphere represents one of the leaks in the karst system. e observed CO 2 concentra tions represent a balance between the rate of CO 2 degas sing from seepage waters and the rate of CO 2 loss by air exchange with the surface. Because of the application of speleothems as paleo climate archives, there is much contemporary interest in dripwater. Measurements are dicult because the chem istry, particularly the CO 2 content, changes rapidly by de gassing when the drips emerge into the cave atmosphere. In a classic paper, Holland et al. (1964) measured the water chemistry of 47 drip points and drip pools in Lu ray Caverns, Page County, Virginia. A selection of these measurements that were suciently complete for calcu lation gave dissolved carbon loadings in the range of 70 to 140 mg/L. Dripwater chemistry from Cave W ithout A Name, Kendall County, Texas (Veni 1997) (Fig. 4) pro duced a cool season carbon loading only slightly greater than that of the sinking streams, ere was a warm sea son maximum in this relatively arid region that reached to the lower end of the range measured in Luray Caverns, in an area with higher rainfall and thicker soils. Two factors indicate that speleothem precipitation is not a major loss term in the carbon ux. (1) Although the seepage waters carry a high carbon loading, only a fraction of the load is deposited as speleothem calcite. (2) Air-lled conduits intercept only a small fraction of the water and dissolved carbon descending from the epikarst. W orthington et al. (2000) have estimated that conduits underlie only on the order of one percent of karstic aquifers. Although the conduit systems act as master drains for the aquifer relatively little of the inl tration water from the epikarst passes through an open conduit. Fig. 4: Dissolved carbon loading in dripwater from Cave Without a Name, Kendall County, Texas. Segments of two J ulian years are included. Original data drawn from V eni (1997). CARBON FLU X ES IN KARST SPRINGS SEASONAL CYCLES Rarely does an investigation of karst springs provide chemical analyses of the water matched with quantita tive measurement of discharge. An exceptionally good data set was collected from Tippery Cave, Huntingdon County, Pennsylvania (Hull 1980) that provides 69 oneweek intervals with carefully measured discharge (Fig. 5). Tippery Spring is fed by a conduit system resulting in a hydrograph with many individual storm peaks. e hy drograph peak at week 47 (February 17, 1976) reached a discharge of 1088 L/sec. In contrast, the carbon load ing was only weakly dependent on discharge. e intense peak at week 47 produced only a dilution eect of about 50% of the carbon loading. e carbon loading shows a broad seasonal maximum at about week 22 (late Au gust). WILLIAM B. WHITE

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ACTA CARSOLOGICA 42/2-3 2013 183 Rio Camuy in Puerto Rico has no signicant seasonal dependence. Rather than corresponding to the maxima in the temperate climate karst data, the Rio Camuy data are very similar to the loadings measured for the other springs during the winter season. STORM FLO W S Storm water moves though the conduit system of karst aquifers as a pulse of fresh surface water. Storm eects on aquifer chemistry are varied. Storm pulses dilute the dissolved carbonates resulting in a dip in hardness dur ing storm events. In contrast, storm pulses may ush soil contaminants resulting spikes of nitrate or agricultural chemicals during storm events. ere appears to be only a small response of the carbon ux to storm pulses in the data collected at Tippery Spring. e intense storm shown on Fig. 5 where the discharge rose to about 50 times the base ow value resulted in only about a factor of two decrease in the carbon loading. Rock Spring, Centre County, Pennsylvania hap pened to be gauged and was being sampled on a regu lar basis when the Hurricane Agnes storm swept across Pennsylania in June, 1972 (Jacobson, 1973). Rock Spring is fed by a strike-oriented conduit with a large recharge from mountain runo. e Agnes storm produced a very ashy hydrograph (Fig. 7-A). e carbon loading, however, dipped only slightly as the storm pulse passed (Fig. 7-B). e carbon ux, therefore, is dominated by the storm hydrograph and also exhibits a large peak co incident with the hydrograph peak (Fig. 7-B). SPRING RUNS, CO 2 DEGASSING, AND TUFA DEPOSITION Karst spring water is usually undersaturated when it is drains from a conduit system and may be close to satura tion when it drains from a fracture aquifer. In most spring waters, the carbon dioxide pressure is well above the at mospheric background. As the water from the spring ows down the spring run, carbon dioxide is degassed, the pH of the water rises as does the saturation index. An example (Shuster 1970) (Fig. 8), measured on Elk Creek which rises from a large conduit-fed spring, shows that pH rises in an S-shaped curve reaching a steady state value about 800 m below the spring mouth. Ca 2+ and bi carbonate ion concentrations remain constant implying that dissolved CO 2 is lost but no carbonate has precipitat ed. e saturation index also rises on an S-shaped curve reaching a steady state value of about +0.3, less than the +0.5 supersaturation typically needed to nucleate calcite precipitation. e concentration of dissolved carbon at the spring mouth was calculated to be 25.9 mg/L which decreased to 24.2 mg/L downstream. For this specic example, e seasonal eect in the carbon loading is dem onstrated with data on three karst springs (Fig. 6). e Penns Cave rise is the discharge point for a large car bonate basin that includes both inltration and sinkhole inputs in a karst valley and a component of sandstone mountain runo from streams that sink along the valley margins. ere is a broad maximum that peaks at about Julian day 250 (early September). A much more promi nent seasonal maximum occurs in the carbon loading of Graham Spring. Graham Spring, near Bowling Green, Kentucky, drains a large area of the Sinkhole Plain south of Mammoth Cave. ere is some sinking stream input but most of the water originates from inltration and sinkhole drains in the karst. e maximum is shied slightly to late September compared with the Pennsylva nia data. In contrast, the carbon loading in the rise of the Fig. 5: H ydrograph and carbon loading for Tippery Spring, H unt ingdon County, P ennsylvania. Week 1 begins April 1, 1975. Note that the hydrograph peak on week 47 is for a severe storm with a discharge of 1088 L/sec. Original data drawn from H ull (1980). Fig. 6: Carbon loadings in three large karst springs. Data for P enns Cave Rise from Shuster (1970). Data for Graham Spring from H ess (1974). Data for the Rise of the Rio Camuy from Troester (1994). C ARBON FLU X ES IN K ARST A QUIFERS : S OURCES SINKS AND THE EFFECT OF STORM FLO W

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ACTA CARSOLOGICA 42/2-3 2013 184 Fig. 7: (A) H ydrograph for Rock Spring, Centre County, P ennsylvania during the H urricane Agnes storm in J une, 1972. Elapsed days are measured from J une 1. (B) Carbon loading and carbon ux from Rock Spring during H urricane Agnes. Original data drawn from J acobson (1973). 6-7 % of the dissolved carbon was lost to the atmosphere by CO 2 -degassing. Many karst streams, particularly those in tropical regions, precipitate excess carbonate as tufa deposits (Ford & Pedley 1996). is reverses the carbonate disso lution chemistry and releases CO 2 into the atmosphere. Tufa deposition is an additional loss term in the overall carbon balance. CO 2 degassing from spring runs raises the pH and the saturation index. If the CO 2 pressure in the stream water reaches the atmospheric value before the saturation index reaches the critical value for car bonate precipitation, the reaction will stop, the steadystate condition illustrated in Fig. 8. However, if the CO 2 partial pressures and dissolved carbonate concentra tions are high, as is typical of tropical karst, the satu ration index will reach the critical value before all ex cess CO 2 is degasses and carbonates will be precipitated (W hite 1997). Fig. 8: Changes is pH and carbonate species along the spring run of Elk Creek Rise in Centre County, P ennsylvania. From Shuster (1970). CONCLUSIONS It has been suggested that the reaction of carbon dioxide with carbonate rocks in karst aquifers is a potential sink for atmospheric CO 2 and should be considered by those who are constructing global carbon cycles. is paper ex amines some of the details in the ux of carbon through karst aquifers. e ux of carbon leaving a karst drainage basin can be described in mass balance terms: WILLIAM B. WHITE A B

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ACTA CARSOLOGICA 42/2-3 2013 185 ACKNO WLEDGEMENTS As cited in the text, this study draws heavily on data collected in the course of the thesis research of many Penn State students who, themselves, had quite dier ent objectives. eses are a useful repository of raw data which rarely appear in the published versions of the dis sertation research. FLUX Exit = [C Epk + C Atm ] Q i + C SS Q SS C Spm Q i f Con C Run Q Spr C Tuf Q Spr (9) e C terms are carbon loadings in units of mass/ unit volume with subscripts Epk = epikarst, Atm = at mospheric, Spm = speleothem, SS = sinking stream, Run = spring run, Tuf = tufa. e Q terms are water ows in units of volume/unit time. Q i = overall inltration re charge into the karst aquifer; Q SS is total recharge from sinking streams; Q Spr is the discharge from the springs draining the karst aquifer. f Con is the fraction of the to tal karst area underlain by air-lled conduits. e carbon loading terms can usually be estimated from available geochemical data. Data for the Q-terms are sparse. Carbon loadings vary with season in temperate karst but show little variation in tropical karst. ey are strongly dependent on the details of the specic karst drainage basin but are only weakly dependent on dis charge. Storm ows dilute the carbon loadings but usu ally by a factor of two or less. Carbon uxes essentially scale with discharge. Speleothem precipitation is a loss term but a small one because of the small fraction of the aquifer that is actually underlain by air-lled conduits. Degassing from spring runs represents a loss of 5 10 percent of the car bon load with an additional loss if the spring run pre cipitates tufa deposits. Overall, it appears that although karst systems are leaky in the sense that a fraction of the extracted CO 2 is returned to the atmosphere, they do represent a carbon sink that should be considered in global models. REFERENCES Brook, G.A., Folko, M.E. & E.O. Box, 1983: A world model of soil carbon dioxide.Earth Surface Pro cesses and Landforms, 8, 79. Crowther, J., 1984: Soil carbon dioxide and weathering potentials in tropical karst terrain, peninsular Ma laysia: A preliminary model.Earth Surface Pro cesses and Landforms, 9, 397. De Jong, E. & H.J.V. Schappert, 1972: Calculation of soil respiration and activity from CO 2 proles in the soil.Soil Science, 113, 328. Drever, J.I., 1997: e Geochemistry of Natural Waters. Prentice Hall, pp. 436, Upper Saddle River, NJ. Dyer, J.M. & G.A. Brook, 1991: Spatial and temporal variations in temperate forest soil carbon dioxide during the non-growing season.Earth Surface Processes and Landforms, 16, 411. Ek, C. & M. Gewelt, 1985: Carbon dioxide in cave atmo spheres. New results in Belgium and comparison with some other countries.Earth Surface Processes and Landforms, 10, 173. Ford, T.D. & H.M. Pedley, 1996: A review of tufa and travertine deposits of the world.Earth Science Re views, 41, 117. Franck, S., Kossacki, K. & C. Bounama, 1999: Modelling the global carbon cycle for the past and future evo lution of the earth system.Chemical Geology, 159, 305. Gaillardet, J., Dupr, B., Louvat, P. & C.J. Allgre, 1999: Global silicate weathering and CO 2 consumption rates deduced from the chemistry of large rivers.Chemical Geology, 159, 3. Gombert, P., 2002: Role of karstic dissolution in global carbon cycle.Global and Planetary Change, 33, 177. Gunn, J., 1984: A world model of soil carbon dioxide: A discussion.Earth Surface Processes and Land forms, 9, 83. Hess, J.W ., 1974: H ydrochemical investigations of the cen tral Kentucky karst aquifer system. Ph.D. esis, e Pennsylvania State University, pp. 218. C ARBON FLU X ES IN K ARST A QUIFERS : S OURCES SINKS AND THE EFFECT OF STORM FLO W

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ACTA CARSOLOGICA 42/2-3 2013 186 Holland, H.D., Kirsipu, T.V., Huebner, J.S., & U.M. Ox burgh, 1964: On some aspects of the chemical evolution of cave waters.Journal of Geology, 72, 36. Hull, L.C., 1980: Mechanisms controlling the inorganic and isotopic geochemistry of springs in a carbonate terrain. Ph.D. esis, e Pennsylvania State Uni versity, pp. 259. Jacobson, R.L., 1973: Controls on the quality of some car bonate ground waters: Dissociation constants of cal cite and CaH CO 3 + from 0 to 50 C. PhD thesis. e Pennsylvania State University, pp. 131. Jakucs, L., 1977: Morphogenetics of Karst Regions. John W iley, pp. 284, New York, NY. Kempe, S., Neander, F., Hartmann, J. & W Klughart, 1998: CO 2 -Druck der Lu in der Hhle ohne Na men, Steinamwasser (Nrdliche Frankenalb).Mit teilungen der Verbandes der deuschen Hhlenund Karstforscher, 44, 2, 39. Kiefer, R.H., 1990: Soil carbon dioxide concentrations and climate in a humid subtropical environment.Professional Geographer, 42, 182. Krawczyk, W .E. & D.C. Ford, 2006: Correlating specic conductivity with total hardness in limestone and dolomite karst waters.Earth Surface Processes and Landforms, 31, 221. Kump, L.R., Brantley, S.L. & M.S. Arthur, 2000: Chemi cal weathering, atmospheric CO 2 and climate.An nual Reviews of Earth and Planetary Science, 28, 611. Langmuir, D., 1997: Aqueous Environmental Geochemis try. Prentice Hall, pp. 600, Upper Saddle River, NJ. Liu, Z. & J. Zhao, 2000: Contribution of carbonate rock weathering to the atmospheric CO 2 sink.Environ mental Geology, 39, 1053. Liu, Z., Dreybrodt, W & H. Liu, 2011: Atmospheric CO 2 sink: Silicate weathering or carbonate weathering?Applied Geochemistry, 26, S292S294. Miotke, F.-D., 1974: Carbon dioxide and the soil atmo sphere.Abhandlungen zur Karst-und Hhlen kunde, A9, pp. 49. Moosdorf, N., Hartmann, J., Lauerwald, R., Haedorn, B. & S. Kempe, 2011: Atmospheric CO 2 consumption by chemical weathering in North America.Geo chimica et Cosmochimica Acta, 75, 7829. Shuster, E.T., 1970: Seasonal variations in carbonate spring water chemistry related to ground water ow.M.S. esis, e Pennsylvania State University, pp. 148. imnek, J. & D.L. Suarez, 1993: Modeling of carbon dioxide transport and production in soil 1. Model development.W ater Resources Research, 29, 487 497. Suarez, D.L. & J. imnek, 1993: Modeling of carbon di oxide transport and production in soil 2. Parameter selection, sensitivity analysis, and comparison of model predictions to eld data.W ater Resources Research, 29, 499. Troester, J.W ., 1994: e geochemistry, hydrogeology, and geomorphology of the Rio Camuy drainage basin, P uerto Rico: A humid tropical karst.Ph.D. esis, e Pennsylvania State University, pp. 344. Troester, J.W & W .B. W hite, 1984: Seasonal uctuations in the carbon dioxide partial pressure in a cave at mosphere.W ater Resources Research, 20, 153 156. Veni, G., 1997: Geomorphology, hydrogeology, geo chemistry, and evolution of the karstic Lower Glen Rose aquifer, south-central Texas.Texas Speleolog ical Survey Monographs 1, pp. 409. W hite, W .B., 1988: Geomorphology and H ydrology of Karst Terrains. Oxford University Press, pp. 464, New York. W hite, W .B., 1997: ermodynamic equilibrium, kinet ics, activation barriers, and reaction mechanisms for chemical reactions in karst terrains.Environ mental Geology, 30, 46. W ood, W .W ., 1985: Origin of caves and other solution openings in the unsaturated (vadose) zone of car bonate rocks: A model for CO 2 generation.Geol ogy, 13, 822. W orthington, S.R.H., Ford, D.C. & G.J. Davies, 2000: Matrix, fracture and channel components of storage and ow in a Paleozoic limestone aquifer. In: Sa sowsky, I.D. & C.M. W icks (eds.) Groundwater Flow and Contaminant Transport in Carbonate Aquifers A.A. Balkem, pp. 113-128, Rotterdam. WILLIAM B. WHITE



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D O CARBONATE KARST TERRAINS AFFECT THE GLOBAL CARBON CYCLE? A LI KRA KA OBMO JA NA KARBONATIH VPLIVAJO NA GLOBALNO KROENJE OGLJIKA? Jonathan B. MARTIN 1 Amy B ROW N 1 & John EZELL 1 Izvleek UDK 551.44:546.26 Jonathan B. Martin, Amy Brown & John Ezell: Ali kraka obmoja na karbonatih vplivajo na globalno kroenje oglji ka? Kljub temu, da so karbonati so najveje skladie ogljika v li tosferi, velja splona domneva, da nimajo pomembnega vpliva na globalno kroenje ogljika, ker sta raztapljanje in izloanje karbonatov uravnoteen izvor in ponor atmosferskega oglji ka. Kopenski in morski karbonati se v izbranem hidrolokem siste mu soasno raztapljajo in izloajo. V vseh primerih, ki jih obravnavamo, sta raztapljanje in izloanje povezana s primarno produkcijo, ki uskladii atmosferski CO 2 kot organski ogljik in prehajanje organskega ogljika v raztoljeni CO 2 ob remineralizaciji v krakih vodonosnikih. Odlaganje karbonatov v morju predstavlja prehod CO 2 v ozraje. Raztapljanje karbonatov je ponor CO 2 e raztaplja ogljikova kislina oziroma izvor CO 2 e raztapljanje poteka preko oksidacije sulda na meji med slano in sladko vodo. Ker je raztapljanje in izloanje odvisno od tevilnih okoljskih parametrov, lahko spremembe okolja, kot so intenziteta in pogostost padavin, temperatura povrja in spremembe morske gladine, spremenijo velikosti izvorov in ponorov atmosferskega CO 2 iz krakih procesov in vodijo v nove povratne zanke v kroenju ogljika, ki se razlikujejo od dananjih. Kljune besede: globalno kroenje ogljika, skladienje or ganskega ogljika, remineralizacija, raztapljanje karbonatov, izloanje karbonatov. 1 University of Florida, Department of Geological Sciences, PO Box 112120, 241 W illiamson Hall, Gainesville Florida 32605, 352, e-mail: jbmartin@u.edu, amy.brown@u.ed, ezelljohn1@gmail.com Received/Prejeto: 20.3.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 187, POSTOJNA 2013 Abstract UDC 551.44:546.26 Jonathan B. Martin, Amy Brown & John Ezell: Do carbonate karst terrains aect the global carbon cycle? Carbonate minerals comprise the largest reservoir of carbon in the earths lithosphere, but they are generally assumed to have no net impact on the global carbon cycle if rapid dissolution and precipitation reactions represent equal sources and sinks of atmospheric carbon. Observations of both terrestrial and marine carbonate systems indicate that carbonate minerals may simultaneously dissolve and precipitate within dierent portions of individual hydrologic systems. In all cases reported here, the dissolution and precipitation reactions are related to primary production, which xes atmospheric CO 2 as organic carbon, and the subsequent remineralization in watersheds of the organic carbon to dissolved CO 2 Deposition of carbonate minerals in the ocean represents a ux of CO 2 to the atmo sphere. e dissolution of oceanic carbonate minerals can act either as a sink for atmospheric CO 2 if dissolved by carbonic acid, or as a source of CO 2 if dissolved through sulde oxida tion at the freshwater-saltwater boundary. Since dissolution and precipitation of carbonate minerals depend on ecological processes, changes in these processes due to shis in rainfall patterns, earth surface temperatures, and sea level should also alter the potential magnitudes of sources and sinks for atmo spheric CO 2 from carbonate terrains, providing feedbacks to the global carbon cycle that dier from modern feedbacks. Keywords: Global carbon cycle, carbonate terrains, organic carbon xation, remineralization, carbonate mineral dissolu tion, carbonate mineral precipitation.

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ACTA CARSOLOGICA 42/2-3 2013 188 Models of the global carbon cycle attempt to identify the magnitude of carbon in all earth, ocean, and atmospheric reservoirs, quantify the uxes between those reservoirs, and assess chemical transformations of the carbon within reservoirs and in the ow paths between reservoirs (e.g., Solomon et al. 2007, Fig. 1). Estimates of carbon cycling are particularly important because the concentration of atmospheric CO 2 is thought to have approximately dou bled over the past few hundred years due to land use changes and increased uxes of carbon from the terres trial biosphere and burning of fossil fuels with industri alization (Solomon et al. 2007). Atmospheric CO 2 con centration has been directly measured over the past 67 years and these measurements show that the partial pres sure of CO 2 in the earths atmosphere has increased from around 318 ppmv to the most recent values of around 395 ppmv (Keeling & W horf 2005). is increase is sim ilar in magnitude to variations in CO 2 partial pressure that occurred throughout Pleistocene glacial-interglacial cycles as estimated from CO 2 concentrations measured from the Vostok ice core in Antarctica (Falkowski et al. 2000). Changing land use and burning of fossil fuels are estimated to have increased annual uxes of carbon to the atmosphere by approximately 8 petagrams (10 15 g) of C per year (PgC/yr), an increase of slightly more than 4% over the pre-industrial uxes of carbon (Solomon et al. 2007). is new ux of carbon, and the increase in atmospheric carbon abundance, as well as associated in creases in the oceans, are presumed to lead to a variety of environmental eects, including global warming (Mann et al. 1998), ocean acidication (Orr et al. 2005, HoeghGuldberg et al. 2007), and melting of land-based glaciers (Alley et al. 2005, Overpeck et al. 2006). Melting of landbased glaciers, along with thermal expansion, lead to ris ing sea level (Lambeck et al. 2002). e major pre-industrial uxes of carbon between reservoirs include exchange between the atmosphere and terrestrial biosphere, which was approximately 120 PgC/yr and between the atmosphere and surcial ocean, which was approximately 70 PgC/yr (Solomon et al. 2007). Of the reservoirs of carbon that typically are in cluded as participating in global carbon cycling, the ocean is the largest reservoir, containing approximately 3.8 x 10 4 Pg C, compared with the current abundance of carbon in the atmosphere of around 7.6 x 10 2 Pg C (Fig. 2). e other major reservoir of carbon that partici pates in carbon cycling is the terrestrial biosphere, which is estimated to contain around 2.3 x 10 3 Pg C. By far, the largest reservoir of carbon in the earth system is car bonate rocks, which contain as much as 4 x 10 7 Pg C, or more than three orders of magnitude more carbon than the oceanic reservoir (Fig. 2). ese carbonate rocks crop out over approximately 20% of the earths ice-free surface (e.g., Ford & W illiams 2007), but may be mantled by soils containing variable amount of organic carbon. Carbonate rocks undergo dis solution due to the action of CO 2 which when dissolved I NTRODUCTION JONATHAN B. M ARTIN A MY B RO W N & JOHN E ZELL Fig. 1: A generalized version of the modern global carbon cycle model. e numbers in the box are in petagrams (10 15 g) C (Pg C) and the uxes are PgC/yr and include uxes from anthropo genic sources. For simplicity, the model shows only uxes between the atmosphere and biosphere and oceans. e grey polygons within the carbonate rock reservoir represent karst features in cluding caves, conduits and blue holes. Modied from Solomon et al. (2007) and Kump et al. (2010). Fig. 2: H istogram of the magnitude of carbon stored in various global reservoirs showing carbonate terrains with the largest abundance of carbon (data from Falkowski et al. 2000).

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ACTA CARSOLOGICA 42/2-3 2013 189 A TMOSPHERIC CO 2 FLUXES IN SILICATE AND CARBONATE WEATHERING Silicate weathering reactions can be represented by the following general reaction mechanism: CaSiO 3(s) + 2CO 2(g) + H 2 O Ca 2+ + 2HCO 3 + SiO 2(s) (1) in which a Ca silicate mineral (here shown as wollas tonite) reacts with carbonic acid formed by the hydra tion of atmospheric CO 2 and water to generate dissolved calcium and bicarbonate, along with silica. e dissolved calcium and bicarbonate can react to form calcium car bonate (either aragonite or calcite), water, and CO 2 ac cording to Ca 2+ + 2HCO 3 CaCO 3(s) + H 2 O + CO 2(g) (2). e silica in reaction 1 and carbonate mineral in re action 2 are commonly precipitated by autotrophic and heterotrophic organisms in the oceans, and to a lesser degree in lakes and other terrestrial water systems. e resulting solid phases commonly sink to the bottom of the water bodies in which they have formed, thereby se questering the silica and carbon in a process referred to as the biological pump. e summation of reactions 1 and 2 yields the following reaction CaSiO 3 + CO 2 CaCO 3 + SiO 2 (3) which shows explicitly that for each mole of Ca silicate minerals dissolved, a mole of atmospheric CO 2 is seques tered through the biologic pump. e forward reaction, which is a low-temperature weathering reaction, tends to be faster than the reverse reaction, which occurs at metamorphic temperatures pressures. Consequently, the weathering of silicate minerals is typically considered an D O CARBONATE KARST TERRAINS AFFECT THE GLOBAL CARBON CYCLE ? and hydrated forms carbonic acid, the most common weathering agent for all rock types, including carbonate as well as silicate rocks. e CO 2 driving this dissolution can come from surface waters in equilibrium with atmo spheric CO 2 or from the remineralization of the organic carbon which may increase CO 2 concentrations over that of the atmosphere resulting in increased dissolution. ese dissolution reactions result in the transformation of atmospheric CO 2 into dissolved bicarbonate, thereby acting as a sink for atmospheric CO 2 Since carbonate minerals tend to be more soluble, with faster dissolution kinetics, than silicate minerals, they may be expected to play a larger role in dissolution reactions than silicate minerals and thus sequestration of atmospheric CO 2 Carbonate dissolution is neglected in most global car bon cycle models, however, because precipitation of car bonate minerals also liberates CO 2 resulting in no net change in the cycle (Berner et al. 1983). No net change in atmospheric CO 2 would occur only if both precipitation and dissolution occur at the same rates within the times cale of interest (e.g., Liu et al. 2010, 2011). Recent work in karst systems has suggested that at small spatial and short time scales, cycling of carbon ate minerals could be more important for drawdown of atmospheric carbon than silicate dissolution when carbon is buried as solid mineral and organic phases (e.g., Liu et al. 2010, 2011). For example, rapid precipi tation of marine carbonates at elevated sea level and their dissolution at low sea level during glacial times may, in part, control atmospheric CO 2 concentrations at glacial-interglacial timescales (Berger 1982; Mylroie 1993). e soluble nature of carbonate minerals can lead to large dissolution voids (speleogenesis and for mation of karstic features), reecting the exchange of carbon between the atmosphere and solid carbonate minerals. Speleogenesis also impacts other valuable re sources such as potable water and ecosystem services that rely on karst waters (e.g., Ford & W illiams 2007), reecting the importance of this process on human social systems. Consequently, understanding how dis solution and reprecipitation of carbonate minerals may be linked to the global carbon cycle is critical to evalu ations of impact to ecological function and water re sources during changes to the global climate. In the following paper, we present case studies that link reactions and exchange with atmospheric CO 2 in karst hydrologic processes with carbonate dissolution and precipitation. e case studies represent both terres trial and marine systems. In all cases, the reactions of the carbonate mineral phases are linked to ecological pro cesses, including photosynthetic xation of atmospheric CO 2 at the land surface and in subaquatic vegetation, and remineralization of the xed carbon. Case studies pre sented here are used to develop a general understanding of processes that provide sources or sinks of atmospheric CO 2 through reactions with solid carbonate phases. e case studies are somewhat restricted geographically and through time, which limits our ability to extract global magnitudes of the uxes from the processes that are identied as sources or sinks of atmospheric CO 2

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ACTA CARSOLOGICA 42/2-3 2013 190 important control in the sequestration of atmospheric CO 2 and has been suggested to link major tectonic events, such as the upli of the Himalayas, to long-term changes in atmospheric CO 2 partial pressures and global climate (Raymo et al. 1988). e dissolution of carbonate minerals by carbonic acid may be considered by the following reaction CaCO 3(s) + CO 2(g) + H 2 O Ca 2+ + 2HCO 3 (4) in which carbonic acid is formed by hydration of one mole of atmospheric CO 2 and dissolves carbonate min erals to generate dissolved calcium and bicarbonate. is reaction diers from silicate dissolution (e.g., reaction 3) because the reaction products remain dissolved so that the reverse of reaction 4 (e.g., reaction 2) releases one mole of atmospheric CO 2 Consequently, the net result of reactions 2 and 4 is that CO 2 is not permanently se questered from the atmosphere, and instead is returned to the atmosphere during precipitation of the calcium carbonate minerals. e reaction rates of the forward and reverse reactions are similar, compared with the large variations in forward and reverse reaction rates of reaction 3, and thus carbonate minerals are thought to have no net eect on the global carbon cycle (Berner et al. 1983). Dissolution and precipitation of marine car bonates is balanced at a global scale only over glacialinterglacial time scales because of dierences in the tim ing of their deposition and precipitation with sea level variations (Mylroie 1993). If the net dissolution and pre cipitation reactions dier locally, imbalances between reactions 2 and 4 could impact local carbon cycling (e.g., Liu et al. 2010, 2011). C ASE STUDIES Florida a terrestrial system. Limestone of the Floridan Aquifer, an Eocene aged eogenetic (Vacher & Mylroie 2002) karst aquifer, is exposed across a small window of north-central Florida, but elsewhere throughout the re gion, the aquifer is conned by a siliciclastic unit called the Hawthorn Group (Scott 1988, Scott 1992). e boundary between the conned and unconned Floridan Aquifer is locally referred to as the Cody Scarp and represents an erosional edge of the Hawthorn Group rocks. is boundary provides an important control on the hydrolo gy and hydrogeology of the region by separating terrains that are characterized by large amounts of surface runo from areas in which most surface water sinks into the Floridan Aquifer (Gulley et al. 2012). e composition of surface water draining o of the Hawthorn Group tends to be enriched in dissolved organic carbon. Oxidation of this dissolved organic carbon forms CO 2 which hydrates to carbonic acid and dissolves carbonate minerals (reac tion 2). Although this CO 2 is not directly dissolved from atmospheric CO 2 its ultimate source is atmospheric CO 2 that has been photosynthetically xed to organic carbon, and thus also remains a sink of atmospheric CO 2 e tendency for water to dissolve or precipitate mineral phases is commonly evaluated based on calcula tions of the saturation index, a measure of the saturation state of the water with respect to mineral phases. e saturation index is dened here as log(IAP/K sp ), where IAP is the ion activity product of solutes in the water that take part in mineral dissolution reactions, and the K sp is the thermodynamic equilibrium constant relative to the mineral phase of the reaction, in this case calcite (e.g., Stumm & Morgan 1996). W hen the IAP = K sp the sys tem is in equilibrium and the log(IAP/K sp ) equals zero. Negative values indicate undersaturation with respect to the particular mineral and a tendency for that mineral phase to dissolve and positive values indicate supersatu ration and a tendency for that mineral to precipitate. Along the Cody Scarp in north-central Florida, ow of organic carbon and CO 2 -rich water into the Flor idan Aquifer drives dissolution of the aquifer (Martin & Dean 2001). Although all of the CO 2 dissolved in the wa ter could ultimately react with carbonate minerals, the newly identify mechanism for calcite dissolution results from ooding surface water owing into spring vents that normally discharge to rivers during baseow condi tions (Gulley et al. 2011). ese events are common at springs along the Suwannee River, particularly springs in the unconned Floridan Aquifer immediately down stream from the Cody Scarp. One such event at the Peacock Spring system, lo cated approximately 40 km downstream of the scarp allowed water that was characterized by saturation in dices with respect to calcite of less than to ow into the Floridan Aquifer (Fig. 3). River water owed into the spring during the ood for approximately 3 weeks as shown by the presence of low conductivity river wa ter in the spring system, which has relatively constant specic conductivity value of approximately 400 S/cm during base ow. e rapid decrease of specic conduc tivity during the rising limb of the hydrograph reects JONATHAN B. M ARTIN A MY B RO W N & JOHN E ZELL

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ACTA CARSOLOGICA 42/2-3 2013 191 the initial intrusion of river water and the rapid return to elevated specic conductivity during the recession limb reects renewed discharge from the spring. Grab samples collected from the Suwannee River and two karst windows (Orange Grove Sink and Challenge Sink) connected to the Peacock Spring system indicate that the water owing into the subsurface remained undersatu rated with respect to calcite throughout the ood and recession curve although the saturation index increased through time to values of around (Fig. 3). Since the saturation index did not indicate a return to equilib rium during the recession, the water is unlikely to have dissolved enough calcite to reach equilibrium with the minerals making up the Floridan Aquifer. Nonetheless, the increase in the value of the saturation index, and its low initial value, suggests that some calcite was dissolved during this ood. Dissolution of this calcite would oc cur through reaction 4, sequestering atmospheric CO 2 as dissolved bicarbonate. is bicarbonate could ultimately be transported to the ocean, thereby impacting its alka linity. Numerous springs discharge from the unconned portion of the Floridan Aquifer in north-central Florida (Rosenau et al. 1977, Scott et al. 2004) and during times of discharge, the water typically has low dissolved or ganic carbon concentrations, elevated P CO2 values, and is at equilibrium with calcite within the Floridan Aquifer. e elevated P CO2 in the spring water typically degasses to the atmosphere, which can drive the water toward supersaturation with respect to carbonate minerals, in cluding calcite (M. Khadka, Pers. comm.). In addition, metabolism of in-stream submerged vegetation reduces the CO 2 concentrations during daytime photosynthesis (de Montety et al. 2011). is metabolism also drives cy cling of other solutes, in particular, nutrient concentra tions (e.g., Heernan & Cohen 2011, Cohen et al. 2012). ese changes in the chemical composition of the spring runs alter the saturation state of the water with respect to calcite at a 24 hour (i.e., diel) frequencies (Fig. 4). e saturation state always remains supersaturated, and tem poral variations in saturation states show an inverse re lationship to changes observed in specic conductivity, which are largely driven by changes in the Ca 2+ concen trations (de Montety et al. 2011). ese changes suggest that calcite precipitates in the spring runs as CO 2 is re moved through degassing to the atmosphere and uptake by the submerged vegetation. Furthermore, the precipi tated calcite is not balanced by subsequent dissolution in the river since the river never has a saturation index reecting undersaturation. ese studies from the terrestrial freshwater system of north-central Florida indicate that carbonate miner als (largely calcite) will both dissolve and precipitate, depending on the particular location within the system. W hether dissolution or precipitation is the primary pro cess depends largely on whether atmospheric and/or dissolved inorganic carbon is being xed during photo synthesis, or if dissolved organic carbon is being respired to CO 2 Regardless of the pathway, the ultimate source of carbon that precipitates as carbonate minerals is at mospheric, but precipitation provides a source of CO 2 Fig. 3: Time series plot of stage of Suwannee River near P eacock Springs (dashed line), variation in specic conductivity water at the entrance to P eacock Springs (solid line), and the saturation index of the Suwannee River water (gray diamonds) and water at two sink holes connected to P eacock Springs via conduits: Or ange Grove Sink (black dots) and Challenge Sink (white squares). Modied from Gulley et al. (2011). Fig. 4 : Time series measurements of the Ca 2+ concentration and the saturation index with respect to calcite in the Ichetucknee River, Florida between 16:00 hrs on March 26, 2009 and 16:00 hrs March 27, 2009 (Modied from de Montety et al. 2011). D O CARBONATE KARST TERRAINS AFFECT THE GLOBAL CARBON CYCLE ?

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ACTA CARSOLOGICA 42/2-3 2013 192 to the atmosphere (e.g., reaction 2), while dissolution of carbonate through the oxidation of organic carbon se questers carbon from the atmosphere (e.g., reaction 4). e relative magnitude of atmospheric carbon that is released or sequestered by these processes is unknown. Both sequestration and release of carbon to the atmo sphere occur simultaneously at dierent locations within a single hydrologic system, indicating that the magnitude of CO 2 lost from or gained by the atmosphere through dissolution and precipitation of carbonate minerals in terrestrial carbonate karst terrains could be balanced at small spatial and temporal scales. e north Florida car bonate platform is currently being denuded as shown by a net ux of bicarbonate to the oceans and geomorphic modeling (Opdyke et al. 1984, Adams et al. 2010), indi cating this system currently represents a long term sink for atmospheric CO 2 e Bahamas a marine system. e bicarbonate that is generated during mineral weathering (e.g., reac tions 1, 3 and 4) is carried to the oceans through river systems and groundwater. is bicarbonate is an impor tant control on alkalinity of the oceans along with cir culation through mid-ocean ridge crests and anks. e bicarbonate may precipitate as solid carbonate by microorganisms (e.g., foraminifera and cocolithophorids) and by calcifying algae such as H alimeda and P enicillus and to a smaller extent by reef-forming corals. e microorgan isms are deposited on the seaoor and buried, but much of the calcifying algae form deposits of large carbonate banks such as the Bahamas (Sealey 2006). Regardless of the ultimate depositional site for these carbonate miner als, for each mole of calcium carbonate deposited, one mole of CO 2 is released to the atmosphere (reaction 2), a process referred to by Berger (1982) as the coral reef hy pothesis. Carbonate banks are not solely sites of deposi tion, however; much dissolution occurs in the carbonate systems as shown by widespread cave formation, such as ank margin and vertical caves (Mylroie et al. 1995). At modern sea level, vertical caves are commonly lled with water, in which case they are referred to as blue holes in the Bahamas and cenotes in the Yucatan (Mylroie et al. 1995, Schwabe & Herbert 2004). Dissolution in carbon ate platforms, if resulting from carbonic acid generated through dissolution of atmospheric CO 2 or remineral ization of organic carbon that had as its ultimate source CO 2 xed from the atmosphere, should sequester atmo spheric carbon as bicarbonate in the oceans. Blue holes (and cenotes) are open to the surface and thus act as repositories for organic carbon generated within surface ecological systems. W here deep enough, these features may penetrate through the fresh-water lens of low-lying carbonate platforms. e contact between the fresh water and salt water underlying the fresh-water lens traps organic carbon because of density dierences in the two dierent salinity waters. Remineralization of this organic carbon can consume all the available oxygen at that horizon. e underlying saltwater typically has sulfate concentrations of seawater, which are elevated over those of the freshwater and can be used as a ter minal electron acceptor for organic carbon remineral ization following the consumption of oxygen (Stoessell et al. 1989). Sulfate reduction produces hydrogen sulde, which is re-oxidized to sulfate when diused or advected into portions of the water column containing oxygen, thereby generating sulfuric acid, reducing the saturation state of the water within the blue hole, and potentially driving dissolution of the carbonate walls of the blue hole by reaction mechanisms such as H 2 S + 2O 2 + CaCO 3 (s) SO 4 2 + Ca 2+ + H 2 O + CO 2 (5). is mechanism for cave formation was perhaps rst rec ognized in terrestrial systems at Carlsbad Caverns, New Mexico (e.g., Hill 1987, Hill 1990). Dissolution requires that the undersaturated water comes into contact with the wall rock of the blue hole and thus exchange must occur between water within the blue holes and the wall rock. In terrestrial systems, hydrologic gradients can be driven by topography such that during ooding, the hydraulic head within karst features (e.g., caves and conduits) may be greater than the hydraulic head with in the pore spaces of the aqui fer (Martin & Dean 2001). Alternatively, water may ow upward from geothermal heating, as in the case of Carls bad Cavern, where sulfur is introduced into the cave sys tem through migration of sulfur in oil eld brines in the neighboring Permian Basin (Hill 1987, Hill 1990). Topo graphic relief of carbonate platforms is rarely sucient to drive exchange between voids such as blue holes and the surrounding matrix porosity, although geothermal heating may drive vertical circulation in carbonate plat forms (Kohout 1965, W hitaker & Smart 1990, Martin & Moore 2008). Suciently large dierences in hydraulic head to drive small-scale (e.g., a few meters) exchange between blue holes and the surrounding matrix have been shown to result from tidal lags between water in the matrix porosity and water in blue holes that are connected to the ocean via conduits (W hitaker & Smart 1997, Mar tin et al. 2012). e tidal lags result from dierences in lower hydraulic conductivity in the matrix porosity than in the conduits. Exchange allows water undersaturated with respect to carbonate minerals from dissociation of either carbonic acid or sulfuric acid to contact and dis solve the minerals. W hether dissolution occurs from carbonic or sulfuric acid is critical to the cycling of at JONATHAN B. M ARTIN A MY B RO W N & JOHN E ZELL

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ACTA CARSOLOGICA 42/2-3 2013 193 mospheric carbon within these systems. Dissolution by carbonic acid acts to sequester CO 2 from the atmosphere (e.g., reaction 4) in carbonate platforms, counteracting the source of atmospheric CO 2 from precipitation of car bonate minerals (e.g., reaction 2). In contrast, carbonate mineral dissolution caused by sulfuric acid produced through the oxidation of sulde (reaction 5) provides CO 2 to the atmosphere that otherwise would have been sequestered as carbonate minerals via the biological pump. An example of how sulfur may play a role in disso lution of carbonate minerals is shown by variations in the specic conductivity and sulde concentrations through tidal cycles at the edge and in the center of Ink W ell Blue Hole on San Salvador Island Bahamas (Fig. 5). Ink W ell Blue Hole, which has a diameter at the land surface of about 20 m, has a tidal range similar to ocean tides with little lag, indicating it has a high-permeability connec tion with the ocean. e sulde and specic conductivity data shown in gure 5 were measured in water samples that were collected using tubing suspended from oats exactly 2 m below the water surface at three locations: about 1 m of the edge of the blue hole, in the center of the blue hole and midway between the edge and center of the blue hole. e change in specic conductivity though the tidal cycle suggests that the thickness of the freshwater lens varies through the tidal cycle, probably as the sides of the blue hole slope slightly outward resulting in a wider diameter of the blue hole at higher elevations. As water levels rise during the ooding tide, the water within the blue hole ows into a wider portion of the blue hole, thereby thinning the fresh-water lens, which retains an approximately constant volume, and increases the specic conductivity 2 m below the water surface. e dierence in the sulde concentrations from the edge to the center of the blue hole reects the exchange of water from the blue hole water column with sulderich water within the matrix at the edge of the blue hole. At low tide, the head in the aquifer is elevated above the water level in the blue hole, creating a hydrologic gradi ent from the aquifer to the blue hole (Martin et al. 2012). is gradient causes water to ow from the aquifer into the blue hole. Sulde is also generated within the water column because of sulfate reduction. Sulde generated in situ as well as provided from the aquifer would be oxidized by dissolved oxygen in the upper portion of the water column. Only water at the edge of the blue hole with reduced pH values from sulde oxidation is likely to exchange with matrix water and drive dissolution re actions. W ater in the center of the blue hole is unlikely to dissolve calcite of the aquifer because of its lack of con tact with the wall rock. Exchange of water with low pH between the wa ter column and the aquifer matrix derived from the oxidation of sulde will allow dissolution through re action mechanism similar to reaction 5. is reaction, although it dissolves carbonate, generates atmospheric CO 2 rather than sequestering atmospheric CO 2 such as occurs via a dissolution a reaction mechanism such as reaction 4. A source of CO 2 from sulde-driven disso lution could be important for the global carbon cycle, depending on the global magnitude of this process. If blue holes are the only location where this type of disso lution reaction occurs, the magnitude of CO 2 produced is likely to be small. Reducing conditions also occur at the fresh water-salt water interface within the aquifer, which will cause sulfate reduction and subsequent sul de oxidation (e.g., Bottrell et al. 1993). Sulde-driven dissolution could thus be more widespread than the limited footprint provided by blue holes. A broad sur vey of well water chemistry could provide an estimate of the impact of this process outside of areas aected by blue holes. A more thorough sampling of the changes in the redox state beneath carbonate islands is required to estimate the net ux of these processes and whether dis solution by sulfuric acid provides an important source of atmospheric CO 2 Fig. 5: Sulde concentrations versus specic conductivity at Ink Well Blue H ole, San Salvador Island, Bahamas. e data were measured on samples collected at a constant depth of 2 m below the water level through a tidal cycle in a transect from the center of the blue hole (gray triangles) to the edge (black dots) with a sample point intermediate between the other two sampling points (gray squares). Samples with the low specic conductivity and sulde concentrations were sampled at low tide and with high specic conductivity at high tide. e lines represent linear re gression of the data from each sample site, with the slope (m), r 2 and p values shown for each regression. D O CARBONATE KARST TERRAINS AFFECT THE GLOBAL CARBON CYCLE ?

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ACTA CARSOLOGICA 42/2-3 2013 194 REFERENCES Adams, P.N., Opdyke, N.D. & J. M. Jaeger, 2010: Isostatic upli driven by karstication and sea-level oscilla tion: Modeling landscape evolution in north Flori da.Geology 38, 531. Alley, R.B., Clark, P.U., Huybrechts, P. & I. Joughin, 2005: Ice-sheet and sea-level changes. Science 310, 456. Berger, W .H., 1982: Increase of carbon dioxide in the atmosphere during deglaciation: the coral reef hy pothesis.Naturwissenschaen 69, 87. Berner, R.A., Lasaga, A.C. & R.M. Garrels, 1983: e carbonate-silicate geochemical cycle and its eect on atmospheric carbon dioxide over the past 100 million years.Am. J. Sci 283, 641. e case studies described here reect both sources and sinks of atmospheric CO 2 in both marine and terrestrial carbonate karst systems. In addition, all of the dissolu tion and precipitation reactions involve either direct or indirect ecological control of the carbon concentrations through photosynthetic xation of inorganic carbon and remineralization of organic carbon. e dissolution and precipitation of the solid carbonate minerals are also closely linked to the hydrology and hydrogeology of the region. Interactions between undersaturated water and the carbonate minerals of the aquifers require variabil ity of head gradients allowing renewal of undersaturated water in contact with the minerals, or release of CO 2 to the atmosphere to increase the saturation state of the car bonate minerals and to drive their precipitation. Although these results indicate that carbonate min erals are linked to the global carbon cycle (e.g., Fig. 1), it is not clear if these reactions result in net changes in the uxes of carbon throughout the cycle and thus will im pact atmospheric concentrations. W ith changing global climates, however, the linkages between dissolution and precipitation within carbonate terrains may change. In the past, lower sea level would have exposed much of the surface of the platforms, which are currently submarine, to the atmosphere (Mylroie 1993). In these conditions, limited surface water and low freshwater lenses would be expected to alter ecological processes on the land surface, perhaps reducing primary productivity, concen tration of organic carbon and its remineralization, and thus dissolution of the carbonate minerals through the formation of carbonic and sulfuric acids. Changes in sea level will have an important impact on the hydrogeol ogy of low-lying carbonate platforms. Rising sea level will initially ood low-lying carbonate terrains and raise the water table, possibly increasing primary productivity, xation of atmospheric CO 2 and dissolution reactions as this organic carbon is remineralized. As sea level contin ues to rise, for example to elevations higher than modern sea level such as in the last interglacial (Lambeck et al. 2002), the low lying areas of what are now islands will be ooded with seawater and increase the deposition of ma rine carbonate. Consequently, even if the modern global carbon cycle has no net impact from the dissolution and precipitation of carbonate minerals, changes in ecologi cal systems resulting from climate change impacts such as sea level variations and shis in global precipitation patterns, may alter feedbacks between dissolution and precipitation within carbonate terrains, resulting in a net impact to the global carbon cycle. S UMMARY AND REMAINING QUESTIONS A CKNO WLEDGEMENTS anks to omas and Erin Rothfuss and sta at the Gerace Research Centre for their help and logistical sup port during eld work in the Bahamas. anks also to John Mylroie and Bill Jones for extremely helpful reviews of an early version of this manuscript. W e acknowledge nancial support for this work from the National Sci ence Foundation, grant numbers EAR and EAR, as well as Geological Society of America, grant number 9821. Acknowledgment is also made to the Donors of the American Chemical Society Petroleum Research Fund for partial support of this research. JONATHAN B. M ARTIN A MY B RO W N & JOHN E ZELL

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ACTA CARSOLOGICA 42/2-3 2013 195 Bottrell, S.H., Carew, J.L. & J.E. Mylroie, 1993: Inorganic and bacteriogenic origins for sulfate crusts in ank margin caves, San Salvador Island, Bahamas, P ro ceedings of the Sixth Symposium on the Geology of the Bahamas, B. W hite, San Salvador Island, Baha mas, Bahamian Field Station, Ltd. 6, 17. Cohen, M.J., Heernan, J.B., Albertin, A. & J.B. Mar tin, 2012: Inference of riverine nitrogen pro cessing from longitudinal and diel variation in dual nitrate isotopes.Jour. Geophys. Res. 117, doi:10.1029/2011JG001715. de Montety, V., Martin, J.B., Cohen, M.J. Foster, C.R. & M.J. Kurz, 2011: Inuence of diel biogeochemical cycles on carbonate equilibrium in a karst river.Chemical Geology 283, 31. Falkowski, P., Scholes, R.J., Boyle, E., Canadell, J., Can eld, D., Elser, J., Gruber, N., Hibbard, K., Halgberg, P. & S. Linder, 2000: e global carbon cycle: a test of our knowledge of earth as a system.science 290, 291. Ford, D. & P. W illiams, 2007: Karst H ydrogeology and Geomorphology, John W iley and Sons, pp. 562, W est Sussex, Gulley, J., Martin, J.B., Screaton, E.J. & P.J. Moore, 2011: River reversals into karst springs: A model for cave enlargement in eogenetic karst aquifers.Geologi cal Society of America Bulletin 123, 467. Gulley, J.D., Martin, J.B., Spellman, P., Moore, P.J. & E.J. Screaton, 2012: Inuence of partial connement and Holocene river formation on groundwater ow and dissolution in the Florida carbonate platform.Hydrological Processes. Heernan, J.B. & M.J. Cohen, 2011: Direct and indirect coupling of primary production and diel nitrate dy namics in a subtropical spring-fed river.Limnol ogy and Oceanography 55, 677. Hill, A.C., 1987: Geology of Carlsbad Cavern and other caves in the Guadalupe Mountains, New Mexico and Texas.Signature 142735, 36. Hill, C.A., 1990: Sulfuric Acid Speleogenesis of Carls bad Cavern and Its Relationship to Hydrocarbons, Delaware Basin, New Mexico and Texas (1).AAPG Bulletin 74, 1685. Hoegh-Guldberg, O., Mumby, P.J., Hooten, A.J., Steneck, R.S., Greeneld, P., Gomez, E., Harvell, C.D., Sale, P.F., Edwards, A.J. & K. Caldeira, 2007: Coral reefs under rapid climate change and ocean acidica tion.Science 318, 1737. Keeling, C.D. & T.P. W horf, 2005: Atmospheric CO 2 re cords from sites in the SIO air sampling network.Trends: A compendium of data on global change 2009. Kohout, F.A., 1965: A hypothesis concerning the cyclic ow of salt water related to geothermal heating in the Florida Aquifer.New York Academy of Science Transactions 28, 249. Lambeck, K., Esat, T.M. & E.K. Potter, 2002: Links be tween climate and sea levels for the past three mil lion years.Nature 419, 199. Liu, Z., Dreybrodt, W & H. Liu, 2011: Atmosheric CO 2 sink: silicate weatering or carbonate weathering.Applied Geochemistry 26, 5292. Liu, Z., Dreybrodt, W & H. W ang, 2010: A new direc tion in eective accounting for the atmospheric CO2 budget: Considering the combined action of carbonate dissolution, the global water cycle and photosynthetic uptake of DIC by aquatic organ isms.Earth-Science Reviews 99, 162. Mann, M.E., Bradley, R.S. & M.K. Hughes, 1998: Globalscale temperature patterns and climate forcing over the past six centuries.Nature 392, 779. Martin, J.B. & R.W Dean, 2001: Exchange of water be tween conduits and matrix in the Floridan Aquifer.Chemical Geology 179, 145. Martin, J.B., Gulley, J. & P. Spellman, 2012: Tidal pump ing of water between Bahamian blue holes, aquifers, and the ocean.Journal of Hydrology 416, 28. Martin, J.B. & P.J. Moore, 2008: Sr concentrations and isotope ratios as tracers of ground-water circulation in carbonate platforms: Examples from San Salva dor Island and Long Island, Bahamas.Chemical Geology 249, 52. Mylroie, J.E., 1993: Carbonate deposition/dissolution cycles and carbon dioxide ux in the Pleistocene.In: B. W hite (eds), Proceedings of the Sixth Sym posium on the Geology of the Bahamas, Bahamian Field Station, San Salvador, Bahamas, Bahamian Field Station, Ltd. Mylroie, J.E., Carew, J.L. & A.I. Moore, 1995: Blue Holes: Denition and genesis.Carbonates and Evaporites 10, 225. Mylroie, J.E., Carew, J.L. & H.L. Vacher, 1995: Karst de velopment in the Bahamas and Bermuda.In: H. A. Curran and B. W hite (eds), Terrestrial and Shallow Marine Geology of the Bahamas and Bermuda. Geo logical Society of America. 300: 251, Boulder, Colorado. Opdyke, N.D., Spangler, D.P., Smith, D.L., Jones, D.S. & R.C. Lindquist, 1984: Origin of the epeirogenic up li of Pliocene-Pleistocene beach ridges in Florida and development of the Florida karst.Geology 12, 226. D O CARBONATE KARST TERRAINS AFFECT THE GLOBAL CARBON CYCLE ?

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ACTA CARSOLOGICA 42/2-3 2013 196 Orr, J. C., V. J. Fabry, O. Aumont, L. Bopp, S. C. Doney, R. A. Feely, A. Gnanadesikan, N. Gruber, A. Ishida & F. Joos, 2005: Anthropogenic ocean acidication over the twenty-rst century and its impact on cal cifying organisms.Nature 437, 681. Overpeck, J. T., B. L. Otto-Bliesner, G. H. Miller, D. R. Muhs, R. B. Alley & J. T. Kiehl, 2006: Paleoclimatic evidence for future ice-sheet instability and rapid sea-level rise.Science 311, 1747. Raymo, M. E., W F. Ruddiman & P. N. Froelich, 1988: Inuence of late Cenozoic mountain building on ocean geochemical cycles.Geology 16, 649. Rosenau, J. C., G. L. Faulkner, J. Charles W Hendry & R. W Hull, 1977: e Springs of Florida, Department of Natural Resources, Bulletin #31, pp. 461, Talla hassee, Florida, Schwabe, S. & R. A. Herbert, 2004: Black Holes of the Bahamas: what they are and why they are black.Q uaternary International 121, 3. Scott, T. M., 1988: e lithostratigraphy of the H awthorn Group (Miocene of Florida), Florida Geological Sur vey Bulletin, pp. 147, Scott, T. M., 1992: A Geological Overview of Florida, Florida Geological Survey, pp. 78, Scott, T. M., G. H. Means, R. P. Meegan, R. C. Means, S. B. Upchurch, R. E. Copeland, J. Jones, T. Roberts & A. W illet, 2004: Springs of Florida, Tallahassee, Florida, Florida Geological Survey. 66, 377. Sealey, N. E., 2006: Bahamian landscapes; an introduction to the physical geography of the Bahamas, Macmillan Publishers, pp. 174, Oxford, Solomon, S., D. Qin, M. Manning, Z. Chen, M. Marquis, K. B. Averyt, M. Tignor & H. L. Miller, 2007: IPCC.Intergovernmental Panel on Climate Change. Cli mate Change. Stoessell, R. K., W C. W ard, B. H. Ford & J. D. Schuert, 1989: W ater chemistry and CaCO3 dissolution in the saline part of an open-ow mixing zone, coastal Yucatan Peninsula, Mexico:.Geological Society of America Bulletin 101, 159. Stumm, W & J. J. Morgan, 1996: Aquatic Chemistry, John W iley and Sons, Inc., pp. 1022, New York, Vacher, H. L. & J. E. Mylroie, 2002: Eogenetic karst from the perspective of an equivalent porous medium.Carbonates and Evaporites 17, 182. W hitaker, F. F. & P. L. Smart, 1990: Active circulation of saline ground waters in carbonate platforms: Evi dence from the Great Bahama Bank.Geology 18, 200. W hitaker, F. F. & P. L. Smart, 1997: Groundwater circula tion and geochemistry of a karstied bank-margin al fracture system, South Andros Island, Bahamas.Journal of Hydrology 197, 293. JONATHAN B. M ARTIN A MY B RO W N & JOHN E ZELL



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QUATERNARY GLACIAL CYCLES : K ARST PROCESSES AND THE GLOBAL CO 2 BUDGET KVARTARNE POLEDENITVE: KRA KI PROCESI IN GLOBALNA BILANCA CO 2 Erik B. L ARSON 1 & John E. MYLROIE 2 Izvleek UDK 551.435.8:546.264-31 546.264-31:551.79 Erik B. Larson & John E. Mylroie : Kvartarne poledenitve: kraki procesi in globalna bilanca CO 2 O pomenu krakih procesov v globalni bilanci ogljika je bilo narejenih veliko tevilo raziskav, le malo pa jih govori o pov ezavi med poledenitvami ter z njimi povezanimi spremembami morske gladine in izpostavljenost krakega povrja zakrasevanju. Med poledenitvami je zaradi padca morske gladine veina karbonatnih platform izpostavljenih meteornemu zakrase vanju. Preko raztapljanja se CO 2 iz zraka pretvarja v bikarbonat v raztopini. Karbonatne platforme so grajene preteno it ara gonita, ki je bolj topen od kalcita, kar e dodatno prispeva k omenjeni pretvorbi. Po drugi strani pa so med poledenitvami karbonati v arktinih in v nekaterih zmernih klimatskih paso vih pokriti z ledom in na ta nain izkljueni iz procesov zakra sevanja. Po naih ocenah se uinek obeh mehanizmov iznii v okviru nekaj tisoink gigatone ogljika letno, kar je manj kot 1 % celotnega atmosferskega CO 2 ki ga sicer odstranijo kraki pro cesi. Na letni bazi je torej koliina atmosferskega ogljika, ki ga odstranijo kraki procesi danes, enaka kot v obdobju poledenitvenega vika. Po drugi strani pa je veji dele kvartarja pri padal poledenitvam. Raziskava ima pomen za razumevanje globalnega ravnoteja ogljika v kvartarju. Kljune besede: Kras, globalna bilanca ogljika, kvartar, zadnji poldedenitveni maximum, raztapljanje karbonatov. 1 Department of Geosciences, Mississippi State University, Mississippi State, MS 39762, e-mail: ebl47@msstate.edu, 2 Department of Geosciences, Mississippi State University, Mississippi State, MS 39762, e-mail: mylroie@geosci.msstate.edu, Received/Prejeto: 19.1.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 197 POSTOJNA 2013 Abstract UDC 551.435.8:546.264-31 546.264-31:551.79 Erik B. Larson & John E. Mylroie : Quaternary glacial cycles: Karst processes and the global CO 2 budget Extensive research has been conducted investigating the re lationship between karst processes, carbonate deposition and the global carbon cycle. However, little work has been done looking into the relationship between glaciations, subsequent sea level changes, and aerially exposed land masses in rela tion to karstic processes and the global carbon budget. During glaciations sea-level exposed the worlds carbonate platforms. W ith the sub-aerial exposure of the platforms, karst processes can occur, and the dissolution of carbonate material can com mence, resulting in the drawdown of CO 2 from the atmosphere as HCO 3 Furthermore, the material on the platform surfaces is primarily aragonite which is more readily soluble than calcite allowing karst processes to occur more quickly. During glacia tions arctic carbonates and some of the temperate carbonates are blanketed in ice, eectively removing those areas from karst processes. Given the higher solubility of aragonite, and the ex tent of carbonate platforms exposed during glaciations, this dissolution balances the CO 2 no longer taken up by karst pro cesses at higher latitudes that were covered during the last gla cial maximum e balance is within 0.001 GtC / yr, using soil p CO 2 (0.005 GtC / yr assuming atmospheric p CO 2 ) which is a dierence of <1% of the total amount of atmospheric CO 2 re moved in a year by karst processes. Denudation was calculated using the maximum potential dissolution formulas of Gombert (2002). On a year to year basis the net amount of atmospheric carbon removed through karstic processes is equivalent be tween the last glacial maximum and the present day, however, the earth has spent more time in a glacial conguration during the Q uaternary, which suggests that there is a net drawdown of atmospheric carbon during glaciations from karst processes, which may serve as a feedback to prolong glacial episodes. is research has signicance for understanding the global carbon budget during the Q uaternary. Keywords: Karst, Global Carbon Budget, Q uaternary, Last Glacial Maximum, Carbonate Dissolution.

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ACTA CARSOLOGICA 42/2-3 2013 198 Karst processes have long been known to result in the drawdown of atmospheric CO 2 (e.g. Ford & W illiams 2007; W hite 1988). e process results in the sequestra tion of one molecule of CO 2 for each molecule of calcium carbonate dissolved (equation 1). CaCO 3 + CO 2 + H 2 O Ca 2+ + 2HCO 3 (Equation 1) is process has been proposed to result in signi cant removal of atmospheric CO 2 (e.g. Gombert 2002; Liu et al. 2011; Mylroie 1993; 2008). Gombert (2002) proposed that on the global scale karst denudation re sulted in the removal of 0.3 Gt of C per year from the atmosphere. is 0.3 Gt C per year represents 21% of the unknown carbon sink in the continental biosphere suggesting that karst processes play a signicant role in balancing the global carbon budget (Gombert 2002; Liu et al. 2011; Schimel et al. 1996). e goal of this study is to determine the amount of carbon that is removed from the atmosphere during karst dissolution on a global scale for the present day and during the last glacial maximum (LGM) and to deter mine if the amount of carbon withdrawn during these two time periods balances out. e only other study to examine this issue looked only at a local scale in the Ba hamas (Mylroie 1993). Mylroie (1993) discovered that the sequestration of carbon during the LGM by carbon ate dissolution and release of carbon during the present day from carbonate deposition resulted in a net balance when viewed over the average length of glacial and inter glacial cycles. e material presented below represents the con tinuation of Mylroies (1993) eorts applied to a global scale. I NTRODUCTION E RIK B. L ARSON & JOHN E. M YLROIE METHODS Karst denudation was modeled using Gomberts (2002) maximal potential denudation model that is based on W hites (1984; 1988) maximum dissolution model. Go mberts (2002) model is a climatic model which requires as inputs only eective precipitation, p CO 2 temperature to correct the equilibrium constants and karstic area. Go mberts (2002) model was applied to the dierent Kppen climate zones in an eort to model global karst denuda tion more accurately. e Kppen climatic zones used were: polar, cold, cold temperate, warm temperate, Mediterranean, desert, subtropical, wet tropical and equatorial as suggested by Gombert (2002). Eective precipitation in the modern day was cal culated using the precipitation data of Miller (1949) and evapotranspiration data from Mu et al. (2011). Eective precipitation during the LGM was calculated using rain fall data from Miller (1949), which was corrected for the LGM as suggested by Clark et al. (1999), and evapotrans piration data were derived from Bush & Philander (1999). p CO 2 data were calculated using Brooks (1983) model for soil p CO 2 using eective precipitation. Atmo spheric p CO 2 was also used to provide an end member condition that would be similar to no soil cover at all (Palmer 2007). Equilibrium constants for calcite and aragonite in the present day were corrected for temperature using Millers (1949) temperature data. e equilibrium con stants used came from Plummer and Busenberg (1982). e equilibrium constants used during the LGM were modied using Clark et al.s (2009) temperature correc tion. Karstic area was calculated using Ford and W illiams (2007) global karst map; the Kppen climatic zones were overlaid on this resulting in karst areas for the dierent climate zones. During the glaciations when sea-level dropped, the carbonate platforms and coral reefs around the world would have become sub-aerially exposed re sulting in an increase in karstic areas of 1.25 million km 2 (Smith 1978). ese newly gained areas were then placed into their respective climatic zones. ese new areas were treated to be aragonite instead of calcite (as all other ar eas were treated) as the sediments in the coral reefs and on the carbonate platforms are predominately aragonite (Reijmer et al. 2009). As a result of glaciations several high latitude and high altitude karst areas were covered with glaciers ef fectively removing them from karstic processes. A total of 4.7 million km 2 of karst area was removed from the global karst areas and each subset was then subtracted from their respective climatic zones (Clark & Mix 2002; Velichko et al. 1997). W ith all these parameters now dened, Gomberts (2002) maximal potential dissolution model was applied

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ACTA CARSOLOGICA 42/2-3 2013 199 RESULTS Denudation rates in the dierent climatic zones were determined to vary from 10 mm per thousand years, up to 100 mm per thousand years depending on the p CO 2 and climatic conditions. e extrapolation of these denudation rates to volume of carbon drawn from the atmosphere, assuming soil p CO 2 and all calcite mineral ogy, can be found in Tab. 1. e volume of carbon re moved from the atmosphere assuming soil pCO 2 and both calcite and aragonite mineralogy can be found in Tab. 2. e volume of carbon taken out of the atmos phere assuming atmospheric p CO 2 and only calcite can be found in Tab. 3. Finally, the volume of carbon drawn from the atmosphere assuming atmospheric p CO 2 and calcite and aragonite mineralogy can be found in Tab. 4. A summary table of all the dierent volumes of carbon removed can be found in Tab. 5. In summary about 0.21 Gt C per year are with drawn under all the various karst dissolution models, for both the present day and the last glacial maximum. In QUATERNARY GLACIAL CYCLES : K ARST PROCESSES AND THE GLOBAL CO 2 BUDGET to the dierent climatic zones. e model was run us ing either atmospheric or soil p CO 2 and either assum ing all the carbonate material was calcite or that some of the newly exposed carbonate areas during the LGM were aragonite, resulting in four model runs. ese models and the assumptions within them can cause signicant spread in the nal denudation rates. W hite (2007) re ports that karst denudation rates in the present can vary over an order of magnitude within a geographic area. However, when viewed over a large scale these deviations should cancel out, and as long as there is consistency in the calculations these errors should cancel out between the present and the LGM resulting in a net dierence that is precise. Tab. 1: e calculated maximal potential dissolution of karst pro cesses around the world. ese data were created using soil pCO 2 and assuming all the CaCO 3 was calcite. Present Day LGM Dierence Polar 2.80E+10 7.86E+09 2.01E+10 Cold 5.58E+10 5.33E+10 2.49E+09 Cold Temperate 2.42E+11 1.98E+11 4.45E+10 Warm Temperate 1.26E+11 1.36E+11 .65E+09 Mediterranean 3.46E+10 3.79E+10 .26E+09 Desertic 1.31E+09 2.57E+09 .26E+09 Subtropical 1.39E+10 2.35E+10 .58E+09 Wet Tropical 1.73E+11 2.04E+11 .14E+10 Equitorial 6.43E+10 6.60E+10 .72E+09 Total (m 3 / 1000yr) 7.40E+11 7.29E+11 1.03E+10 Total (GtC / yr) 0.222 0.219 0.003 Tab. 2: e calculated maximal potential dissolution of karst pro cesses around the world. ese data were created using soil pCO 2 and assuming that the currently exposed CaCO 3 is calcite, but the coral reefs and carbonate platforms are covered with arago nite sediment. Present Day LGM Dierence Polar 2.80E+10 7.86E+09 2.01E+10 Cold 5.58E+10 5.33E+10 2.49E+09 Cold Temperate 2.42E+11 1.98E+11 4.45E+10 Warm Temperate 1.26E+11 1.36E+11 .65E+09 Mediterranean 3.46E+10 3.79E+10 .28E+09 Desertic 1.31E+09 2.57E+09 .26E+09 Subtropical 1.39E+10 2.38E+10 .90E+09 Wet Tropical 1.73E+11 2.19E+11 .62E+10 Equitorial 6.43E+10 6.60E+10 .72E+09 Total (m 3 / 1000yr) 7.40E+11 7.44E+11 .82E+09 Total (GtC / yr) 0.222 0.223 .001 Tab. 3: e calculated maximal potential dissolution of karst processes around the world. ese data were created using atmo spheric pCO 2 and assuming all the CaCO 3 was calcite. Present Day LGM Dierence Polar 4.43E+10 1.24E+10 3.18E+10 Cold 7.25E+10 6.93E+10 3.23E+09 Cold Temperate 2.19E+11 1.78E+11 4.02E+10 Warm Temperate 1.39E+11 1.50E+11 .06E+10 Mediterranean 4.13E+10 4.51E+10 .88E+09 Desertic 2.32E+09 4.80E+09 .48E+09 Subtropical 8.98E+09 1.55E+10 .57E+09 Wet Tropical 1.28E+11 1.60E+11 .20E+10 Equitorial 3.55E+10 3.56E+10 .92E+07 Total (m 3 / 1000yr) 6.91E+11 6.71E+11 1.96E+10 Total (GtC / yr) 0.207 0.201 0.006

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ACTA CARSOLOGICA 42/2-3 2013 200 D ISCUSSION: e total 0.21 Gt C per year that are removed from the atmosphere by karst processes in the modern day cor relates well within the 0.11 0.47 Gt C per year that have been calculated by previous workers (e.g. Gomb ert 2002; Liu et al. 2011). Given that the current data matches well with previous work, some of which used dierent methods, provides condence for the rest of the data and the conclusions regarding the LGM. e amount of carbon removed from the atmosphere may not actually be sequestered; the carbon as CO 2 could degas from surface water bodies back to the atmos phere, or be deposited as tufas (Equation 1) before even reaching the oceans where the carbon could be depos ited as carbonates. T HE L AST G LACIAL M A X IMUM e amount of carbon withdrawn from the atmosphere during the present day and during the LGM is essentially equal, at about 0.21 Gt C per year (Tab. 5). During the LGM sea-level was about 125 meters lower than today resulting in the sub-aerial exposure of coral reefs and carbonate platforms, while covering up high latitude and high altitude areas with glaciers. Some carbonate plat forms, such as the Bahamas, become exposed with a sealevel fall of only 20 m; almost all carbonate platforms are exposed by a sea-level fall of 60 m. Given that the amount of carbon removed from the atmosphere through karst processes balances between these two times with signi cant geographic changes implies that land area uctua tions, in conjunction with modied climate are sucient to make up for the glacial impacts. Taking into account the mineralogy changes helps to account for some of the dierence between the present day and the LGM, but it has relatively insignicant control in the total carbon vol ume calculations (Tab. 5). During the cycling from glacial to non-glacial peri ods there would be regional variation in the amount of karst denudation (and subsequent draw down of atmo spheric carbon) as W hite (2007) demonstrated, occurs during the present day. is variation is in part due to the heterogeneous nature of the rocks, the pCO 2 and precipitation conditions. However, if the same assump tions are made in the calculations for karst denudation in the present and during the LGM these errors should eectively cancel out, and the net dierence (or lack of dierence in the case of this study) becomes signicant in demonstrating the net amount of carbon removed from the atmosphere during the present and LGM are equivalent. During the Q uaternary, glacial periods have lasted about 10 times longer than the interglacial periods (e.g. Mylroie 1993). is is signicant because when looked at over the entire Q uaternary it becomes obvious that there would be a net removal of atmospheric carbon due to dissolution processes occurring from carbonate platform exposure. is may result in a feedback mechanism re Tab. 5: e summary data from tables 1. GtC / yr Scenario Present Day LGM Dierence Soil pCO 2 Calcite 0.222 0.219 0.003 Soil pCO 2 Calcite & Aragonite 0.223 .001 Atmospheric pCO 2 Calcite 0.207 0.201 0.006 Atmospheric pCO 2 Calcite & Aragonite 0.202 0.005 most cases the amount of carbon withdrawn from the at mosphere during the LGM is slightly less than that in the present day, but the dierence is less than 1%. Tab. 4: e calculated maximal potential dissolution of karst processes around the world. ese data were created using atmo spheric pCO 2 and assuming that the currently exposed CaCO 3 is calcite, but the coral reefs and carbonate platforms are covered with aragonite sediment. Present Day LGM Dierence Polar 4.43E+10 1.24E+10 3.18E+10 Cold 7.25E+10 6.93E+10 3.23E+09 Cold Temperate 2.19E+11 1.78E+11 4.02E+10 Warm Temperate 1.39E+11 1.50E+11 .06E+10 Mediterranean 4.13E+10 4.52E+10 .91E+09 Desertic 2.32E+09 4.80E+09 .48E+09 Subtropical 8.98E+09 1.58E+10 .78E+09 Wet Tropical 1.28E+11 1.63E+11 .44E+10 Equitorial 3.55E+10 3.56E+10 .92E+07 Total (m 3 / 1000yr) 6.91E+11 6.74E+11 1.70E+10 Total (GtC / yr) 0.207 0.202 0.005 E RIK B. L ARSON & JOHN E. M YLROIE

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ACTA CARSOLOGICA 42/2-3 2013 201 sulting in the prolonging of glacial periods, as proposed by Mylroie (1993). roughout this research the deposition of carbon ates has been ignored from the net balance, but now it will be addressed. During LGM the sea-level drop would cause the carbonate factory to become sub-aerially ex posed; the shutdown of the carbonate factory would prevent the deposition of signicant carbonate sedi ments and prevent the release of that CO 2 into the at mosphere (Equation 1). At the present day the carbon ate factory is actively releasing CO 2 into the atmosphere, and based on this qualitative statement there should be a net sequestration of carbon during glacial periods and a net release of carbon during interglacial periods caused by karst processes. For the Bahama platform, Mylroie (1993; 2008) demonstrated that the carbon ux through a glacial-interglacial cycle was equal. e rapid release of CO 2 by carbonate deposition during a ~ 10ka interglacial was compensated almost exactly by the ten times slower CO 2 sequestration during the ~100 ka of glacioeustatic platform exposure. is loss of atmospheric CO 2 due to global processes (e.g. primary productivity) including karstic processes during the Q uaternary is documented in the Vostok ice core by higher atmospheric CO 2 dur ing interglacial periods and lower atmospheric CO 2 dur ing glacial periods (e.g. Falkowski et al. 2000; Petit et al. 1998). C ONCLUSIONS Given the maximal potential denudation formula of Go mbert (2002) the global carbon budget was calculated with respect to karst dissolution for both the present day and the LGM. Karstic processes result in the removal of atmospheric carbon through dissolution (equation 1). Both the present day and LGM were found to remove 0.21 Gt C per year from the atmosphere, within 1% of each other regardless of the model used. is represents 16% of the unknown carbon sink of the continental bio sphere (Schimel et al. 1996). Furthermore, this research demonstrates that through land area changes, climatic changes and mineralogy dierences the amount of car bon removed through karstic dissolution processes in the present is equivalent to the amount removed from the atmosphere during the LGM. W hat is dierent, given the steep-sided nature of most carbonate platforms, is that carbonate deposition is greater during an interglacial than during a glacial cycle, perhaps compensated for by the longer duration of glacial cycles compared to inter glacial conditions. Furthermore, the yearly balance be tween todays atmospheric carbon removal by karst proc esses and the LGMs carbon drawdown by karst processes indicates the importance of climate in karst dissolution. e loss of 4.7 million km 2 of karst area in high latitude and high altitude regions to ice cover during a glaciation is exactly compensated by the gain of 1.25 million km 2 of karst area in tropical and subtropical regions due to sea level fall. Finally, this research is signicant as it is the rst to quantify the eect the Q uaternary glaciations had on karstic processes with respect to atmospheric carbon removal. A CKNO WLEDGMENTS e authors wish to thank W ill W hite and an anonymous reviewer for comments that improved the paper. Discus sion with Philippe Gombert was useful in developing the project. Finally, the Department of Geosciences at Mis sissippi State University is thanked for nancial support. QUATERNARY GLACIAL CYCLES : K ARST PROCESSES AND THE GLOBAL CO 2 BUDGET

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ACTA CARSOLOGICA 42/2-3 2013 202 Brooks, G.A., Folko, M.E. & E.O. Box, 1983: A world model of soil carbon dioxide.Earth Surface Pro cesses and Landforms, 8, 79. Bush, A.B.G. & S.G.H. Philander, 1999: e climate of the last glacial maximum: results from a coupled atmo sphere-ocean general circulation model.Journal of Geophysical Research, 104, D20, 24509. Clark, P.U. & A.C. Mix, 2002: Ice sheets and sea level of the last glacial maximum.Q uaternary Science Re views, 21, 1. Clark, P.U., Alley, R.B. & D. Pollard, 1999: Northern hemisphere ice-sheet inuences on global climate change.Science, 286, 1104. Clark, P.U., Dyke, A.S., Shakun, J.D., Carlson, A.E., Clark, J., W ohlfarth, B., Mitrovica, J.X., Hostetler, S.W & A.M. McCabe, 2009: e last glacial maxi mum.Nature, 325, 710. Falkowski, P., Scholes, R.J., Boyle, E., Canadell, J., Can eld, D., Elser, J., Gruber, N., Hibbard, K., Hgberg, P., Linder, S., Mackenzie, F.T., Moore III, B., Peder sen, T., Rosenthal, Y., Seitzinger, S., Smetacek, V. & W Steen, 2000: e global carbon cycle: a test of our knowledge of Earth as a system.Science, 290, 291. Ford, D. & P. W illiams, 2007: Karst hydrogeology and geo morphology.W iley, pp. 562, W est Sussex. Gombert, P., 2002: Role of karstic dissolution in global carbon cycle.Global and Planetary Change, 33, 177. Liu, Z., Dreybrodt, W & H. Liu, 2011: Atmospheric CO 2 sink: silicate weathering or carbonate weathering?.Applied Geochemistry, 26, S292S294. Miller, A.A., 1949: Climatology .Methuen and Company, pp. 325, London. Mu, Q ., Zhao, M. & S.W Running, 2011: Improvements to a MODIS global terrestrial evapotranspiration algorithm.Remote Sensing of Environment, 115, 1781. Mylroie, J.E., 1993, Carbonate deposition/dissolution cy cles and carbon dioxide ux in the Pleistocene.In: W hite, B. (ed.) P roceedings of the sixth symposium of the geology of the Bahamas, 11 th th June 1992, San Salvador, Bahamas, Bahamian Field Station, 103, Port Charlotte. Mylroie, J.E., 2008: Late Q uaternary sea-level position: evidence from Bahamian carbonate deposition and dissolution cycles.Q uaternary International, 183, 61. Palmer, A.N., 2007: Cave geology.Cave Books, pp. 454, Dayton. Petit, J.R., Jouzel, J., Raynaud, D., Barkov, N.I., Barnola, J.M., Basile, I., Bender, M., Chappellaz, J., Davis, M., Delaygue, G., Delmotte, M., Kotlyakov, V.M., Legrand, M., Lipenkov, V.Y., Lorius, C., Ppin, L., Ritz, C., Saltzman, E. & M. Stienenard, 1998: Cli mate and atmospheric history of the past 420,000 years from the Vostok ice core, Antarctica.Nature, 399, 429. Plummer, L.N. & E. Busenberg, 1982: e solubilities of calcite, aragonite and vaterite in CO 2 H 2 O solutions between 0 and 90C, and an evaluation of the aque ous model for the system CaCO 3 CO 2 H 2 O.Geo chimica et Cosmochimica Acta, 46, 1011. Reijmer, J.J.G., Swart, P.K., Bauch, T., Otto, R., Reuning, L., Roth, S. & S. Zechel, 2009: A re-evaluation of fa cies on Great Bahama Bank I: new facies maps of western Great Bahama Bank.In: Swart, P.K., et al. (eds.) P erspectives in carbonate geology: a tribute to the career of Robert Nathan Ginsburg. Blackwell, pp. 29, W est Sussex. Schimel, D., Alves, D., Enting, I., Heimann, M., Joos, F., Raynaud, D., W igley, T., Prather, M., Derwent, R., Ehhalt, D., Fraser, P., Sanhueza, E., Zhou, X., Jonas, P., Charlson, R., Rodhe, H., Sadasivan, S., Shine, K.P., Fouquart, Y., Ramaswamy, V., Solomon, S., Srinivasan, J., Albritton, D., Derwent, R., Isaksen, I., Lal, M. & D. W uebbles, 1996: Radiative forcing of climate change.In: Houghton, J.T., et al. (eds.), Cli mate Change 1995. e Science of Climate Change. Contribution of Working Group I to the Second As sessment Report of the Intergovernmental P anel on Climate Change. Cambridge University Press, pp. 69, Cambridge. Smith, S.V., 1978: Coral reef area and the contributions of reefs to processes and resources of the worlds oceans.Nature, 273, 225. Velichko, A.A., Kononov, Y.M. & M.A. Faustova, 1997: e last glaciation of earth: size and volume of icesheets.Q uaternary International, 41/42, 43. W hite, W .B., 1984: Rate processes: chemical kinetics and karst landform development.In: LaFleur, R.G. (ed.) Groundwater as a geomorphic agent. AllenUnwin, pp. 227, Boston. W hite, W .B., 1988: Geomorphology and hydrology of karst terrains.Oxford University Press, pp. 464, New York. W hite, W .B., 2007: Evolution and age relations of karst landscapses.Acta Carsologica, 36, 45. REFERENCES E RIK B. L ARSON & JOHN E. M YLROIE



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A FRAME W ORK FOR ASSESSING THE ROLE OF KARST CONDUIT MORPHOLOGY HYDROLOGY AND EVOLUTION IN THE TRANSPORT AND STORAGE OF CARBON AND ASSOCIATED SEDIMENTS OSNOVA ZA OCENJEVANJE VLOGE MORFOLOGIJE, HIDROLOGIJE IN RAZVOJA KRA KIH KANALOV PRI PRENOSU IN SKLADIENJU OGLJIKA IN SORODNIH SEDIMENTOV George VENI 1 Izvleek UDK 551.444:546.26 George Veni: Osnova za ocenjevanje vloge morfologije, hidrologije in razvoja krakih kanalov pri prenosu in skla dienju ogljika in sorodnih sedimentov Kraki vodonosniki in kanali nastajajo z raztapljanjem karbonatnih mineralov in poasnim sproanjem anorganskega ogljika. Z razvojem morfologije, velikosti in poloajem znotraj vodonosnika, se njihova funkcija in zmogljivost skladienja ter prenosa anorganskega in organskega ogljika spremenita. Kanali sedimentom sluijo predvsem kot transportni mehanizmi. Merjeni podatki so redki, vendar je za uinkovito delova nje kanalov potrebno vsaj ravnoteje med koliino sedimentov, ki vstopijo in izstopijo iz vodonosnika. Ko koliina sedimentov presee vstopne vrednosti, bo nekaj sedimenta ostalo v podzem lju. Ko naravno upadanje dosee odstranitev sedimentov in prihaja le do odlaganja sige, se ta material uskladii vse do popolne denudacije kamninske mase. Skladienje se dimenta je v hidroloko aktivnih kanalih veinoma prehodno, vendar prihaja do relativnih razlik v razlinih vodonosnikih. Vodonosniki s kanali, razvitimi v ve nivojih ali kot labirinti poplavnih voda, skladiijo sorazmerno veje koliine sedi mentov. Hipogeni sistemi hranijo veje koliine usedlin kot epigeni vodonosniki, saj veinoma prevajajo raztopljeni mate rial in ne toliko kombinacije raztopljenih in suspendiranih kla stinih usedlin. Poleg tega se hipogeni vodonosniki veinoma napajajo s padavinami in tako prejmejo in skladiijo le malo sedimentov s povrine. Koliina sedimentov in organskega ogljika shranjenega v krakih vodonosnikih na globalni ravni je v tej tudiji ocenjena na 2x10 4 km 3 and 2x10 2 km 3 Koliina organskega ogljika shranjenega v paleokrasu ni ocenjena, ven dar razpololjivi podatki kaejo, da je bistveno veja od koli ine, shranjene v aktivnih krakih vodonosnikih. Ureditev teh podatkov lahko nakae, da bi zaloge nae v paleokrasu lahko sluile kot uinkoviti ponori ogljika za globalno skladienja ogljika. Paleokrake zaloge z manjo vsebnostjo ogljikovodikov morajo zagotoviti bistveno ve skladienja kot je to znailno za nepaleokrake kamnine. Kljune besede: sediment, ogljik, paleokras, morfologija kana lov, kraka hidrologija, skladienje. 1 National Cave and Karst Research Institute, 400-1 Cascades Avenue, Carlsbad, New Mexico 88220-6215, USA, e-mail: gveni@nckri.org Received/Prejeto: 3.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 203, POSTOJNA 2013 Abstract UDC 551.444:546.26 George Veni: A framework for assessing the role of karst con duit morphology, hydrology, and evolution in the transport and storage of carbon and associated sediments Karst aquifers and conduits form by dissolution of carbonate minerals and the slow release of inorganic carbon to the sur face environment. As conduits evolve in size, morphology, and position within the aquifer, their function and capacity change relative to the storage and transport of inorganic and organic carbon as sediment. Conduits serve mostly as trans port mechanisms in relation to sediments. Q uantied data are sparse, but for conduits to function eectively there must be at least equilibrium in the amount of sediment entering and exiting the aquifer. If sediment discharge exceeds input, little sediment will remain underground. W hen natural declines in base level cease removing sediments and only deposit calcite speleothems, these materials are stored until the rock mass is denuded. W hile sediment storage is mostly transient in hydro logically active conduits, relative dierences occur. Aquifers with conduits developed at multiple levels or as oodwater mazes store proportionately greater volumes of sediment. Hy pogenic systems should store greater volumes of sediment than epigenic aquifers because they mostly discharge a dissolved load as opposed to both dissolved and suspended clastic loads. However, some hypogenic aquifers are diusely recharged and receive and store little sediment from the surface. e global volume of sediment and organic carbon stored in karst aquifers is estimated in this study to be on the order of 2x10 4 km 3 and 2x10 2 km 3 respectively. e amount of organic carbon stored in paleokarst is not estimated, but available data indicate it is substantially greater than that stored in modern karst aquifers. Development of such data may suggest that paleokarst petro leum reservoirs might serve as ecient carbon sinks for global carbon sequestration. Hydrocarbon-depleted paleokarst reser voirs should provide substantially more storage per injection well than sequestration in non-paleokarstic rocks. Keywords : Sediment, carbon, paleokarst, conduit morphology, karst hydrology, sequestration.

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ACTA CARSOLOGICA 42/2-3 2013 204 Caves are commonly thought of as dirty places, so it is ironic that cave sediments have seen relatively little study as compared to topics such as cave hydrology, morphol ogy, and mineralogy. Reviews by Ford (2001), Gillieson (2004), Sasowsky (2004), and Springer (2005) nicely summarize how sediments accumulate in karstic caves as passages dissolutionally and mechanically enlarge, leav ing collapsed rock and insoluble materials behind (auto chthonous sediments). ey further explain how as cave entrances and sinkholes enlarge, increasingly greater amounts of sediment move underground through the action of water, wind, and animals (allochthonous sedi ments). Most detailed studies focus on the mechanism of cave sedimentation, generally in regard to a specic cave or karst area (e.g. Bottrell et al. 1999), while others investigate sediment stratigraphy or content (typically bone or mineral) to interpret modern to past hydrolo gies, contaminant transport, paleoclimate, and potential economic resources. Sasowsky and Mylroie (2004) pro vide several examples. Cave sediment generally includes inorganic carbon, oen in the form of speleothems, and organic carbon, oen plant and animal detritus. e study of carbon in cave sediments is a largely uninvestigated topic but one which is gaining attention. e purpose of this paper is to use the published information available on cave car bon to establish a conceptual framework on how organic and inorganic carbon enters, is stored, in some cases produced in, and ultimately returned to the surface from karst caves and aquifers. Testing and renement of this framework by future authors is encouraged. is paper focuses on the role of conduit morphol ogy, aquifer hydrology, and changes in sediment and car bon movement and distribution, as individual conduits and aquifers evolve. One section estimates the critically important total volume of sediment and carbon stored in karst conduits; paleokarst is considered separately. For the purposes of this paper, sediment within the aquifer is considered uvial in the sense of how most of it is trans ported and deposited; sediment transport by wind and animals, while at times locally signicant, rarely extends far into the aquifer system. Conduit is generally used rather than cave to include caves as well as voids too small for human access but formed by turbulent ground water ow and which collectively play signicant roles relative to sediment and carbon. Similarly, sediment is used to address the total mass of material; its application to carbon is implied. Carbon is used when specicity is required. Chemolithoautotrophic microbes potentially produce a major source of organic carbon in karst sys tems but are not considered in this paper. INTRODUCTION GEORGE VENI CONDUIT MORPHOLOGY AND HYDROLOGY One of the dening characteristics of a karst conduit is its ability to transmit turbulent groundwater ow and hence sediment. A conduits tendency to deposit, erode, or transport sediment depends on two primary factors: conduit morphology and conduit and aquifer hydrology. See Veni (2005) for a summary of conduit types and ori gins. Certain basic principles of sediment movement and distribution relative to conduit morphology and hydrol ogy must be acknowledged. e potential for sediment transport, as opposed to deposition, increases with: 1) conduit slope; 2) hydraulic gradient; 3) smoother conduit cross sections; 4) frequency of ood events; 5) magnitude of ood events; 6) position with respect to the water table; and 7) decreases in conduit width that focus erosion on passage oors. All of these factors assume at least intermittently turbulent groundwater ow through the conduit to carry sediment. Position with respect to the water table is the least reliable predictor of sediment transport; conduits high in the vadose zone have the greatest competence to erode sediments while those closer to the water table are more likely to carry converged groundwater ow with a far greater capacity to move sediment. Herman et al. (2012) provide a detailed review of factors aecting sedi ment transport through karst aquifers. Based on the above principles, rough predictions can be made about gross sediment distribution through out a particular cave or karst aquifer if some of those pa rameters are known. e availability of sediment on the surface is generally not a major factor in those principles. W hile it has an eect on short-term deposition and ero sion, the long-term function of the conduit will be deter mined by those principles; dierences in sediment input will mostly aect the degree of how sediment is trans

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ACTA CARSOLOGICA 42/2-3 2013 205 A FRAME W ORK FOR ASSESSING THE ROLE OF KARST CONDUIT MORPHOLOGY HYDROLOGY AND EVOLUTION ... ity and type of recharge. In general, mazes and multitiered caves, no matter their origin and in proportion to their active groundwater ow, have the potential to store greater volumes of sediment than other types of caves. ese types of caves have a high proportion of passage volume for sediment storage relative to the volume of their active or seasonally active streams. In both cases, ood waters rise into higher level passages and mazes where sediment is deposited. Sedimentation is especially enhanced in maze caves where the injected ood water becomes ponded or nearly ponded in the conduits, al lowing sediment to settle with little potential for remov al; high passages that can sustain signicant currents provide relatively less opportunity for sedimentation. Palmer (1991) recognized three main types of re charge: karst depressions, diuse, and hypogenic. Grav ity-drained epigenic aquifers with either depression-fo cused or diuse recharge discharge dissolved, suspended, and bedload sediments. In principle, hypogenic conduits should proportionally store more sediment as discussed below at the end of this section. Palmers (1991) classication system is also useful in assessing the type of sediment in each type of aquifer. ose recharged by sinkholes and other focused sources have the greatest sediment diversity, containing alloch thonous and autochthonous sediments with the alloch thonous sediments potentially including non-carbon ates/non-evaporites washed in from adjacent rocks. ey also contain the greatest percentage of organic carbon from soils, plants, and animals carried into the conduits by water, wind, or gravity. Diusely recharged karst aquifers contain few al lochthonous clastic sediments. Organic carbon is pri marily limited to tree roots that may penetrate conduits, and organic particles and dissolved carbon moving down fractures (e.g. Toth 1998). However, diusely re charged aquifers that discharge near valley oors may contain substantial organic carbon deposited by back ooding of surface streams into the conduits. Although not a diusely recharged karst aquifer, Mammoth Cave, USA, provides an excellent example of extensive sedi ment transport that resulted from backooding. Hen drickson (1961) and Collier and Flint (1964) found clay and silt deposited up to 6 cm thick and at least 2 km into the cave. Hypogenic aquifers discharge rising water under pressure, with gravity holding much of the suspended load and bedload in the aquifer, consequently resulting in a relatively greater accumulation of insoluble sedi ment with the soluble fraction removed in the dissolved load. Because most conduits in hypogenic aquifers are deep below the water table and not readily observable, this proportionately greater sediment accumulation is ported or deposited, not the general principles aecting transport and deposition. roughout the history of a conduit, its primary function is to transport water and to a generally lesser degree, sediment, in, through, and out of the karst aqui fer; storage is a relatively minor function. Data which quantify the volume of sediment stored in a karst aquifer are sparse, relative to the sediment that ushes through it, but W orthington (1984) provides an instructive ex ample. Analogous information is also available for com parison and to support the point. A useful study is that of W orthington et al. (2000) who found in studying four distinct types of unconned karst aquifers that the vol ume of water stored in the conduits ranged from 0.052.8% while 94-99.7% of the groundwater transported through the aquifer over time moved through the con duits. W hen considering that the same water carries sediment and dissolved solids, even though the energy and stresses needed to move sediment are greater than just moving water, it is reasonable to assume that only a small percentage of that sediment is stored in the hydro logically active part of the aquifer at a given time, while a vastly greater volume is washed through. Groves and Meiman (2001) give an example of a cave system that is discharging stored sediments and limestone eroded from cave walls in a far greater volume than the sedi ment which enters that aquifer. For most of their histories, karst aquifers must func tion close to equilibrium or at a sediment decit relative to the amount of sediment entering and exiting; how this equilibrium relates to the increase in total stored sedi ment as conduits enlarge is discussed later. Prolonged sediment surplus will result in the lling of the conduits and a nonor poorly-functional aquifer; Herman et al. (2012) oer an equation that quanties this behavior. Typically, a cave stream will erode or deposit sediment until equilibrium is reached relative to the volume, gra dient, grain size, and other conditions of the stream. Ma jor episodes of erosion or deposition may take millen nia to recover, or may recover within the time frame a single storm event. For example, Van Gundy and W hite (2009) describe a ood dislodging 1,800 m 3 of soil from a sinkhole into a cave. Aer the ood, no measureable accumulation of sediment was found inside, which con sidering the conditions of the cave strongly suggest that most was transported through. Bosch and W hite (2004) oer an equation for measuring sediment ux through a karst groundwater drainage basin. Sediment storage is mostly transient in hydrologi cally active conduits, but relative dierences occur due to dierences in conduit morphology, hydrology, and origin. Palmer (1991) identied 15 major types of con duit patterns based on a karst aquifers dominant poros

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ACTA CARSOLOGICA 42/2-3 2013 206 oen not apparent in the humanly-accessible portions of the conduit system. Hypogenic karst aquifers with focused recharge may contain the same variety of sediments at their up gradient, unconned ends as epigenic aquifers. Most of this material is deposited in the conduits prior to reach ing the downgradient end, or dissolved and discharged from the aquifer. Diusely recharged hypogenic aquifers receive few allochthonous clastic sediments, but unless CONDUIT AND A QUIFER EVOLUTION Changes in sediment transport and storage, and hence carbon transport and storage, occur over time as karst conduits and aquifers evolve. Most fundamentally, the ca pacity for sediment storage increases as conduits enlarge. Even when short-term measurements suggest a mass bal ance in sediment deposition and erosion, storage capac ity continues to increase over geologic time as conduits increase in size. Paragenetically formed conduits are ex cellent examples where sediment accumulates to armor the oor and focus conduit dissolution and enlargement upward. Natural declines in base level diversify the means and locations for sediment deposition in conduits in four basic stages as illustrated in Fig. 1: Stage 1 All water and sediment ow occurs through phreatic conduits along a roughly single zone of eleva tion. Stage 2. As base level and the karst water table de cline, the conduits are either incised to lower elevations and/or lower-elevation conduits form to carry the phre atic ow. e Stage 1 level conduits, as either high-level ledges or as distinct conduits, transmit and store sedi ments from vadose water as well as from ood-stage ris es in the water table. Vadose speleothems may begin to form; speleothems on sediments reduce the sediments susceptibility to erosion. Phreatic rises from ooding into the Stage 1 levels will deposit sediments and transport sediments downgradient through these now episodically active routes. Many Stage 1 conduits may ll completely with sediment as groundwater ow patterns shi. Stage 3. Base level and the water table continue to decline, either deepening the existing conduits further and/or creating another level of conduits at an even lower elevation. e processes described for Stage 2 now apply to the two lowest levels. e Stage 1 level now continues to vadosely transmit and store sediment but not phre atic ood ows. In many cases, this level has decreas ing hydrologic activity as more direct vertical routes to the water table form and bypass older, more horizontal conduits. Consequently, sediment deposition typically exceeds erosion in the Stage 1 level, especially sediment high in organic carbon brought in by animals and fallin/slide-in from cave entrances and sinkholes. Sedimentoccluded conduits in the Stage 1 level become common. Stage 4. Base level declines and aquifer conditions change to the point where a water table is no longer sus tainable. Flows decrease and are typically episodic, mi nor, and focused along a few active routes. More sedi diagnostic allochthonous trace minerals are transported through the system it may be impossible to distinguish focused from diuse recharge at the aquifers downgra dient end. Analyses of sediment at the downgradient end of hypogenic aquifers may suggest a predominantly autochthonous origin (e.g. Veni & Heizler 2009), which could be misleading without detailed knowledge of the recharge area and type. GEORGE VENI

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ACTA CARSOLOGICA 42/2-3 2013 207 QUANTIFYING SEDIMENT AND CARBON STORAGE Critical questions for this paper are how much sediment is stored in karst conduits and what percentage is com posed of organic carbon? W ith the current high level of interest in global carbon balance and searches for carbon sinks to sequester carbon from the atmosphere, only or ganic carbon plays a notable role in karstic conduit sedi ments. Karst conduit development and landscape denuda tion release tremendous volumes of inorganic carbon over geologic time. Groves and Meiman (2001) measured inor ganic carbon discharging from the Mammoth Cave Sys tems Logsdon River 25-km 2 drainage basin at an annual rate of 7.8x10 3 (.9x10 3 ) kg ha 1 Only 1% of that total entered the aquifer with recharging water, demonstrating that most is being actively dissolved from the cave walls and/or stored sediment. Only a small amount of this in organic carbon is recaptured in the karst environment as speleothems in conduits and travertine and tufa at springs. However, while it is so much volumetrically greater than organic carbon, inorganic carbon in karst is of less im portance relative to climate change because it is naturally sequestered in karst aquifers at geologic time scales while organic carbon is produced biologically and released for storage into karst aquifers at far more rapid rates. Herman et al. (2012) provide a detailed overview of sediment transport studies in karst aquifers and oer an insightful analysis of the role of uid hydrodynanics. Un fortunately, they do not distinguish between organic and inorganic carbon in the sediments. In fact, few data are available to accurately quantify the amount of organic car bon stored in karst conduit sediments. Most early studies of sediment deposits describe organic components qualita tively as thin beds, bones, and/or black staining visible in proles, but do not quantify their volumes (e.g. Davies & Chao 1959; Bramwell 1964). Most published analytical studies of conduit sediments have focused on the miner alogy and either do not analyze for organic carbon or do not report such results (e.g. Gospodari 1974). Several re cent reports examine total and dissolved organic carbon in karst groundwater as tracers and indicators of contamina tion (e.g. Batiot et al. 2003). Fig. 1: Conceptual four-stage model of karst aquifer evolution illustrating increases in sediment storage capacity and depositional pat terns as related to conduit development and groundwater ow conditions. ment is deposited than removed through most of the conduit system, with extensive sections of the system po tentially lled completely. W hile primarily a deposition al stage, the rate of deposition is relatively slow due to the minor and episodic transport mechanisms of avail able water, wind, and biota. e sediment is primarily stored in the conduits until the aquifer is hydrologically reactivated by rises in base level or released by denuda tion of the karst landscape. Mihevc et al. (1998) describe examples of sediment-lled caves exposed by landscape denudation. is four-stage model is deliberately kept simple for conceptual understanding of how conduit evolu tion aects sediment/carbon transport and storage. Lo cal variations in geology and hydrology will accentuate or diminish various features associated with each stage. ey are illustrated in Fig. 1 under epigenic conditions, but similar changes occur in hypogenic aquifers. As wa ter levels decline in hypogenic karst aquifers, comparable stages take place as total storage capacity increases, ow routes are modied, and in sediment and calcite deposi tional patterns. A FRAME W ORK FOR ASSESSING THE ROLE OF KARST CONDUIT MORPHOLOGY HYDROLOGY AND EVOLUTION ...

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ACTA CARSOLOGICA 42/2-3 2013 208 However, there are a few studies that quantify the volume of organic carbon. For example, Herman et al. (2007) investigated suspended sediment from three karst springs where carbon was found only in trace concentra tions. Simon et al. (2007) provide a conceptual model for the ow organic carbon through karst aquifers. Study ing two caves in temperate climates, one in Europe and the other in North America, they determined that most organic carbon in karst aquifers is dissolved and enters at mean respective rates of 4.36 and 7.67 mg C L for sinking streams and 0.70 and 1.1 C L for epikarstic drips. ese studies and the literature in general suggest that while karstic sediments commonly contain carbon ate rock fragments and minerals, organic carbon is less dense and decomposes far more rapidly than inorganic carbon dissolves, and thus is removed and released into the global environment at far more rapid rates. e modern interest in sediment storage stems from the desire to understand total sediment and car bon stored in karstic conduits and its potential role in the global carbon budget. Unfortunately, an accurate es timate is impossible given the paucity of data and wide margins of error with the available information. In an ticipation of better data, and as a guide to which data are critical to obtain, the following equations are proposed for quantifying the volumes: Gcs = Ak x Dk x Vc x Sv (eq. 1) where: Gcs = global conduit sediment volume Ak = global karst area Dk = average depth of karstication Vc = average percent of conduit volume Sv = average percent of conduit volume lled with sediment and Goc = Gcs x Co (eq. 2) where: Goc = global organic carbon in conduit storage Co = average percent of organic carbon in conduits Despite the few available data, the equations are used below to roughly estimate global sediment and car bon stored in karst aquifers. Ak equals 1.7 x 10 7 km 2 based on Hollingsworths (2009) mapping of global carbonate and evaporite karst areas. is should be considered a minimum gure. Hol lingsworth recognizes signicant evaporite karst areas that could not be included in her study. W hile she did map pseudokarst areas, they are not considered here be cause their caves are not known to accumulate signi cant sediment volumes and their distribution of caves re mains both poorly understood and unquantied. Lastly, Hollingsworths totals reect only karst outcrops and do not account for signicant karstication of rocks over lain by non-karstic units. Dk is estimated as 0.2 km as a global average. To date there has been no measured average global thick ness of carbonate and evaporite units. e United Na tions (2011) reports on the mean thickness of various limestone and dolomite units in Europe as ranging from 150 m. W hile some karstied units are much thin ner and thicker, 200 m is suggested as a probable con servative average, not just for bedrock thickness but also as a probably conservative estimate of the thickness of the zone of most signicant karstication within those units. Vc is set at 1.33%. As previously mentioned, W or thington et al. (2000) studied four distinct types of un conned karst aquifers and found conduit porosity ranged from 0.05.8%, or an average of 1.33%. Given the broad range of hydrologic conditions those aquifers represent, their average is used as a possibly conserva tive estimate of global mean potential conduit storage; a qualitative review of the worlds karst outcrops suggests that most are similar to the aquifers with higher conduit porosities. Sv is the least understood value. W hile conduits are known to span the range of containing eectively no sediment to being totally occluded, most cave sedi ment studies and observations are made with little or no knowledge of how much of the cave is actually lled. For this study, Sv is proposed as 50%. is may instinctively seem too high a proportion of sediment ll, but that in stinct is almost certainly biased from predominantly ob serving caves that have relatively little ll and are thus accessible. Casual observation of road cuts and quarries in karstied rock around the world suggest that sedi ment lled conduits are generally as common as open conduits, if not more so. Calculating the above values with equation 1, the global volume of sediments stored in conduits is esti mated roughly as 2.2 x 10 4 km 3 is gure is potentially very conservative based on conservative estimation of all of the parameters, but this is not assured consider ing the uncertainty of the datas precision. Taking this a step further, Co is grossly approximated as an average of 1% of the volume of global conduit sediment based on the discussion earlier in this section. erefore, fol lowing equation 2, the estimated total global volume of organic carbon stored as sediments in karst conduits is 2.2 x 10 2 km 3 GEORGE VENI

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ACTA CARSOLOGICA 42/2-3 2013 209 PALEOKARST: CARBON SOURCE AND SINK Much of this paper has focused on the movement and oen short-term storage of sediment through karst aqui fers. However, once hydrological processes become inac tive and/or the area buried under other deposited rocks, the sediment could be stored for long periods of geologic time in paleokarst. Perhaps the oldest example is in the W estern Transvaal of South Africa with the age of karsti cation dating to 2.2 Gy (Martini, 1981). Paleokarst areas occur throughout the world. ey are included in this papers discussion because they are well recognized sources of carbon, much of which is be ing burned as fuel oil and gas, which are widely consid ered among the primary factors for climate change as the carbon accumulates in the atmosphere. e reason for addressing them here is to provide a rough framework for the volume of carbon they hold in long-term geologic storage as a contrast to the relatively short-term geologic storage described in the preceding sections. Many of the worlds most productive petroleum reservoirs are paleokarst. Notable examples include the Cretaceous Mishrif Formation in the Persian Gulf region (Farzadi & Hesthammer 2007) and the Ordovician El lenburger Limestone (Kerans, 1989) and Permian San Andres Dolomite (Craig, 1988) in the Permian Basin of west Texas, USA. However, published global values on hydrocarbons stored in paleokarst seem rare. Schlum berger (2007) reports that carbonate rocks respectively store 60% and 40% of the global oil and gas reserves; they further report that 70% and 90% of respective Mid dle East oil and natural gas reserves are held within car bonate reservoirs but do not report the actual volumes. Data on hydrocarbon production are more available. As an example, the Permian Basin is one of the worlds lon gest major oil-producing regions and since production began in 1921 through the end of 2012, about 29 billion barrels of oil (4.61 km 3 ) have been recovered (Railroad Commission of Texas, 2012), with most yielded from pa leokarst. is production gure for this reservoir alone is more than an order of magnitude greater than the total organic carbon stored in modern karst as calculated in the previous section. A focused and in-depth study is needed to precise ly estimate the volume of organic carbon stored in pa leokarst, but what purpose would such research serve? eoretically, much of the volume of oil and natural gas that is removed from paleokarst reservoirs could be re placed by concentrated carbon dioxide for sequestra tion. Data on the global volume of paleokarst hydro carbon storage could be used to compare their potential carbon sequestration storage potential relative to other reservoirs. Oil and gas production data from known pa leokarst reservoirs, such as the Permian Basin, suggest that the high porosity and permeability of paleokarst, as compared to many fracture and porous media systems, may allow for much faster injection and greater storage capacity per well or unit area. W hile this hypothesis re mains to be tested, unlike the mostly short-term carbon storage in modern, hydrologically active karst systems, sequestration in deeply buried paleokarst would be ef fectively permanent relative to human history. How does this potential volume compare to car bon in the global atmosphere? e W orld Meteorologi cal Organization (2011) estimated that 3.8 x 10 9 metric tons of carbon as carbon dioxide were released into the atmosphere since the onset of the Industrial Revolution in 1750 through 2010. In gas form, as a fraction of car bon dioxide, that totals 6.5 x 10 3 km 3 of carbon. How ever, organic carbon sediment stored in karst conduits is concentrated in solid form. e density of carbon in such sediments varies and has not been quantied. How ever, carbon as graphite has a maximum density prob ably more than an order of magnitude greater than the density of typical solid organic carbon in cave sediments. erefore, the volume of organic carbon released into the atmosphere since 1750 is probably approximately equivalent to the volume of organic carbon estimated in the previous paragraph as stored in karst. CONCLUSIONS Sediment is a well recognized feature that occupies and lls many karst conduits, but its study is not as common as its appearance. is paper conceptually evaluates sedi ment occurrence in karst aquifers and nds it a transient A FRAME W ORK FOR ASSESSING THE ROLE OF KARST CONDUIT MORPHOLOGY HYDROLOGY AND EVOLUTION ...

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ACTA CARSOLOGICA 42/2-3 2013 210 REFERENCES Batiot, C., Emblanch, C. & B. Blavou, 2003: Total Or ganic Carbon (TOC) and magnesium (Mg2+): two complementary tracers of residence time in karstic systems.Comptes Rendus Geoscience, 335, 2, 205. Bosch, R.F. & W .B. W hite, 2004: Lithofacies and trans port of clastic sediments in karstic aquifers.In: Sasowsky, I.D. & J. Mylroie (eds.) Studies of Cave Sediments P hysical and Chemical Records of P aleo climate, Klewer Academic/Plenum Publishers, pp. 1, New York. Bottrell, S., Hardwick, P. & J. Gunn, 1999: Sediment dy namics in the Castleton Karst, Derbyshire, United Kingdom.Earth Surface Processes and Landforms, 24, 745. Bramwell, D., 1964: e excavations at Elder Bush Cave, W etton, Staordshire.North Staordshire Journal of Field Studies, 4, 46. Collier, C.R. & R.F. Flint, 1964: Fluvial Sedimentation in Mammoth Cave, Kentucky. US Geological Survey Professional Paper 475D, 141. Craig, D.H., 1988: Caves and other features of Permian karst in San Andres Dolomite, Yates Field reservoir, west Texas.In: James, N.P. & P.W Choquette (eds.) P aleokarst Springer-Verlag, pp. 342, New York. Davies, W .F. & E.C.T. Chao, 1959: Report on the Sedi ments in Mammoth Cave, Kentucky.U.S. Geologi cal Survey Administrative Report to the U.S. Na tional Park Service, W ashington. Farzadi, P. & J. Hesthammer, 2007: Diagnosis of the Up per Cretaceous palaeokarst and turbidite systems from the Iranian Persian Gulf using volume-based multiple seismic attribute analysis and pattern rec ognition.Petroleum Geoscience, 13, 227. Ford, T.D., 2001: Sediments in Caves.BCRA Cave Stud ies No. 9, Derbyshire. Gillieson, D., 2004: Sediments: Allochthonous Clastic.In: Gunn, J. (ed.), Encyclopedia of Caves and Karst Fitzroy Dearborn, pp. 633, New York. Gospodari, R., 1974: Fluvial sediments in Krizna Jama.Acta Carsologica, 4, 327. material with well over 90% moved through the aquifer over time. e remaining percentage is stored in conduits for long periods, especially as the aquifer evolves and be comes hydrologically inactive, until the aquifer regains sucient groundwater ow to erode the sediments or until the karstic bedrock and its conduits are removed by landscape denudation. Maze and multilevel caves have the capacity to store greater volumes of sediment than most other cave types. Equations are proposed, using the area of mapped global karstic outcrops, mean thickness of karstic units, average conduit volume within karst aquifers, estimated fraction of that volume, and its estimate percentage of organic car bon to set a lower limit on the total global conduit sedi ment at 2.2 x 10 4 km 3 and its total organic carbon content as 2.2 x 10 2 km 3 is volume of organic carbon is similar to the volume released into the atmosphere since the on set of the Industrial Revolution in 1750. Sediment and carbon volumes were not estimated for paleokarst petro leum reservoirs, but the rate and volume they release oil suggests they may prove excellent reservoirs to eciently and essentially permanently sequester large volumes of atmospheric carbon dioxide. ACKNO WLEDGEMENTS e paper was prepared for the Carbon and Boundaries in Karst Conference co-organized by the Karst W aters Insti tute and the National Cave and Karst Research Institute. I appreciate their proposing that thought-provoking sub ject, which prompted this paper. Discussions with and information provided by Drs. Derek Ford and W illiam W hite during and aer the conference proved helpful. I especially thank Dr. E. Calvin Alexander Jr. for his care ful review and insightful comments on the manuscript. GEORGE VENI

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ACTA CARSOLOGICA 42/2-3 2013 211 Groves, C. & J. Meiman, 2001: Inorganic carbon ux and aquifer evolution in the South Central Ken tucky Karst.In: Kuniansky, E.L. (ed.), U.S. Geologi cal Survey Karst Interest Group P roceedings, USGS W ater-Resources Investigations Report 01, 13 February 2001, St. Petersburg, Florida, pp. 99, W ashington. Hendrickson, G.E, 1961: Sources of W ater in Styx and Echo Rivers, Mammoth Cave, Kentucky.US Geol Survey W ater Supply Paper, 424, 41. Herman, E.K., Tancredi, J.H., Toran, L. & W .B. W hite, 2007: Mineralogy of suspended sediment in three karst springs.Hydrogeology Journal, 15, 2, 255. Herman, E.K., Toran, L. & W .B. W hite, 2012: Clastic sediment transport and storage in uviokarst aqui fers: an essential component of karst hydrogeology.Carbonates and Evaporites, 27, 3, 211. Hollingsworth, E., 2009: Karst regions of the world (KROW)populating global karst datasets and gen erating maps to advance the understanding of karst occurrence and protection of karst species and habi tats worldwide. Masters thesis, University of Arkan sas, pp. 138. Kerans, C., 1989: Karst-controlled Reservoir heteroge neity and an example from the Ellenburger Group (Lower Ordovician) of W est Texas.Report of In vestigations No. 186, University of Texas at Austin Bureau of Economic Geology, Austin. Martini, J., 1981: Early Proterozoic paleokarst of the Transvaal, South Africa.In: Beck, B.F. (ed.), P ro ceedings of the Eighth International Congress of Spe leology, volume 1, National Speleological Society, pp. 6, Huntsville. Mihevc, A., Slabe, T. & S. ebela, 1998: Denuded caves an inherited element in the karst morphology, the case from Kras.Acta Carsologica, 27(1). 165. Palmer, A.N., 1991: Origin and morphology of limestone caves.Geological Society of America Bulletin, 103(1), 1. Railroad Commission of Texas, 2012: Permian Basin in formation.[Online] Available from: http://www. rrc.state.tx.us/permianbasin/index.php [Accessed 5 January 2013]. Sasowsky, I.D., 2004: Sediments: Autochthonous Clas tic.In: Gunn, J. (ed.), Encyclopedia of Caves and Karst, Fitzroy Dearborn, pp. 634, New York. Sasowsky, I.D. & J. Mylroie (eds.), 2004: Studies of Cave Sediments P hysical and Chemical Records of P aleo climate.Klewer Academic/Plenum Publishers, pp. 329, New York. Schlumberger, Inc., 2007: Carbonate reservoirs Meeting unique challenges to maximize recovery.Schlum berger, Inc., pp. 16. Simon, K.S, Pipan, T. & D.C. Culver, 2007: Journal of Cave and Karst Studies, 69, 2, 279. Springer, G.R., 2005: Clastic sediments in caves.In: Cul ver, D.C. & W .B. W hite (eds.) Encyclopedia of Caves, Elsevier, pp. 102, Burlington. Toth, V.A., 1998: Spatial and temporal variations in the dissolved organic carbon concentrations in the va dose karst waters of Marengo Cave, Indiana.Jour nal of Cave and Karst Studies, 60, 3, 167. United Nations, 2011: Assessment of transboundary riv ers, lakes and groundwaters discharging into the Mediterranean Sea.Economic and Social Council, Economic Commission for Europe, Meeting of the parties to the convention on the protection and use of transboundary watercourses and international lakes, pp. 85. Van Gundy, J.J. & W .B. W hite, 2009: Sediment ushing in Mystic Cave, W est Virginia, USA, in response to the 1985 Potomac Valley ood.International Jour nal of Speleology, 38, 2, 103. Veni, G., 2005: Passages.In: Culver, D.C. and W hite, W .B. (eds.) Encyclopedia of Caves, Elsevier, pp. 436, Burlington. Veni, G. & L. Heizler, 2009: Hypogenic origin of Rob ber Baron Cave: implications on the evolution and management of the Edwards Aquifer, central Texas, USA.In: Staord, K.W et al. (eds.) NCKRI Sym posium 1, Advances in H ypogene Karst Studies, Na tional Cave and Karst Research Institute, pp. 85, Carlsbad. W orld Meteorological Organization, 2011: e state of greenhouse gases in the atmosphere based on glob al observations through 2010.Greenhouse Gas Bulletin, 7, 1. W orthington, S.R.H., 1984: e paleodrainage of an Ap palachian uviokarst: Friars H ole, West V irginia. Masters thesis, McMaster University, pp. 218. W orthington, S.R.H., Ford, D.C. & P.A. Beddows, 2000: Porosity and permeability enhancement in uncon ned carbonate aquifers as a result of solution.In: Klimchouk A.B. et al. (eds.) Speleogenesis: Evolution of Karst Aquifers, National Speleological Society, pp. 463, Huntsville. A FRAME W ORK FOR ASSESSING THE ROLE OF KARST CONDUIT MORPHOLOGY HYDROLOGY AND EVOLUTION ...



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B IOLOGICAL C ONTROL ON A CID GENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : QUANTIFICATION THROUGH G EOCHEMICAL M ODELING VPLIV BIOLO KIH PROCESOV NA IZVOR KISLIN OB MEJI MED KRA KIM PREVODNIKOM IN MATI NO KAMNINO V ZALITI CONI: KVANTITATIVNA OCENA NA OSNOVI GEOKEMI NEGA MODELIRANJA Janet S. HERMAN 1* Alexandria G. H OUNSHELL 1 Rima B. FRANKLIN 2 & Aaron L. MILLS 1 Izvleek UDK UDK:550.46:551.444 Janet S. Herman, Alexandria G. Hounshell, Rima B. Franklin & Aaron L. Mills: Vpliv biolokih procesov na izvor kislin ob meji med krakim prevodnikom in matino kamnino v zaliti coni: kvantitativna ocena na osnovi geokeminega modeliranja Jama No-mount je zalita sladkovodna jama v dravnem parku W ekiva Springs, Florida, ZDA. Skozi jamo se proti povrju pretaka voda iz Zgornjega Floridskega apnenastega vodonosnika eocenske starosti. V afotinem okolju jame so kolonije oksidirajoih bakterij v obliki vlaknastih skupkov. Na osnovi terenskih, laboratorijskih in modelskih raziskav smo ocenje vali moen speleogenetski pomen bakterijske oksidacije sul dov. Ocenjevali smo raztapljanje kalcita v razlinih scenarijih meanja vode iz apnenaste matrice in jamske vode ob prisotnosti oksidacije sulda. Pri geokeminem modeliranju smo uporabili modul reakcijskih poti v modelskem okolju PHRE EQ CI, kot vhodni podatek modela smo vzeli sestavo vode iz matrice in jamskega rova. Za vzorenje vode v matrici apnenca smo v okviru te raziskave razvili nov vzorevalnik. Laboratorijske raziskave so temeljile na pretonem reaktorju, pri emer smo v reaktorskem stolpcu uporabili zdrobljen apnenec z mesta raziskav in v njem meali laboratorijske replike vode po rozne matrice obogatene s suldom in jamske vode z dodanim kisikom. Ob odsotnosti bioloke aktivnosti ni bilo raztapljanja ob meanje matrine in jamske vode. Do obutnega raztaplja nja pa je prilo ob dovajanju protonov, ki se sprostijo pri bioloki pretvorbi v vodi raztopljenih veplovih spojin. Iz irega spektra rezultatov smo izluili relevantno vrednost. Ocenili smo, da se v jamsko vodo sprosti 158 mg kalcija na vsak liter vode, ki iz porozne matrice preide v jamski rov in se mea z jamsko 1 Department of Environmental Sciences, University of Virginia, P.O. Box 400123, Charlottesville, VA 22904-4123 2 Department of Biology, Virginia Commonwealth University, 1000 W est Cary Street, Room 126, Richmond, VA 23284-2012 *Corresponding Author, e-mail: jherman@virginia.edu Received/Prejeto: 1.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 213, POSTOJNA 2013 Abstract UDC UDK:550.46:551.444 Janet S. Herman, Alexandria G. Hounshell, Rima B. Franklin & Aaron L. Mills: Biological Control on Acid Generation at the Conduit-Bedrock Boundary in Submerged Caves: Quanti cation through Geochemical Modeling No-mount Cave, located in W ekiwa Springs State Park in cen tral Florida, USA, is an aphotic, submerged, freshwater cave in which large colonies of sulfur-oxidizing bacteria live in la mentous microbial mats. Upwardly discharging groundwater enters the cave from the Upper Floridan aquifer, specically the Eocene-aged Ocala Limestone. W e undertook a combined eld, laboratory, and modeling study in which we sought to deter mine the amount of calcite dissolution attributable to the gen eration of protons by microbially mediated sulde oxidation. e chemical compositions of groundwater within the lime stone formation collected through a newly designed sampling device and of water in the cave conduit were used in geochemi cal modeling. W e used the reaction-path model PHREEQ CI to quantify the amount of calcite dissolution expected under various plausible scenarios for mixing of formation water with conduit water and extent of bacterial sulde oxidation. Labora tory experiments were conducted using ow-through columns packed with crushed limestone from the study site. Replicate columns were eluted with articial groundwater containing dissolved HS in the absence of microbial growth. W ithout bio logically mediated sulde oxidation, no measurable calcite dis solution occurred in laboratory experiments and no additional amount of speleogenesis is expected as formation water mixes with conduit water in the eld. In contrast, signicant calcite dissolution is driven by the protons released in the biological transformation of the aqueous sulfur species. Although a range of results were calculated, a plausible amount of 158 mg Ca 2+ released to conduit water per liter of groundwater crossing the formation-conduit boundary and mixing with an equal volume

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ACTA CARSOLOGICA 42/2-3 2013 214 J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS of conduit water was predicted. Our modeling results indicate that signicant cave development can be driven by microbially mediated sulde oxidation under these hydrogeochemical con ditions. Keywords : calcite dissolution, microbial sulde oxidation, geo chemical model. vodo v enakem razmerju. Na model kae, da lahko mikrobsko pogojena oksidacija suldov pomembno vpliva na speleogenezo v pogojih podobnih naim. Kljune besede : raztapljanje kalcita, mikrobska oksidacija suldov, geokemini model. INTRODUCTION Hypogenic cave communities, devoid of any light and based on energy produced by chemoautotrophs, have been described in environments ranging from subma rine caves in Italy to large caverns in Mexico (Bottrell et al. 1991; Hose et al. 2000; Mattison et al. 1998; Sarbu et al. 1996; Vlasceanu et al. 2000). In each cave, microbial mats were observed in areas with high hydrogen sulde concentrations, and sulde oxidation to elemental sul fur or to sulfate is the energy source for carbon xation into biomass. Mats of sulfur-oxidizing bacteria can occur along the water-atmosphere interface in cave pools and streams ( e.g. Hubbard et al. 1986) or, more commonly, at the interface between the cave conduit and the limestone cave walls where sulde-rich groundwater seeps through the bedrock (Mattison et al. 1998). e biogeochemical reactions carried out by sulfur-oxidizing bacteria could have a signicant eect on limestone dissolution and cave enlargement (Engel & Randall, 2011; Hose et al. 2000; Macalady et al. 2006; Sarbu et al. 1996; Vlasceanu et al. 2000). Calcite, the main reactive mineral in limestone, is dissolved in the presence of H + ions in solution and is, therefore, sensitive to the acidity of natural waters. e dissolution of CO 2 gas to form carbonic acid in natural waters is the most common source of acid that dissolves calcite (W hite 1988). In an environment in which sul de is undergoing oxidation, however, there is another signicant source of acid. One of the products of sulfur oxidation (Eqns. 1 and 2) is sulfuric acid which ultimate ly dissociates to H + ions in aqueous solutions (Erlich & Newman 2009). H 2 S (aq) + O 2 (g) S 0 (s) + H 2 O (1) S 0 (s) + 1 O 2 (g) + H 2 O 2H + + SO 4 2 (2) Due to the production of acid, an increase in calcite dis solution (Eqn. 3) is expected when sulde oxidation, ei ther biotic or abiotic, is present. CaCO 3 (s) + H + Ca 2+ + HCO 3 (3) A recent study conducted in the sulde-rich, saline portion of the karstic Edwards Aquifer in Central Texas found increased extent of calcite dissolution in the pres ence of sulde-oxidizing bacteria when compared to abi otic calcite dissolution in situ (Engel & Randall, 2011). e aquifer is developed in limestone and includes abundant, lamentous, microbial mats which were de termined to contain sulde-oxidizing bacteria along the limestone wall of the open-hole wells. Microcosms com prising calcite chips were suspended in wells, one treat ment inoculated with sulde-oxidizing bacteria and one that was not. e amount of dissolution was determined by weight loss of the calcite chips. e results were direct evidence of an increased extent of calcite dissolution due to sulde-oxidizing bacteria in limestone aquifers. e extent to which sulfuric-acid speleogenesis in uences cave enlargement may be great, as indicated by the acidity of that environment. In several subaerial caves containing microbial mats, very low pH readings (be tween 0 and 1) were measured within the water droplets known as snottites hanging from the mats (Hose et al. 2000; Sarbu et al. 1996), yet the cave streams remain near neutral indicating neutralization by calcite dissolution. Very few studies, however, have included quantication of the extent of calcite dissolution due to microbially me diated sulde oxidation in cave development. Most caves with abundant chemoautotrophic bac teria that have been studied to date were either subaerial or submerged by saline waters. Some freshwater-lled caves, however, have been shown to contain extensive microbial mats as well. One such setting, No-mount Cave feeding the W ekiwa Spring in central Florida, is the focus of the present study in which we seek to quantify the extent of calcite dissolution attributable to sulde-ox idizing bacteria. Our conceptual model of the geochemi cal conditions leading to speleogenesis is illustrated in Fig. 1, in which sulde-containing groundwater within the limestone formation enters a cave conduit through a bacterial mat covering the cave walls. W e believe that sulde oxidation is occurring at the base of the bacterial mat near the cave wall, where the exceptionally so tex ture of the limestone is in sharp contrast with the harder,

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ACTA CARSOLOGICA 42/2-3 2013 215 METHODS FIELD S ITE Groundwater of the Upper Floridan aquifer discharges at W ekiwa Springs located in W ekiwa Springs State Park, near Apopka (28 42 42.79 N, 81 27 37.52 W ) north of Orlando, Florida, USA (Ferguson et al. 1947). e spring discharge is derived from two observed features: a vertical sha and a horizontal cavern, both of which are disconnected from surface recharge features. e main boil is essentially a vertical sha (down to at least 20 m, the maximum depth reached by this research team) and is responsible for the majority of the total discharge from W ekiwa Springs. is ow combines with discharge from a smaller horizontal vent (approximately 4.2 0.6 m) lo cated about 5 m below the water surface, which serves as the entrance to No-mount Cave. e term no mount refers to the dive-gear conguration required to access the cave passage. No-mount Cave is developed in bedrock of the Oc ala Limestone, an Eocene-aged formation of predomi nantly limestone and occasional dolostone layers that is conned by the overlying Hawthorn Formation com posed mostly of lower permeability clay, silt, and sand beds (Miller 1986). e upper facies of the Ocala Lime stone is poorly indurated fossiliferous limestone (Scott 1991) and is an important, highly conductive component of the Floridan Aquifer System (Miller 1990). B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ... more competent limestone that is not covered in bac terial mat. A simulation of abiotic dissolution was car ried out experimentally in laboratory experiments, and the extents of abiotic and of bacterially mediated calcite Fig. 1: In our conceptual model of the eld situation, groundwater (herein termed formation wa ter) containing dissolved reduced sulfur (depicted here as HS ) passes from the porous limestone rock walls of the cave where it is acted upon by sulfur-oxidizing microbes in the abundant mats growing on the cave walls. In the presence of dissolved oxygen from the conduit water, the bac teria oxidize the sulde releasing SO 4 2 and protons which then can accelerate the dissolution of the limestone at the cave wall. dissolution were modeled using a geochemical reactionpath model constrained by dissolved ion concentrations observed in the eld and in the laboratory.

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ACTA CARSOLOGICA 42/2-3 2013 216 e submerged entrance of the cave lies along a hor izontal bedding plane near the bottom of the spring ba sin. ere is a small photic zone at the entrance followed by a narrow opening into a larger (5 2 2 m) aphotic cavern that penetrates down into the limestone bedrock at a small angle o vertical. e ceiling and walls of the aphotic cavern are covered by lamentous microbial mats with long white streamers that oen extend 5-10 cm from the rock surface and that support an ecosystem of small invertebrates (Franklin et al. 2005). Similar mats have been observed in all areas of No-mount Cave that have thus far been explored and can also be found in sporadic locations in the sha of the nearby main boil. Greatest biomass accumulation is observed in sheltered areas and crevices with low ow turbulence. e microbial mats are dominated by sulfur-oxidiz ing bacteria (Franklin et al. 2011; Hill & Franklin 2012). Initially, the nature of the bacteria was conrmed visual ly by the presence of sulfur granules within these bacte rial cells using phase-contrast microcopy. Subsequently, molecular genetic analyses have been conducted using whole-community DNA extracted from the mat sam ples. Cloning and sequencing of the 16S-rRNA gene has demonstrated the presence of several groups of sulfuroxidizing bacteria including: iobacillus, iospira, and Beggiatoa (Hill & Franklin 2012). FIELD SAMPLING W ater and microbial mat samples were collected from No-mount Cave on ve dates: January 19 and July 15 2005, January 16 and July 7 2006, and July 23 2007. Sam ples were obtained by technical divers from the Cam brian Foundation ( www.cambrianfoundation.org ) using the best sterile technique possible given the challenging environment. e divers utilized no mount tank con guration, closed and semi-closed circuit re-breathers, and Nitrox gas, all in an eort to minimize their impact on the cave structure, avoid unnecessary destruction of the microbial mats, and limit oxygen and carbon diox ide bubble release into the water column. Five sampling stations were established within No-mount Cave during the January 2005 dive and visited on each subsequent date. In addition, water samples were obtained from the main boil of W ekiwa Springs in July 2007 and from near by groundwater wells during January 2005. e ground water wells were located ~500 m away from the spring pool and are maintained by the St. Johns River W ater Management District (Palatka, FL). e land surface at the well locations is about 20 m above sea level, and the cave is within a meter of sea level. For the purposes of this study, wells are referred to as Deep (SJR WMD ID ORO 547, which reaches the Lower Floridan aquifer; total depth: 200 m, cased to 135 m), Mid (ORO 548, which reaches the Upper Florian aquifer the source of water to W ekiwa Cave according to Ferguson et al. (1947); total depth: 50 m, cased to 30 m), and Shallow (ORO 546; total depth: 20 m, cased to 15 m). CHEMICAL ANALYSIS At each station in No-mount Cave, divers collected a single large sample of water from the bulk conduit ow (~1.5 L) for immediate analysis on the surface using a water quality sonde (Hydrolab Q uanta, Loveland, CO) to determine temperature (C), specic conductivity (SpC, mS/cm), pH, dissolved oxygen (DO, mg/L), oxidationreduction potential (ORP, mV), and salinity (PSS). Trip licate water samples (250 mL) were collected at each sta tion for eld analysis using Chemets kits (CHEMetrics, Inc., Midland, VA) to determine sulde (Kit Catalog No. K-9510), iron (K-6210), ammonia (K-1510), and alka linity (K-9810 and K-9815). Triplicate 50-mL samples were collected, ltered on the surface using a 0.2-m syringe-tip lter (Millipore, Billerica, MA), and stored frozen until laboratory analysis using a Dionex-120 Ion Chromatograph (Sunnyvale, CA) to determine the con centration (mg/L) of chloride (Cl ), sulfate (SO 4 2 ), ni trate (NO 3 ), phosphate (PO 4 3 ), calcium (Ca 2+ ), sodium (Na + ), magnesium (Mg 2+ ), potassium (K + ), and ammo nium (NH 4 + ). An identical suite of analyses were per formed on water samples from the groundwater wells and main spring boil. L ABORATORY C OLUMN REACTORS To determine the extent of the eects of abiotic sulde oxidation on calcite dissolution, replicate columns were constructed to imitate the environmental conditions in No-mount Cave. Samples of limestone obtained as loose material from the cave oor were drained of wa ter, air-dried, and stored dry in the laboratory for several months before being crushed using a mortar and pestle and then sieved to grain size between 0.5 and 1.0 mm. Each column was built from a piece of 7.5-cm diameter PVC pipe about 20 cm in length. e bottom of each cyl inder was lled with a 10-cm-thick layer of quartz sand grains sieved as the carbonate upon which a 1.5-cm-thick layer of the crushed limestone was placed. e sand acted as a hydraulic diuser to ensure even ow through the column, while the crushed limestone creates the inter face between the groundwater within the limestone host ing No-mount Cave and the conduit water lling the cave (Fig. 2). Articial groundwater (Fig. 2) was prepared by dis solving analytical salts in the laboratory in proportions that yielded a chemical composition similar to that of the water owing through the cave (Tab. 1). e pH of J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS

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ACTA CARSOLOGICA 42/2-3 2013 217 the articial groundwater was adjusted by small addi tions of strong acid or base. Laboratory temperature was similar to that in the eld in Florida. In order to reduce oxygen saturation in the articial groundwater, the solu tion was sparged with N 2 gas to displace the oxygen, re sulting in a headspace of N 2 above the groundwater res ervoir. A mylar balloon was used to provide makeup N 2 that displaced water as it was pumped from the reservoir. e resulting dissolved oxygen concentrations were low enough to accurately reect the low oxygen concentra tions observed in the groundwater in the limestone aqui fer surrounding No-mount Cave. Na 2 S salt was added to Fig. 2: Schematic of a laboratory column reactor. Articial ground water sparged of O 2 and contain ing HS owed through the lime stone layer to represent formation water, herein called groundwater. Articial conduit water began as the same composition as ground water except without HS and with O 2 and it was recirculated following mixing with formation water and reaction with the sur face of the calcite layer. B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ...

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ACTA CARSOLOGICA 42/2-3 2013 218 the articial groundwater to achieve a dissolved HS con centration similar to eld values. In contrast, the arti cial conduit water was chemically identical to the arti cial groundwater except it was not sparged with N 2 nor was Na 2 S salt added (Tab. 1). Groundwater (anaerobic, articial conduit wa ter with HS ) was pumped from its reservoir into the bottom of the column and through the sand and crushed limestone layers at a discharge rate of about 2.73 10 2 mL/min, resulting in a 9.5-day residence time of water in the sediment-lled portion of the col umn and a 30.5-hr residence time within the limestone layer. In the head space above the crushed limestone, aerobic articial conduit water was pumped at a veloc ity of 3.03 mL/s resulting in about a 2.0-cm-deep layer of water on top of the limestone layer with a residence time of about 30 seconds. Once the conduit water owed across the surface of the limestone layer, mixed with the impinging groundwater, and owed out of the reaction column, it was returned to the conduit-water reservoir and was subsequently recirculated. About 50 mL of the simulated groundwater and conduit water were collected twice each week for four weeks. e samples were taken via a glass tube inserted into the cap covering each cave-water reservoir. CHEMICAL ANALYSIS e pH, alkalinity, and Ca 2+ concentration of each sam ple was measured within 24 hours of collection, while the remaining sample was refrigerated until further analysis. e pH was measured immediately upon collection us ing a pH electrode. Alkalinity was determined through the inection-point titration method (Rounds 2006). Calcium concentrations were determined by titrating the water sample with EDTA (American Public Health As sociation 1995), with 0.4 g/titration of the crushed Eri chrome Blue indicator. e anions Cl SO 4 2 and NO 3 were determined by ion chromatography within 2 weeks of sampling. Sulde concentrations were measured using the colorimetric method described by Otte and Morris (1994) and Cline (1969) within one week of sample col lection. GEOCHEMICAL M ODELING Calcium concentrations and alkalinity were used to de termine the saturation state with respect to calcite. e chemical speciation model W ATEQ4F (Ball & Nord strom 1991) was used to determine saturation indices for all column samples. e saturation index is an expres sion of degree of underor over-saturation of an aqueous solution with respect to a mineral of interest (Langmuir 1997), where SI < 0 indicates an undersaturated solution capable of calcite dissolution, SI = 0 indicates the solution is at equilibrium with calcite, and SI > 0 indicates a super saturated solution from which calcite will precipitate. PHREEQ C (Parkhurst & Appelo 1999) is a reactionpath geochemical model based on thermodynamic data and equations for aqueous speciation and heterogeneous reactions among gas, liquid, and solid phases for calcula tion of the equilibrium state of a natural water of known composition with respect to various minerals. e pro gram calculates the resulting chemical composition of aqueous solutions aer two solutions mix, equilibrate with a gas phase, and equilibrate with specied miner als. In this study, PHREEQ CI (Charlton & Parkhurst 2002) was used to model a variety of scenarios, includ ing the column system in the laboratory and the cave system in situ considering both abiotic and biotic sulde oxidation. Although the chemistry of our laboratory and simulation experiments was actually suciently simpli ed that hand calculations might have been employed, testing the application of PHREEQ CI in this report al lows for building toward more complex chemical envi ronments likely to be encountered in the eld setting. Modeling the abiotic column and the abiotic cave environment e modeling results for the abiotic experimental column were compared to observed water compositions in order to assess the success of the model assumptions. W ith that success demonstrated, the subsequent model ing results for the abiotic cave system could be assumed to reasonably reect the actual cave environment. Each scenario was entered into PHREEQ CI as a series of Tab. 1: e chemical composition of prepared laboratory solutions meant to simulate groundwater within the porous limestone forma tion and the cave-conduit water. All concentrations are reported in mg/L. Na 2 SO 4 NaNO 3 NaCl, and Na 2 S salts were used to generate the various anion concentrations, and pH was adjusted with H Cl and NaO H. P repared without Ca 2+ the waters are initially greatly undersaturated with respect to calcite. Cl NO 3 SO 4 2 Na + HS pH Articial groundwater 1 4.98 0.573 16.1 12.1 0.78 5.37 Articial conduit water 2 4.97 0.630 16.2 11.0 0.0 8.5 1 Sparged with N 2 so presumed O 2 = 0.0 mg/L. 2 In equilibrium with the atmosphere, so presumed O 2 = 9 mg/L. J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS

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ACTA CARSOLOGICA 42/2-3 2013 219 steps that are outlined in Tab. 2. e amount of calcite dissolved in each reaction step was calculated using the expression for calcite dissolution rate in mmol/cm 2 /s as a function of pH, Ca 2+ and HCO 3 concentrations, and t ted parameters published by Plummer et al. (1978). To mimic the experimental columns using the geo chemical model, the calcite dissolution rate had to be combined in a calculation using an approximate surface area of the calcite in contact with water in the experimen tal system and the residence time of the groundwater in the limestone layer. e residence time of the groundwa ter in the 1.5-cm-deep limestone layer was determined in the laboratory by timing the displacement of one pore volume through the reaction bed at the experimental pumping rate. An average grain diameter of 0.75 mm was used to determine the surface area of calcite in con tact with the water in the 1.5-cm-deep limestone layer. It was assumed that the grains were packed closely enough that only of the surface area of calcite was exposed to dissolution. Although the selection of 25% of the surface area for reaction is a smaller value, probably by about half, than expected from calculation based on a geomet ric packing pattern, we chose a small value to allow for the likely hydraulic situation of incomplete access of well mixed reaction uid to the entire physical surface area. An estimate for the total amount of calcite (in number of moles) dissolved was then calculated from the reaction rate, residence time, and exposed surface area and used in the modeling scenario. A similar procedure was used to determine the total amount of calcite dissolved aer the groundwater solu tion and articial conduit water were mixed and exposed to an oxygenated atmosphere. However, the residence time was calculated for the headspace above the lime stone layer and the surface area of calcite was approxi mated for the thin top layer of calcite grains in the col umn. It was assumed that only half of the calcite grain surface area was exposed to calcite dissolution with re spect to the circulating conduit water now occupying the headspace above the calcite layer. e modeling results were then compared to the results obtained from the circulating water in the two abiotic columns. For each column and each sampling period, a saturation index was calculated based on the measured pH, alkalinity, and Ca 2+ concentration. e sat uration index obtained was then input to PHREEQ CI in order to obtain theoretical values for pH, alkalinity, and Ca 2+ concentration for the corresponding time step and column. is approach allowed for a direct comparison between the results obtained from the abiotic column and the results obtained from PHREEQ CI modeling. Based on the assumptions made in the abiotic col umn modeling scenario and previous groundwater and conduit-water data collected from No-mount Cave, the abiotic cave environment was modeled by the steps elab orated in Tab. 3. Modeling the biotic cave environment: For the biotic cave environment, several dierent scenarios were run each using slightly dierent assump tions with respect to the degree of bacterially mediated sulde oxidation. e general scenario for each model is outlined in Tab. 4, along with the variations. Currently, no stoichiometry or rate of reaction has been formulated for sulde-oxidizing bacteria; however, based on a series of assumptions, the amount of biologi cal sulde oxidation due to sulde-oxidizing bacteria Tab. 2: Steps, assumptions, and explanations for the PHREEQCI modeling simulation of the laboratory abiotic column experiment. Step 1: Articial groundwater in contact with calcite layer a) Solution 1: 1 mL articial groundwater (Tab. 1) b) Equilibrium 1: Articial groundwater was equilibrated with calcite ( i.e., SI = 0). Step 2: Resulting groundwater solution mixing with articial conduit water and equilibrating with an oxygenated atmosphere a) Solution 1 saved after step 1 was used as the groundwater solution b) Solution 2: 1000 mL articial conduit water (Tab. 1) c) Oxygenated atmosphere simulated with 9 mg/L dissolved O 2 in the articial conduit water d) Mix 1: Groundwater solution and articial conduit water mixed in a ratio of 1 mL groundwater to 1000 mL conduit water. Step 3: Mixed groundwater/conduit-water solution allowed to react with a layer of calcite and equilibrate with ambient lab P CO 2 a) Solution 3 saved after step 2 b) Reaction 2: Solution 3 was reacted with 0.2717 mol calcite, as was determined using the rate of calcite dissolution calculated from Plummer et al. (1978), residence time in the column headspace, and surface area of the calcite. c) The resulting solution was equilibrated with a P CO 2 of 10 3.3 atm (Jacobson & Wu, 2009). Step 4: Resulting circulating conduit water allowed to react with calcite layer at varying SI values based on results measured from the abiotic column a) Solution 4 saved from step 3 b) Reaction 3: calcite reacted with the solution at varying SI values. B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ...

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ACTA CARSOLOGICA 42/2-3 2013 220 was approximated. W hile there is no published reaction stoichiometry, the estimated energy obtained by the mi croorganisms during oxidation has been determined. For the partial oxidation from sulde to sulfur, a total of 177.31 kJ/mol is gained while the complete oxidation from sulde to SO 4 2 yields a much greater energy yield of 744.49 kJ/mol for sulde-oxidizing bacteria found in a geothermal, limestone well in Vulcano, Italy (Amend et al. 2004). Based on the high amounts of energy ob tained, assuming there is no limitation on bacterial growth, and assuming hydrogen sulde was the limiting reagent in all reactions, all modeled scenarios oxidized hydrogen sulde to either elemental sulfur or SO 4 2 e rst scenario, termed high-extreme, assumed the bacteria would fully oxidize sulde to SO 4 2 generat ing the maximum amount of acidic protons during com plete sulde oxidation (sum of Eqns. 1 and 2). e sec ond scenario assumed a more conservative stance where the entire sulde concentration of the natural water was only partially oxidized to sulfur and ultimately stored in the cell as sulfur granules resulting in no production of acidic protons (Eqn. 1). e third scenario modeled was based on the combination of the two previous scenarios and observations of iobacillus iothrix and Beggiatoa in the cave environment. In this scenario it was assumed that all sulde in the groundwater was oxidized to ele mental sulfur (Eqn. 1) and stored intracellularly as has been observed in cells acquired from No-mount Cave; then, it was assumed that half of the sulfur granules were subsequently oxidized to SO 4 2 (Eqn. 2). Tab. 3: Outline of modeling scenarios and assumptions used for the abiotic cave simulation. Step 1: Groundwater in contact with calcite a) Solution 1: Groundwater composition as dened by water samples obtained from wells in the surrounding limestone aquifer (Tab. 5). b) Equilibrium 1: the solution was equilibrated with a P CO 2 of 10 -2.5 atm, a typical value for a limestone aquifer (White, 1988). c) Equilibrium 2: the solution equilibrated with the calcite cave walls (SI = 0). It was assumed that the groundwater was in contact with calcite for sucient time to achieve SI=0. Step 2: Groundwater mixing with circulating conduit water a) Solution 1: solution obtained after the groundwater was equilibrated with calcite b) Solution 2: conduit water composition as dened by measurements taken in the cave system (Tab. 5). c) Mixing 1: The mixing ratios were varied to determine the amount of groundwater needed to eect calcite dissolution. Step 3: Resulting circulating conduit water dissolving calcite a) Solution 3: solution obtained after step 2 b) Reaction 1: no calcite dissolution was predicted when a negative rate of reaction was determined. Tab. 4. Assumptions for each step in the model simulations of the biotic cave environment scenarios. Step 1: Groundwater solution in contact with calcite a) Solution 1: Groundwater composition dened by water samples obtained from wells in the surrounding limestone aquifer (Tab. 5). b) Equilibrium 1: solution equilibrated with calcite. It was assumed that the water was in contact with calcite for sucient time to achieve SI = 0. Step 2: Modeling biological sulde oxidation a) Three dierent extents of microbial sulde oxidation were used: i) High extreme: complete oxidation from sulde to sulfate ii) Low extreme: half oxidation from sulde to sulfur granules iii) Mid-scenario: all sulde oxidized to sulfur; half sulfur oxidized to sulfate Step 3: Equilibrium with calcite and P CO 2 a) Equilibrium 2: resulting solution was equilibrated with calcite. It was assumed that the resulting solution was in contact with the calcite for sucient time to achieve SI = 0 b) Equilibrium 3: solution was equilibrated with a P CO 2 of 10 -2.5 atm. RESULTS e chemical composition of water collected from groundwater wells nearby, from the main boil in W eki wa Springs, and from the conduit of No-mount Cave is reported in Tab. 5. Once these chemical data were ob tained, articial solutions representing the groundwater within the limestone formation and the conduit water in No-mount Cave were prepared in the laboratory as de scribed in the Methods (Tab. 1). Monitored chemical composition of the water cir culating through contact with crushed limestone and mixing with dissolved HS over a 4-week period demon strated the feasibility of laboratory simulation of the cave J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS

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ACTA CARSOLOGICA 42/2-3 2013 221 environment and allowed an assessment of the abiotic mechanisms of calcite dissolution. Replicate laboratory columns showed calcite dissolution in the presence of abiotic sulde oxidation and dissolution due to dissolved carbonic acid. e measured HCO 3 and Ca 2+ (initial Ca 2+ = 0.0 mg/L; nal Ca 2+ = 9.82 mg/L) concentrations increased over the course of the experiment. e calcu lated saturation state of the solution increased from an initial great undersaturation toward, but not reaching, equilibrium, indicating calcite dissolution over the entire sampling period (initial SI = .0; nal SI = .2). e experimental column results were compared to the output of the theoretical modeling to determine how well the assumptions made in the model reect the physical experiment. e calculated saturation index for each sample collected from the experimental columns (from measurements of dissolved calcium and alkalin ity concentrations and pH) was used in the PHREEQ CI simulation to determine the expected pH, alkalinity, and calcium concentration. e agreement between mea sured and model-calculated parameters was generally good ( e.g., Fig. 3 for pH). e abiotic cave modeling scenario indicated no dissolution would occur at the limestone wall surface due to abiotic sulde oxidation, even when the mix ing ratio between the groundwater and the cave water was unrealistically high at a volumetric ratio of 1:1. e model does indicate, however, some dissolution within the pore space due to abiotic sulde oxidation occurring as the sulde oxidizes in the presence of the low oxygen concentrations of the groundwater. W ith respect to the biotic modeling, it can be assumed that all dissolution at the groundwater-cave interface will be due to biologic, as opposed to chemical, sulde oxidation. Four dierent scenarios were modeled to deter mine the percent of biological sulfur oxidation ver sus the amount of calcite dissolution. It was assumed that all of the dissolved sulde was oxidized to sulfur (Eqn. 1), and, therefore, only the amount of sulfur sub sequently oxidized to SO 4 2 (Eqn. 2) varied among the scenarios. e dierence in calcium concentration be tween the starting solution and the modeled resultant solution was used as a proxy for the extent of calcite dissolution. A linear regression line was t to the data and can be used to approximate the amount of calcite dissolution due to biological sulde oxidation based on the percent of oxidation of sulfur to SO 4 2 e results showed a positive, linear relationship between amount of sulde oxidation and extent of calcite dissolution (Fig. 4). Even if there was only 25% oxidation from sul fur to SO 4 2 about 99.4 mg of Ca 2+ would be dissolved into each 1 L of groundwater. Fig. 3: Comparison between column results (experimental) and modeling results (theoretical) plotted for each value of SI, calcu lated from measured dissolved calcium and alkalinity concentra tions, for each sampling period. pH, a very sensitive indicator of calcite dissolution, showed good agreement between the theoreti cal and experimental results for both experimental columns. e line represents the 1:1 relationship, i.e., perfect agreement be tween experimental and theoretical results. Fig. 4: e relationship between percent of biologic oxidation from sulfur to sulfate and its eect on calcite dissolution, as indi cated by dissolved calcium concentration. e amount of dissolu tion was modeled for four dierent scenarios of various propor tions of sulde completely oxidizing. e calcite dissolution rate law of P lummer et al. (1978) results in an asymptotic approach to equilibrium with respect to calcite as dissolved calcium con centrations increase. B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ...

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ACTA CARSOLOGICA 42/2-3 2013 222 Tab. 5: Chemical characteristics of water samples collected from Wekiwa Springs. V alues are mean one standard error with N=65 un less otherwise noted. Cells with nd indicate no data are available. Cells with bdl indicate concentrations below the detection limit. Groundwater wells 1 Main boil 2 No-mount Cave 3 Deep Mid Shallow Sonde data Temperature (C) 25.3.0 24.0.1 24.6.2 22.8.0 23.6.2 SpC (S/cm) 240 300 300 340 340 pH 8.2.0 8.1.0 8.3.0 7.4.1 7.5.1 DO (mg/L) 0.1.1 0.4.4 3.8.1 4.2.5 0.5.1 Salinity (PSS) 0.11.00 0.15.00 0.14.00 0.18.00 0.15.01 ORP (mV) 104 199 306 289 324 Field kit data Sulde (mg S/L) 4 1.3.2 bdl bdl bdl bdl 5 Alkalinity (mg CaCO 3 /L) nd nd nd 103 101 Iron (mg/L) 0.00.00 0.02.02 0.03.03 0.00.00 0.00.00 Cations (mg/L) Ca 2+ 31.0 35.2 27.2 39.7 45.3.3 Na + 9.3 10.1 20.4 9.3 9.5.2 Mg 2+ 8.6 9.6 9.7 11.3 10.2.2 K + 0.4 0.6 10.4 1.6 1.4.0 NH 4 + 0.0 0.0 0.0 0.0 0.0.0 Anions (mg/L) Cl 5.3 8.3 6.6 18.0.7 14.2.3 SO 4 2 15.5 18.6 29.3 22.1.8 18.6.3 NO 3 0.1 0.0 0.3 0.5.1 0.7.1 PO 4 3 1.2 1.2 0.6 0.3.2 0.4.0 1 Sampled January 2005, N=3 per well for sonde and eld kit data; N=1 per well for cations and anions. 2 Sampled July 2007, N=3 per well for all parameters. 3 Triplicate samples for each of ve stations on each of the ve dates (N=65) except for sonde data where N=25 (one sample at each station on each date). 4 e detection limit is 0.05 mg S/L. 5 No sulde was detected in conduit water. e limestone formation matrix water was sampled 15 cm deep into the cave wall rock one time within No-mount Cave at three dierent locations (~4, 6, and 10 m linear penetration distance) and yielded sulde concentrations of 0.08, 0.00, and 0.09 mg/L, respectively. DISCUSSION e abiotic-column modeling scenario was mainly con ducted to determine if the assumptions made in the modeling could accurately reect the evolving chemical composition of subterranean water (Fig. 1) mimicked by a physical experiment (Fig. 2). roughout the 4-week laboratory experiment, the modeling results seemed to accurately reect the results obtained in the in vitro col umns. Based on the relatively good agreement between the experimental column results and the theoretical modeling results, we concluded that the model could be used to accurately model the cave environment. Our desire to construct laboratory columns with biological activity failed when multiple attempts to grow suldeoxidizing bacteria from three dierent microbial mat samples obtained from No-mount Cave were each in troduced into three dierent sulde-containing growth media using standard culturing methods resulted in no growth of sulfur-granule-containing bacteria. e abiotic cave modeling results indicate dissolu tion in the pore space, potentially leading to increased porosity within the limestone bedrock, but indicates no dissolution on the limestone cave-wall surface. e oxy gen concentrations in the circulating cave water maybe too low to support enough abiotic sulde oxidation to J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS

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ACTA CARSOLOGICA 42/2-3 2013 223 promote additional dissolution. In addition, the dilu tion factor between the incoming groundwater and the already circulating cave-water may be too high for the high sulde and, therefore, high sulde-oxidizing-poten tial groundwater to make much of an impact on the cal cite-dissolving capabilities of the circulating cave water. A better estimate for abiotic oxidation could be obtained if the ux of groundwater through the entire cave-wall system and the total amount of circulating cave-water was known for the cave environment; however, the mod eling, despite signicantly increasing the ratio between groundwater and circulating cave water, indicates that the amount of calcite dissolution due to abiotic sulde oxidation along the cave-wall surface will be small. Based on the results obtained for the abiotic cave modeling, it is assumed that all oxidation on the cave wall-surface for the biotic cave modeling scenario was due to biologically mediated sulde oxidation. Signi cant dissolution due to sulde oxidation was obtained when the bacteria were modeled to oxidize all HS to SO 4 2 e high amount of dissolution is expected based on the greater production of protons in the complete oxidation of sulde (sum of Eqns. 1 and 2). From eld observations, however, it seems unlikely that complete (100% scenario) sulde oxidation is occurring. Intracel lular sulfur granules were noted in the collected bacte rial mats (Franklin et al. 2005), indicating that not all the dissolved sulde is fully oxidized to SO 4 2 On the other end of the extreme, if the microor ganisms could only oxidize the sulde to sulfur (Eqn. 1) and never produce SO 4 2 (Eqn. 2), the eect on extent of calcite dissolution was notable. For this scenario, no cal cite was dissolved in the geochemical model simulations. is extreme, as was the case with the full sulde oxida tion, is not expected based on eld observations. e pH of water in the micro-environments surrounding sulfuroxidizing bacterial mats has been measured as low as 0 (Hose et al. 2000; Sarbu et al. 1996). e low pH readings are attributed to the generation of acid during the biotic oxidation of sulde at least partially to sulfate (Eqn. 2). Consideration of these two unlikely extremes leads us to expect that the amount of SO 4 2 produced by suldeoxidizing bacteria probably lies somewhere in between. From our various scenarios for simulation modeling, the assumption that all sulde is oxidized to sulfur and subsequently only half of the sulfur is oxidized to SO 4 2 appears to be a reasonable approximation of eld con ditions. is scenario yields a concentration of 158 mg dissolved Ca 2+ per L of groundwater in the limestone formation. is estimate assumes the oxidized water is in contact with calcite long enough to reach equilibrium. An additional factor is the note that Acidithiobacillus has been found in the mat community as well. is organism oxidizes HS or elemental S to SO 4 2 with the concomi tant production of protons, i.e. the sum of Eqn. 1 and Eqn. 2. In this specic modeling scenario, the mixing ra tio was assumed to be 1:1000 for the groundwater and circulating conduit water, however, this assumption has not yet been corroborated with eld data. More informa tion is needed to determine the groundwater ux into the cave and the amount of circulating conduit water to more accurately predict the mixing ratio. A dier ent mixing ratio could result in more dissolution if the groundwater ux is greater than assumed here, because it is the groundwater discharging from the porous lime stone formation that contained the sulde whose oxida tion generates acidity. Conversely, there is the potential for less dissolution if the conduit water is present in large enough amounts to suciently dilute the acidic ground water. In addition, the model also assumes the acidic groundwater is in contact with the limestone wall long enough for calcite to reach equilibrium (SI = 0). is may also be a false assumption, if the circulating cave water is moving at a rapid enough velocity to quickly and su ciently mix the groundwater with the circulating conduit water. It may be that reaching equilibrium calcite solu bility only occurs in slower-ow environments such as in pore spaces or at the microenvironment at the base of a thick biolm. In contrast to a partial approach to equi librium calcite dissolution, maximum biospeleogenesis will derive from sucient contact time to allow for large amounts of calcite dissolution leading to equilibrium. More information on the physical cave environment and the structure of biolms is needed to determine the ef fect that pore spaces and lamentous biolms have on the residence time of the acidic water in contact with the conduit wall. CONCLUSION e results of this study indicate a strong potential for increased calcite dissolution, and ultimately cave en largement due to sulde-oxidizing bacteria. Based on the PHREEQ CI modeling results, up to 229 mg of calcium B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ...

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ACTA CARSOLOGICA 42/2-3 2013 224 REFERENCES Amend, J.P., Rogers, K.L. & D.R. Meyer-Dombard, 2004: Microbially mediated sulfur-redox: Energetics in marine hydrothermal vent systems.In: Amend, J.P. et al. (eds.), Sulfur Biogeochemistry: past and present, Geological Society of America, pp. 17. Boulder, CO. American Public Health Association, 1995: Standard methods for the examination of water and wastewa ter. American Public Health Association, W ashing ton, D. C. Ball, J.W & D.K. Nordstrom, 1991: User's manual for WATEQ4F, with revised thermodynamic data base and test cases for calculation of major, trace, and redox elements in natural waters.U.S. Geological Survey, Open-File Report 91. Bottrell, S.H., Smart, P.L., W hitaker, F. & R. Raiswell, 1991: Geochemistry and isotope systematics of sul phur in the mixing zone of bahamian blue holes.Applied Geochemistry, 6, 97. Charlton, S.R. & D.L. Parkhurst, 2002: PHREEQCI-A graphical user interface to the geochemical model PHREEQC.U.S. Geological Survey, Fact Sheet: FS031. Cline, J.D., 1969: Spectrophotometric determination of hydrogen sulde in natural waters.Limnology and Oceanography, 14, 454. Engel, A.S. & K.W Randall, 2011: Experimental Evi dence for Microbially Mediated Carbonate Disso lution from the Saline W ater Zone of the Edwards Aquifer, Central Texas.Geomicrobiology Journal, 28, 4, 313. 10.1080/01490451.2010.500197. Erlich, H.L. & D.K. Newman, 2009: Geomicrobiology CRC Press, Boca Raton, FL. Ferguson, G.E., Lingham, C.W ., Love, S.K. & R.O. Ver non, 1947: Springs of Florida. Florida Geological Survey, Bulletin: 31. Franklin, R.B., Campbell, A.H., Higgins, C.B., Barker, M.K. & B.L. Brown, 2011: Enumerating bacterial cells on bioadhesive coated slides.Journal of Mi crobiological Methods, 87, 154. Franklin, R.B., Tysall, T.N., Gianotti, A. & A.L. Mills, 2005: Geomicrobiology of phreatic caves associ ated with central Florida springs, Spring Meeting, American Geophysical Union, New Orleans, LA. Hill, R.S. & R.B. Franklin, 2012: Microbial diversity: A spatial study of microbial community assemblages in the Floridian Aquifer, 97 th Annual Meeting of the Ecological Society of America, Portland, OR. Hose, L.D., Palmer, A.N., Palmer, M.V., Northup, D.E., Boston, P.J. & H.R. DuChene, 2000: Microbiol ogy and geochemistry in a hydrogen-sulphide-rich karst environment.Chemical Geology, 169, 3, 399. 10.1016/s0009(00)00217. Hubbard, D.A., Jr., Herman, J.S. & P.E. Bell, 1986: e role of sulde oxidation in the genesis of Cesspool Cave, Virginia, USA, 9th International Congress of Speleology, Barcelona, Spain, 255. Jacobson, A.D. & L. W u, 2009: Microbial dissolution of calcite at T= 28C and ambient pCO 2 .Geochimica et Cosmochimica Acta, 73, 2314. 10.1016/j. gca.2009.01.020. could be dissolved per square centimeter of limestone surface per liter of water, presuming complete oxidation of all available dissolved sulde. e more conservative estimate, based on a presumed partial oxidation of the dissolved sulde, of 158 mg of dissolved calcium per square centimeter per liter of water is expected. Either way, the presence of sulde-oxidizing bacteria in a lime stone aquifer would indicate accelerated calcite dissolu tion. e study also sets groundwork for future research in the actual rate and stoichiometry for biotic sulde oxi dation, subsequent rate of calcite dissolution, and rate of cave enlargement. ACKNO WLEDGEMENTS e authors thank the Cambrian Foundation for local coordination and intellectual and logistical support and Florida Department of Environmental Protection for ac cess to the springs which are located in State Parks. Fund ing from the University of Virginia to AGH and from the National Science Foundation to RBF (DEB 0920398) is gratefully acknowledged. J ANET S. H ERMAN A LE X ANDRIA G. H OUNSHELL R IMA B. F RANKLIN & A ARON L. M ILLS

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ACTA CARSOLOGICA 42/2-3 2013 225 Langmuir, D., 1997: Aqueous Environmental Geochem istry. Prentice-Hall, Inc., Upper Saddle River, NJ, 600. Macalady, J.L., Lyon, E.H., Koman, B., Albertson, L.K., Meyer, K., Galdenzi, S. & S. Mariani, 2006: Dominant microbial populations in limestonecorroding stream biolms, Frasassi Cave system, Italy.Applied and Environmental Microbiology, 72, 5596. Mattison, R.G., Abbiati, M., Dando, P.R., Fitzsimons, M.F., Pratt, S.M., Southward, A.J. & E.C. South ward, 1998: Chemoautotrophic microbial mats in submarine caves with hydrothermal sulphidic springs at Cape Palinuro, Italy.Microbial Ecology, 35, 1, 58. Miller, J.A. 1986: H ydrogeologic framework of the Flori dan aquifer system in Florida and in parts of Geor gia, Alabama, and South Carolina. U.S. Geological Survey Professional Paper: 1403B. Miller, J.A. 1990: Ground Water Atlas of the United States: Alabama, Florida, Georgia, and South Carolina. U.S. Geological Survey HA 730G. Otte, M.L. & J.T. Morris, 1994: Dimethylsulphoniopro pionate (DMSP) in Spartina Alterniora Loisel.Aquatic Botany, 48, 3, 239. Parkhurst, D.L. & C.A.J. Appelo, 1999: Users guide to PHREEQC ( V ersion 2)-A computer program for spe ciation, batch-reaction, one-dimensional transport, and inverse geochemical calculations. U.S. Geologi cal Survey, W ater-Resources Investigations Report: 99. Plummer, L.N., W igley, T.M.L. & D.L. Parkhurst, 1978: Kinetics of calcite dissolution in CO 2 -water systems at 5C to 60C and 0.0 to 1.0 atm CO 2 .American Journal of Science, 278, 2, 179. Rounds, S.A., 2006: Alkalinity and acid neutralizing ca pacity (ver. 3.0), U.S. Geological Survey Techniques of W ater-Resources Investigations, book 9, chap. A6, sec. 6.6. U.S. Geological Survey. Sarbu, S.M., Kane, T.C. & Kinkle, B.K., 1996: A chemoau totrophically based cave ecosystem.Science, 272, 5270, 1953. Scott, T.M., 1991: A Geological overview of Florida.In: Scott, T.M. et al. (eds.), Florida's Ground Wa ter Quality Monitoring P rogramH ydrogeological Framework, pp. 5. Florida Geological Survey Special Publication 32. Vlasceanu, L., Sarbu, S.M., Engel, A.S. & B.K. Kinkle, 2000: Acidic cave-wall biolms located in the Fras assi Gorge, Italy.Geomicrobiology Journal, 17, 2, 125. W hite, W .B., 1988: Geomorphology and H ydrology of Karst Terrains. Oxford University Press, Oxford, 464. B IOLOGICAL C ONTROL ON A CID G ENERATION AT THE C ONDUIT B EDROCK B OUNDARY IN S UBMERGED C AVES : ...



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AN APPROACH FOR COLLECTION OF NEARFIELD GROUND W ATER SAMPLES IN SUBMERGED LIMESTONE CAVERNS P RISTOP ZA ZBIRANJE VZORCEV PODZEMNE VODE V POTOPLJENIH APNEN ASTIH JAMAH Aaron L. MILLS 1* Terrence N T YSALL 2 & Janet S. H ERMAN 1 Izvleek UDK 551.444:543.3 Aaron L. Mills, Terrence N Tysall & Janet S. Herman: Pristop za zbiranje vzorcev podzemne vode v potopljenih apnenastih jamah Stene potopljenih jam, ki napajajo oridske izvire, so ve likokrat pokrite z debelimi oblogami vlaknastih bakterij. Te lahko oksidirajo reducirano veplo v podtalnici, ki pronica iz porozne matrice v jamski kanal. Da bi doloili spreminjanje kemijske sestave vode, ko ta prehaja skozi mikrobsko oblogo, smo v vrtino, izvrtano v jamsko steno, namestili enostavno napravo, narejeno iz perforirane cevi za zaito vrtin ter gu mijastega zapiraa s prilagodljivimi odprtinami za vzorenje. Vzorevalnik smo pritrdili z epoksi lepilom. Primerjali smo koncentracije smo anionov v vodi vzoreni z vzorevalnikom in vodi iz jamskega kanala. Koncentracija veine anionov, npr. Cl NO 3 and PO 4 3 je v kanalu nekoliko veja, kot v porozni matrici. Sulde smo merili le v vodi porozne matrice. Kon centracija SO 4 2 je v prevodniku je bila 22 g/L, v vzorevalniku pa 11 g /L, kar kae na to, da je oksidacija vepla pomemben proces v bakterijskih oblogah, ki prekrivajo apnenaste stene v jamah. Vzorevalnik je uporaben tudi za merjenje pretoka iz lokalne matine kamnine v prevodnik. Kljune besede: oksidacija vepla, bakterije, raztapljanje s kislino, podzemna voda, kemija. 1 Laboratories of Microbial Ecology and Aqueous Geochemistry, Department of Environmental Sciences, University of Virginia, P.O. Box 400123, Charlottesville, VA 22904-4123 2 e Cambrian Foundation, 1572 Lawndale Circle, Ste. A, W inter Park, FL 32792 *Corresponding Author, e-mail: amills@virginia.edu Received/Prejeto: 1.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 227, POSTOJNA 2013 Abstract UDC 551.444:543.3 Aaron L. Mills, Terrence N Tysall & Janet S. Herman: An ap proach for collection of neareld groundwater samples in sub merged limestone caverns W alls of submerged caves feeding Florida springs are oen lined with a heavy mat of lamentous bacteria, many of which are able to oxidize reduced sulfur in groundwater migrating from the porous bedrock into the cave conduit. To determine changes in water chemistry as water passes through the mi crobial mat, a simple device made from standard well screen and sealed with a rubber stopper and controllable vents was installed in a hole drilled in the wall of the cave passage. e sampler was sealed in place with marine epoxy. W e measured anions in water from the sampler and from the water-lled conduit taken just outside the sampler. Most anions measured viz ., Cl NO 3 and PO 4 3 increased slightly between the matrix and conduit waters. However, traces of sulde were measured in the water from the rock matrix, but not in the conduit. SO 4 2 concentrations in the conduit were about twice that measured in the water from the sampler, about 22 and 11 mg SO 4 2 L respectively, providing further evidence that sulfur oxidation is an important process in the bacterial mats attached to the lime stone surfaces in these caves. An additional use of the sampling device is to measure discharge from the local bedrock into the cave conduit. Keywords: sulfur oxidation, bacteria, acid dissolution, ground water chemistry.

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ACTA CARSOLOGICA 42/2-3 2013 228 W ater enriched in reduced sulfur is common in many karst terranes, and communities of microorganisms found in the aphotic portion of subaerial and submerged caves in such locations are oen rich in sulfur-oxidizing, autotrophic bacteria (Brigmon et al. 1994; Chen et al. 2009; Engel et al. 2003; Franklin et al. 2005b; Sarbu et al. 1996; Vlasceanu et al. 2000). Such microbes gain energy from the oxidation of reduced sulfur to intermediate or fully oxidized species. e reactions are well dened (Er lich 1996; Stumm & Morgan 1996), and the products are commonly found in cave waters fed by sulfur-containing groundwater. Reduced sulfur (H 2 S, HS ) is common in ground water in central Florida (Sacks 1996). In many sub merged caves, including the two that are examined in the present report, the conduits are covered with a mat of lamentous microorganisms, and the mats usually contain a large proportion of cells easily identiable as sulfur oxidizers by the presence of visible grains of sul fur within the laments (Franklin et al. 2005b). e most likely scenario to remove sulde from the groundwater involves the oxidation of sulde to S 0 and then to SO 4 2 as the water passes through the mat. Protons resulting from the oxidation reaction culminating in SO 4 2 might accel erate the dissolution of the host limestone, such that spe leogenesis occurs faster where there is prolic growth of sulfur-oxidizing microbes (Engel & Randall, 2011; Engel et al. 2004; Franklin et al. 2005b; Vlasceanu et al. 2000). e microbiological acceleration of conduit enlargement is termed biospeleogenesis (Barton & Luiszer, 2005). To help dene that mechanism, we developed an approach that allows comparison of water from within the conduit to that within the bedrock immediately proximal to the sampling site. To determine the extent of chemical change in the water as it discharges from the limestone formation prox imal to the cave conduit as a means to inform geochemi cal modeling of sulfuric acidogenesis and carbonate dis solution, we describe a simple approach for collecting samples of water from within the formation behind the bacterial mat. W e also describe a means of using the in stalled sampler to estimate discharge of water across the face of the wall, useful in quantifying the extent of geo chemical processes constituting biospeleogenesis. W e re port the results of analysis of samples collected with the device(s) to show the kinds of information that can be gained from these samplers. INTRODUCTION AARON L. MILLS, TERRENCE N T YSALL & J ANET S. HERMAN METHODS SAMPLER DESIGN AND CONSTRUCTION. e sampling device is simply a cased well with appropri ate sampling ports that is installed in the wall of the sub merged cave. e sampler itself was constructed from a length of 1-inch diameter (3.3 cm OD 2.4 cm ID) PVC well screen with slots (0.01 inches [0.254 mm] wide). is screen is standard for monitoring wells, and it can be purchased at any driller supply outlet. For our applica tion, sections of screen were cut about 14 cm in length (See Fig. 1.). One end of the screen section was closed with a rubber stopper that had two 6-inch-long, 14-gauge syringe needles inserted nearly to the hub. One of the nee dles was cut o so that it opened at the inner face of the stopper within the chamber. Because the samplers were to be installed into the face of a rock wall and sealed, one of the needle hubs was bent slightly to allow access of the Fig. 1: Design of the formationwater sampler.

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ACTA CARSOLOGICA 42/2-3 2013 229 A N APPROACH FOR COLLECTION OF NEARFIELD GROUND W ATER SAMPLES IN SUBMERGED LIMESTONE CAVERNS hubs by multiple syringes. Additionally, the shorter hub was marked with a cable tie so that divers could tell the short and long needles apart. e stopper was sealed into the well screen with Plumbing Goop (Eclectic Products, Inc., Eugene, OR, available at most hardware and home improvement stores). Other sealants would likely also be satisfactory, but we have found Goop to be a good adhe sive and excellent sealant for underwater applications. SAMPLER INSTALLATION Samples of conduit water and water from within the cave walls were taken at two freshwater springs in central Florida, W ekiwa Spring (28.711910N, 81.460214W ) and DeLeon Spring (29.134246N, -81.362766W ) by a dive team from e Cambrian Foundation. e springs are unconnected, but both discharge to the St. Johns Riv er. At W ekiwa Spring, three samplers were installed in No-Mount Cave at increasing depth in the cave conduit: 3.7 m, 6.1 m, and 10.1 m below the surface of the water in the spring pool. At DeLeon Spring, a single sampler was deployed about 20 m below the free water surface in the spring pool. e particular points at which the samplers were placed had visible bacterial mat present, a small area of which was scraped away to make space for the sampler chamber to be installed. Placement of samplers into the wall of each sub merged cave was done by experienced cave divers. A pneumatic drill was attached to a scuba tank, and the drill was tted with a 1-inch masonry bit of sucient length to bore a hole about 15 cm into the rock. e divers carried the drill and tank to the proper location, bored a hole in the wall, and tted the sampler snugly into the opening until the top of the well-screen section was about ush with the rock-wall surface. e sampler was then sealed in place with a marine epoxy suitable for underwater work. is installation le the two nee dle hubs protruding from the wall. Between samplings, groundwater discharging from the wall passes freely through the sampler chamber, through the needles and into the cave conduit. W ATER SAMPLING e chamber created by drilling the hole and inserting the sampler has a void volume of about 125 mL. Removal of water from the interior of the embedded samplers was done by displacement of the water with sterile distilled water (SD W ). A 60-mL syringe lled with SD W and closed with a stopcock was carried by the divers who col lected samples. e syringe was connected to the longer of the two needles inserted into the sampler through the rubber stopper (Fig. 2). An empty 10-mL syringe with a closed stopcock was then attached to the short needle. e stopcocks were opened and SD W from the larger syringe was slowly expressed into the back of the sampler forcing the water that the sampler had contained into the smaller syringe. It is oen necessary for a sec ond diver to withdraw the plunger of the smaller syringe slowly to facilitate entry of the water into the syringe. W hen the smaller syringe was lled, the stopcock was closed and the syringe was placed into a bag for return to the surface. en a second 10-mL syringe was attached to the needle hub. is syringe contained 5.0 mL of 10 mM Zn acetate to act as a xative to preserve dissolved sul de for later analysis (Cline 1969). W hen the syringe was full to the 10-mL mark, the stopcock was closed and the assembly was detached from the sampler and placed in the bag for return to the surface. e 60-mL syringe was then closed, and it was also placed in the bag with the others. Additional water samples from the cave con duit immediately adjacent to the sampler were collected in sterile, plastic 50-mL screw-cap centrifuge tubes that were then sealed and placed in the bag with the other samples for return to the surface. On the surface, the samples were placed in coolers with ice or synthetic cold packs and returned to the laboratory for analysis. LABORATORY ANALYSES W ater from the cave conduit and from within the porous limestone was returned to the laboratory in Virginia for analysis. Shipments were sent by overnight courier and were kept cold with frozen, synthetic cold packs. For this study, samples were analyzed for anion concentra tion. e anions Cl SO 4 2 NO 3 and PO 4 3 were ana lyzed by ion chromatography (Dionex ICS2100 equipped with an AS4 anion column). Total dissolved sulde (i.e., sum of H 2 S, HS and S 2 ) concentration was determined by colorimetry using Clines reagent (Cline, 1969) as de scribed by Otte and Morris (1994). Fig. 2: Sampler as congured for sample withdrawal. Sterile dis tilled water is expressed from the larger syringe into the chamber, displacing water collected in that chamber into the smaller syringe. e diagram is exploded to show the various parts used in ef fecting the sample transfer.

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ACTA CARSOLOGICA 42/2-3 2013 230 DETERMINATION OF SPECIFIC DISCHARGE rSEEPAGEf Given that the samplers do not impede ow of water from the wall of the cave into the cave opening, rather they merely channel the ow through the needle hubs, the samplers can be used to determine specic discharge from the walls into the cave (i.e., seepage). e general approach employed here was developed in the late 1970s (Lee 1977; Lock & John, 1978) and has been used in lakes, sandy-bottomed streams, and estuarine and coast al locations to measure inseepage (Flewelling et al. 2012; Holly et al. 2003). To collect the water passing through the needle hubs, a 60-mL syringe was cut to obtain a sec tion of the syringe barrel about 3.5 cm long, including the male Luer-lock tip. A latex condom was then placed over the larger end of the syringe tip and secured in place with two short nylon cable ties placed around the unit such that the clasps were on opposite sides of the bar rel of the syringe tip (Fig. 3). e use of two cable ties is essential to obtain a good seal of the condom to the syringe barrel. Use of several dierent chemical sealants always resulted in disintegration of the condom where it contacted the sealant material. A stopcock (Cole Parmer, YO-30600-25) was attached to the syringe tip, the air ex pelled from the condom and the stopcock closed, and the apparatus taken to the sampler location by a diver. e unit was attached to the short-needle hub in the sampler. A stopcock was also placed on the long-needle hub and closed to prevent water from passing through that open ing (Fig. 4). e stopcock on the discharge-collecting-condom device was then opened and the time recorded. Aer a period of time (variable, up to about 2 hours for this study), the stopcock was closed, the time was recorded, the device was detached and brought to the surface, and the water was expelled into a small graduated cylinder to measure its volume. Specic discharge (q) was determined as the total discharge (Q mL of water collected / total collection time, divided by the cross sectional area of the sampler calculated based on the inner diameter of the well screen being 2.4 cm). Fig. 3: Condom attached to a 60-mL syringe tip as used to collect seepage passing through the sampler. Fig. 4: Deployed sampler tted with condom attachment for measurement of discharge of groundwater across the face of the cave wall. RESULTS ANIONS Most of the anions measured by ion chromatography were slightly higher in the water collected from the cave con duit as compared to the water collected from within the cave wall (Tab. 1). At both W ekiwa and DeLeon Springs, Cl concentrations in the cave conduit were higher than those in water from within the wall, and Cl concentra tions in all samples from DeLeon Springs were higher than all samples from W ekiwa Springs. Nitrate was also higher in the cave than in the wall, with dierences on the order of 1-3 mg L Phosphate was oen not seen in one or both samples, but the chromatographic system used for the analysis was not optimized for PO 4 3 such that low concentrations (<1 mg PO 4 3 L ) were oen not detected (even in standards). Sulfate was also higher in the cave conduit than in the cave wall, but the absolute dierences seen were substantially larger than any of the other anions. For all of the sample pairs, the water gained AARON L. MILLS, TERRENCE N T YSALL & J ANET S. HERMAN

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ACTA CARSOLOGICA 42/2-3 2013 231 8-12 mg SO 4 2 L (2.7-4 mg S L ) between the cave wall and the conduit. Sulde concentrations were very low in all samples. Formation water typically contained 0.080.09 mg sulde-S L although at the mid-depth station in W ekiwa Spring, sulde was below detection. In neither W ekiwa Spring nor DeLeon Spring was sulde detected in water from the cave conduit. Previous sampling of the conduit water feeding both springs has never shown sulde (R. B. Franklin, unpublished data and (Franklin et al. 2005a; 2005b)). Presumably, the sulde was oxidized by the sulfur-oxidizing bacteria that dominate the micro bial mats found lining the walls within the caves that feed both these springs. Identication of the organisms pro vided by Franklin et al. (2005a; 2005b) conrmed that the organisms were, indeed, sulfur oxidizers. Tab. 1: Concentration of selected ions in water from Wekiwa Springs and DeLeon Springs caves ( J uly 2012).V alues in parentheses indi cate the depth of the sampler below the water surface. bdl=below detection limit. Detection limits were: Cl 0.1 mg L ; sulde, 0.05 mg L 1 ; SO 4 2 0.1 mg L ; NO 3 0.1 mg L ; PO 4 3 0.1 mg L 1 ; dissolved O 2 0.1 mg L Location Analyte Wall (mg L 1 ) Conduit (mg L 1 ) Cl 97.5 108.1 Sulde-S 0.09 bdl Deleon Spring (20 m) SO 4 2 11.0 23.0 NO 3 0.8 3.5 PO 4 3 bdl bdl Dissolved O 2 Not measured 0.41 pH 8.0 7.4 Cl 15.8 17.8 Sulde-S 0.08 bdl Wekiwa Deep (10.1 m) SO 4 2 11.5 21.4 NO 3 2.3 3.6 PO 4 3 bdl 0.4 Dissolved O 2 0.12 0.63 pH 8.2 7.4 Cl 16.5 17.9 Sulde-S bdl bdl Wekiwa Mid (6.1 m) SO 4 2 14.2 22.6 NO 3 1.5 3.5 PO 4 3 bdl bdl Dissolved O 2 0.43 0.30 pH 8.1 7.4 Cl 16.6 17.9 Sulde-S 0.09 bdl Wekiwa Shallow (3.7 m) SO 4 2 14.1 22.4 NO 3 2.0 3.5 PO 4 3 bdl bdl Dissolved O 2 3.6 0.33 pH 8.2 7.4 Tab. 2: Groundwater discharge into Wekiwa and Deleon Springs. J uly 2012. Reported values represent discharge across the cross sec tional area of the sampler, viz., 4.52 cm 2 Spring Sampling Depth Sampling duration (min) Water Volume recovered (mL) Q (m 3 min 1 ) q (Q/A) (m min -1 ) 3.7 m 83 1.8 2.2 10 8 4.8 10 9 Wekiwa 6.1 m 25 5.9 6.9 10 8 1.5 10 8 10.1 m 60 21.0 3.5 10 7 7.7 10 8 DeLeon 20 m 145 23.1 1.6 10 7 3.5 10 8 A N APPROACH FOR COLLECTION OF NEARFIELD GROUND W ATER SAMPLES IN SUBMERGED LIMESTONE CAVERNS

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ACTA CARSOLOGICA 42/2-3 2013 232 SPECIFIC DISCHARGE Discharge from the limestone bedrock through the cave walls to the cave conduit was measureable in relatively short time periods (<1 to ca. 2.5 hr.) at every sampling location (Tab. 2). In W ekiwa Spring, where sampling was done at dierent depths, discharge increased with depth. Specic discharge at the 10.1-m sampling point was over an order of magnitude greater than that at 3.7 m below the water surface. DISCUSSION Because of limitations in the number of samplers that could be deployed within the permit granted by the Florida DEP, the results reported for both chemistry and groundwater discharge represent only a single sample from each sampling location. us, statistical evaluation of the data is not possible. e data are entirely consistent with those reported for conduit waters taken at several depths in W ekiwa and DeLeon Spring and from nearby wells by Franklin et al. (2005a; 2005b). e samplers described here enabled acquisition of water samples from the bedrock limestone formation prior to its passing through the microbial mat that cov ers large expanses of the cave walls in the two freshwater submerged caves in which they were deployed. Addition ally, the samplers proved useful in obtaining estimates of discharge into the cave from the surrounding porous bedrock. Lower concentrations of SO 4 2 within the host limestone compared to higher concentrations in the cave conduit suggested that SO 4 2 was being formed between the formation and the conduit, presumably in the mi crobial mat where sulfur-oxidizing bacteria are known to occur. at interpretation is also consistent with the presence of dissolved sulde in the water from the for mation and its absence in the conduit. e mass of sul de oxidized did not balance the amount of SO 4 2 formed, although when expressed as S, the dierences are not as great as the numbers in Tab. 2 suggest. e water lost approximately 0.08 mg sulde-S L -1 and gained around 3.2 mg SO 4 2 S L (based on averaging all samples). e failure to close the mass balance cannot be explained at this point, but loss of sulde from the samples prior to analysis, even though preserved with Zn 2+ is a possibil ity, even though the results for the conduit for both sul de and SO 4 2 agree favorably with those obtained by R. B. Franklin (Herman et al. 2013). Nevertheless, trends in losses of sulde and gains in SO 4 2 are in agreement and suggest the microbial mat is, indeed, generating SO 4 2 and associated protons from the reduced S emanating from the cave wall. Further use of the deployed samplers, and installation of some additional samplers in other posi tions, may help decipher the current imbalance in sulfur. Geochemical modeling of the system supports the idea that sulde oxidation in this system can, indeed, account for substantial calcite dissolution (Herman et al. 2013). us, we have demonstrated that use of the device to re cover samples for chemical analysis provided results that were consistent with the expectations of similar concen trations of some anions (Cl NO 3 and PO 4 3 ) in the wall and conduit, and substantial changes in concentration for others of interest (SO 4 2 and dissolved sulde). Use of the samplers to measure seepage was an af terthought, and use of condoms to collect the water from the walls of the caves to quantify discharge into the con duit was based largely on our previous experience mea suring inseepage in lakes and sandy-bottomed streams (Bruckner et al. 1984; Bruckner et al. 1989; Flewelling et al. 2012; Lehman & Mills 1994; McIntire et al. 1987; McIntire et al. 1988). e use of condoms in measure ment of discharge of water from porous media into open waters such as stream channels has been debated over the years, and there is not a clear consensus of opinion as to the suitability for seepage measurements. Many of the issues with condoms revolve around use for low-vol umes and long time periods (e.g., Fellows & Brezonik, 1980; Schincariol & McNeil 2002). Problems such as de terioration of the condom with time or relaxation over long times as suggested by Schincariol & McNeil (2002) are not issues with short deployment times. In a study comparing several types of bags with a dye-displacement meter that has no bag and therefore none of the prob lems associated with bags, Koopmans and Berg (2011) found that larger more rigid bags (3.8 L zipper bags or 3.8-L twist-tie bags) similar to those used by other in vestigators exerted a positive pressure (expressed as a head dierential, h) that increased with increasing col lection volume. e range of h was from around 0.1 to 70 mm as the bags lled. us, the bags exerted a strong resistance to ow. W hen condoms were used as collec tion bags, the head gradient was very slightly negative at around .1 mm, but the dierential stayed constant until the condom reached its non-elastic limit at about 120 mL. Aer that, the head gradient became strongly positive, as would be expected as the latex was forced to stretch. Koopmans and Berg (2011) concluded that seep AARON L. MILLS, TERRENCE N T YSALL & J ANET S. HERMAN

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ACTA CARSOLOGICA 42/2-3 2013 233 ACKNO WLEDGEMENTS e authors thank their project partners, the Cambrian Foundation for local coordination and intellectual and logistical support, and Florida DEP for permitted access to the springs which are located in State Parks. A number of individuals from the Cambrian Foundation played major roles in the conduct of this work, and the authors thank Renee Power and Kris Shannon for diving into the springs along with Tysall. W e thank Amy Giannotti for overall coordination of the project in Florida, and for act ing as the intermediary with Florida DEP. Special thanks go to Marissa W illiams for campaign management; bet ter organization and support cannot be imagined. e project was motivated by a cooperative eort with Rima Franklin at Virginia Commonwealth University who pro vided initial contact with the Cambrian Foundation and also provided initial logistical support through a grant (DEB 0920398) from the National Science Foundation. age meters that employ bags of any type are less sensitive than their dye-displacement meter, but that condoms represented a small source of error in measurement of smaller volumes (<100 mL) that compared favorably with the bagless version over short time periods (a few hours). In addition to the dye-displacement meters of Koopmans and Berg (2011), other types of meters have been developed that also use open ow paths (e.g., Paulsen et al. 2001; Rosenberry & Morin 2004; Sholko vitz et al. 2003), but the expense and complexity of their operation make them impractical for deployment in numbers in submerged cave situations. An obvious improvement in the device design for use as a discharge meter would be to employ a larger di ameter in order to integrate a greater wall area. Drilling larger holes into the rock in these caves would be di cult at best, might destabilize materials of low cohesive ness, and would likely not be permitted by landowners or stewards in many diveable caves of interest. Although we computed the specic discharge using the inner diameter of the sampler, the lower resistance to ow likely causes some greater volume of water to ow through the hollow chamber than would ow across an equivalent cross section of the porous limestone. us, the eective diameter of the sampler is probably slight ly larger than the actual diameter. Because of the high permeability of the porous rock and the relatively low surface area of the drilled chamber, we believe that dif ference to be small and the true specic discharge to be only very slightly less than that computed with the diam eter of the sampler chamber. A question that arose during the initial installation was how long to wait before returning to collect sample. Based on the observed discharge rates at W ekiwa and DeLeon Springs, and an assumed chamber volume of 125 mL, the sample chambers would require between 6 hr (W ekiwa deep sampler) and 96 hr (W ekiwa shal low sampler) to replace the volume of the sampler with groundwater from the formation (T = Volume / seepage rate), In W ekiwa and DeLeon Springs, the separation of installation and rst sampling by a few days to a week should have resulted in complete turnover of all wa ter inside the sample chamber and, therefore, collected samples are representative of the groundwater proximal to the cave wall surface. Because distilled or deionized water is injected into the sampler chamber to displace the native water, a similar amount of time should be al lowed for purging of the diluted water in the chamber and complete replacement by groundwater prior to sub sequent sample collection. W hile we did not do so, use of a solution containing a tracer not found in the native water, for example, Br could help determine if any di lution by the displacement water had occurred and, if it had, to determine the extent of dilution, allowing for correction of the results. Use of the sample chambers installed in the cave walls of submerged caves can provide information to determine chemical changes in the water as it enters the conduit through the microbial mat. An added advantage of the design is that seepage rates can also be established for purposes of, for example, mixing calculations. Sam pling frequency and the density of the installations are limited only by diver skill and availability, aside from any restrictions imposed by land owners or stewards on studies similar to the ones we are conducting with these devices. A N APPROACH FOR COLLECTION OF NEARFIELD GROUND W ATER SAMPLES IN SUBMERGED LIMESTONE CAVERNS

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ACTA CARSOLOGICA 42/2-3 2013 234 Barton, H.A. & F. Luiszer, 2005: Microbial metabolic structure in a suldic cave hot spring: potential mechanisms of biospeleogenesis.Journal of Cave and Karst Studies, 67, 1, 28. Brigmon, R.L., Martin, H.W ., Morris, T.L., Bitton, G. & S.G. Zam, 1994: Biogeochemical Ecology of io thrix Spp in Underwater Limestone Caves.Geomi crobiology Journal, 12, 3, 141. Bruckner, A.E., Hornberger, G.M. & A.L. Mills, 1984: Groundwater seepage in a piedmont impoundment. Practical applications of groundwater models, Pro ceedings of the National Meeting on Groundwater Modeling. National W ater W ell Association, 570 583. Bruckner, A.E., Hornberger, G.M. & A.L. Mills, 1989: Field measurement and associated controlling fac tors for groundwater seepage in a piedmont im poundment.Hydrological Processes, 2, 223. Chen, Y., W u, L.Q ., Boden, R., Hillebrand, A., Kumare san, D., Moussard, H., Baciu, M., Lu, Y.H. & J.C. Murrell, 2009: Life without light: microbial diver sity and evidence of sulfurand ammonium-based chemolithotrophy in Movile Cave.ISME Journal, 3, 9, 1093. 10.1038/ismej.2009.57. Cline, J.D., 1969: Spectrophotometric determination of hydrogen sulde in natural waters.Limnology and Oceanography, 14, 454. Engel, A.S., Lee, N., Porter, M.L., Stern, L.A., Bennett, P.C. & M. W agner, 2003: Filamentous "Epsilonpro teobacteria" dominate microbial mats from suldic cave springs.Applied and Environmental Microbi ology, 69, 9, 5503. Engel, A.S. & K.W Randall, 2011: Experimental Evi dence for Microbially Mediated Carbonate Disso lution from the Saline W ater Zone of the Edwards Aquifer, Central Texas.Geomicrobiology Journal, 28, 4, 313. 10.1080/01490451.2010.500197. Engel, A.S., Stern, L.A. & P.C. Bennett, 2004: Microbial contributions to cave formation: New insights into sulfuric acid speleogenesis.Geology, 32, 5, 369 372. 10.1130/g20288.1. Erlich, H.L., 1996: Geomicrobiology Marcel Dekker, Inc., New York, NY. Fellows, C.R. & P.L. Brezonik, 1980: Seepage Flow into Florida Lakes.W ater Resources Bulletin, 16, 4, 635. 10.1111/j.1752.1980.tb02442.x. Flewelling, S.A., Hornberger, G.M., Herman, J.S. & A.L. Mills, 2012: Travel time controls the magnitude of nitrate discharge in groundwater bypassing the ri parian zone to a stream on Virginias coastal plain.Hydrological Processes, 26, 1242. DOI: 10.1002/hyp.8219. Franklin, R.B., Giannotti, A.L., Tysall, T.N. & A.L. Mills, 2005a: Geomicrobiology of phreatic caves associ ated with central Florida springs, 30th Anniversary National Cave and Karst Management Symposium, Albany, New York. Franklin, R.B., Tysall, T.N., Gianotti, A. & A.L. Mills, 2005b: Geomicrobiology of phreatic caves associ ated with central Florida springs, Spring Meeting, American Geophysical Union, New Orleans, LA. Herman, J.S., Hounshell, A.G., Franklin, R.B. & A.L. Mills, 2013: Biological control on acid generation at the conduit-bedrock boundary in submerged caves: Q uantication through geochemical modeling.Acta Carsologica, 42, 2, 45. Holly, M.A., Lubetsky, J.S. & C.F. Harvey, 2003: Charac terizing submarine groundwater discharge: A seep age meter study in W aquoit Bay, Massachusetts.Geophysical Research Letters, 30, 6, 1297. 10.1029/2002GL016000. Koopmans, D.J. & P. Berg, 2011: An alternative to tra ditional seepage meters: Dye displacement.W ater Resources Research, 47, 1, W01506. 10.1029/2010WR009113. Lee, D.H., 1977: A device for measuring seepage ux in lakes and estuaries.Limnology and Oceanography, 22, 1, 140. Lehman, R.M. & A.L. Mills, 1994: Field evidence for copper mobilization by dissolved organic matter.W ater Research, 28, 2487. Lock, M.A. & P.H. John, 1978: Measurement of Ground water Discharge into a Lake by a Direct method.Internationale revue der gesamten hydrobiologie, 63, 2, 271. 10.1002/iroh.19780630212. McIntire, P.E., Mills, A.L. & G.M. Hornberger, 1987: A groundwater seepage meter for use in lakes with low groundwater ow and high biogenic gas pro duction, Annual Meeting of the American Geo physical Union, San Francisco, CA. McIntire, P.E., Mills, A.L. & G.M. Hornberger, 1988: Groundwater lake interactions and the occurrence of high sulfate concentrations at depth in the sedi ment of Lake Anna, Virginia.Hydrologic Process es, 2, 207. REFERENCES AARON L. MILLS, TERRENCE N T YSALL & J ANET S. HERMAN

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ACTA CARSOLOGICA 42/2-3 2013 235 Otte, M.L. & J.T. Morris, 1994: Dimethylsulphoniopro pionate (DMSP) in Spartina Alterniora Loisel.Aquatic Botany, 48, 3, 239. Paulsen, R.J., Smith, C.F., O'Rourke, D. & T.F. W ong, 2001: Development and evaluation of an ultrasonic ground water seepage meter.Ground W ater, 39, 6. doi:10.1111/j.1745.2001.tb02478.x. Rosenberry, D.O. & R.H. Morin, 2004: Use of an elec tromagnetic seepage meter to investigate temporal variability in lake seepage.Ground W ater, 42, 1, 68. 10.1111/j.1745.2004.tb02451.x. Sacks, L.A. 1996: Geochemical and isotopic composition of ground water with emphasis on sources of sulfate in the Upper Floridan aquifer in parts of Marion Sumter, and Citrus Counties, Florida U.S. Geologi cal Survey W ater-Resources Investigations Report: 95. Sarbu, S.M., Kane, T.C. & B.K. Kinkle, 1996: A chemoau totrophically based cave ecosystem.Science, 272, 5270, 1953. Schincariol, R.A. & J.D. McNeil, 2002: Errors with small volume elastic seepage meter bags.Ground W ater, 40, 6, 649-651. 10.1111/j.1745-6584.2002. tb02551.x. Sholkovitz, E., Herbold, C. & M. Charette, 2003: An au tomated dye-dilution based seepage meter for the time-series measurement of submarine groundwa ter discharge.Limnology & Oceanography Meth ods, 1, 16-28. Stumm, W & J.J. Morgan, 1996: Aquatic Chemistry: Chemical Equilibria and Rates in Natural Waters John W iley & Sons, Inc., New York, NY, 1022. Vlasceanu, L., Sarbu, S.M., Engel, A.S. & B.K. Kinkle, 2000: Acidic cave-wall biolms located in the Fras assi Gorge, Italy.Geomicrobiology Journal, 17, 2, 125-139. A N APPROACH FOR COLLECTION OF NEARFIELD GROUND W ATER SAMPLES IN SUBMERGED LIMESTONE CAVERNS



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O RGANIC MATTER FLU X IN THE EPIKARST OF THE D ORVAN KARST F RANCE T OK ORGANSKE SNOVI V EPIKRASU: P RIMER KRA KEGA SISTEMA D ORVAN F RANCIJA Kevin S. S IMON 1 Izvleek UDK 551.44:556.33:543.38(445.61) Kevin S. Simon: Tok organske snovi v epikrasu: Primer krakega sistema Dorvan, Francija Kljub temu, da gre za pomembne dejavnike krakih ekosiste mov, obstajajo le redke tudije o sestavi in porazdelitvi organske snovi v krakih vodonosnikih. V lanku predstavimo rezultate dveletnega zveznega opazovanja toka detrita in organizmov (ivali) v prenikli vodi in v jamskem potoku, ki ga napaja epikras, v krakem sistemu Dorvan, Francija. V obeh okoljih smo zaznali veliko asovno spremenljivost opazovanih parametrov, ki pa ni bila sezonske narave. 30-69% spremenljivosti pre toka ivali v obeh habitatih in detrita v epikraki prenikli vodi, lahko poveemo s spremenljivostjo pretoka. Pretok detrita v jamskem potoku pa je povezan z najvejimi mesenimi pretoki. Razline znailnosti pretoka organske snovi v obeh opazovanih habitatih, kaejo na razline dejavnike vpliva. ivi organizmi tvorijo veino toka organskih delcev, kar kae na velik pomen ekolokih procesov v prenosu organske snovi. Kljune besede : ogljik, energija, epikraki tok, sezonskost, po tok. 1 School of Environment, e University of Auckland, Private Bag 92019, Auckland 1142, New Zealand, Fax: 6493737434, e-mail: k.simon@auckland.ac.nz Received/Prejeto: 12.3.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 237, POSTOJNA 2013 Abstract UDC 551.44:556.33:543.38(445.61) Kevin S. Simon: Organic matter ux in the epikarst of the Dorvan karst, France Availability of organic matter plays an important role in karst ecosystems. Somewhat surprisingly, study of the composition and distribution of organic matter in karst aquifers is rare. e most comprehensive study or organic matter ux to date is a two year continuous monitoring of detritus and animal ux in epikarst drip waters and an epikarst-fed cave stream in the Dor van karst, France. Analysis of those data reveals high temporal variation in detritus and animal ux in both habitats, but little evidence of seasonality in ux. W ater ux explained 30-69% of the variation in animal ux in both habitats and detritus ux in the epikarst seepage water. Detritus ux in the cave stream was better explained by peak monthly discharge. Lack of coher ence between organic matter ux in epikarst seepage and the epikarst stream suggests organic matter transport is governed by diering factors in the two habitats. Overall, much of the particulate organic matter ux in the epikarst occurs as living animals suggesting a dominant role of ecological processes in organic matter transport. K eywords : Carbon, energy, epikarst, ux, seasonality, stream. I NTRODUCTION Caves are typically thought to be energy limited eco systems with food webs that rely on either dead organic matter, detritus, imported from surface soils and vegeta tion (Hppop 2000; Poulson & Lavoie 2001; Simon et al. 2007) or on internal chemoautotrophic production (Sar bu et al. 1996). Indeed, a suite of life history characters of cave animals have been considered to be adaptations to low energy availability in caves (see Hppop 2000 for a review). Such a strong role of energy availability in the ecology and evolution of cave animals suggests a clear

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ACTA CARSOLOGICA 42/2-3 2013 238 KEVIN S. S IMON understanding of organic matter dynamics in cave sys tems is critical. Historically, this has not been the case with few studies directly examining the pattern and driv ers of organic matter availability in caves and the subse quent consequences of variation in energy availability in cave systems. is is changing, however, with a growing body of research that addresses various facets of organic matter dynamics in caves. Both dissolved and particulate organic matter ap pear to be important for aquatic cave fauna and the avail ability of these resources is partially a function of geology and surface features (Simon et al. 2007). For example, the presence of large openings and sinking streams should permit entry of large detrital particles (e.g. Souza-Silva et al. 2012) which appear to attract and be used by ani mals (Gibert et al. 1994). In contrast, soil layers and ne fractures in epikarst likely restrict movement of large particles into caves while permitting ne particle and dissolved organic matter entry. Dissolved organic mat ter (DOM) should be more universally available, but likely varies in quantity and composition based on ow paths. e role of DOM for higher trophic levels have been examined (Simon et al. 2003; Cooney et al. 2009). Such studies have documented the importance of micro bial production based in chemoautotrophy (Sarbu et al. 1996) and heterotrophy (Simon et al. 2009) for higher trophic levels. e importance of particulate organic matter (POM) for cave food webs is less well studied, but it also appears to serve as an important energy resource for at least some cave animals (Huntsman et al. 2011, Ve narsky 2012). e pace and drivers of particulate organic matter decomposition has been examined in several cave systems (Simon et al. 2001; Kinsey et al. 2007; Huntsman et al. 2011; Venarsky et al. 2012). Such studies have sug gested strong interactions among consumers and micro bial lms in dictating decomposition rate of particulate detritus and its likely use in food webs. W hile the importance and use of detritus in cave streams has been shown, surprisingly few studies have quantied the amount of detritus in caves, its composi tion, and factors that dictate detritus distribution. Such data are critically important as they determine the avail ability and distribution of suitable energy sources in caves. Examination of dissolved organic matter (DOM) composition (Birdwell & Engel 2010; Simon et al. 2010) has illuminated multiple sources of DOM and temporal and spatial variation in its form and abundance. Stand ing stocks of various types of POM have been measured in a few cave systems (e.g. Simon et al. 2001; Huntsman et al. 2011; Venarsky 2012), but examination of its com position is limited (but see Hutchins et al. this issue). Transport of organic matter into and within caves has been poorly studied despite the fact that it is central to understanding energy distribution in caves. By combin ing transport distances and decomposition rates, Simon et al. (2001) found that litter turnover distances were quite short suggesting transport of detritus deep into caves may be quite limited. is pattern has yet to con rmed beyond a single cave system. Direct measures of transport into and within caves are quite rare. Graening et al. (2003), in what is the most complete energy budget for a cave system to date, reported annual rates of im port and export of organic matter to two cave streams. He showed that ux of DOM exceeded POM, a phenon menon previously shown by Gibert (1986) in a French aquifer. Souza-Silva et al. (2012) measured bi-monthly ux of detritus into a Brazilian cave, showing high varia tion and seasonality of detritus import to a cave stream fed by a sinking stream. e most comprehensive study to date detailing the dynamics of organic matter ux in an epikarst-fed cave system is that of the Dorvan karst by Gibert (1986). Over a two year period she measured ux of particulate de tritus and animals in epikarst drips and a cave stream at high temporal resolution (days to months). Her data from over 30 years ago remain the single most compre hensive study of organic matter ux in a cave system. As such, they provide a unique opportunity to examine the relative magnitude of organic matter ux in dead and living forms and to address factors that dictate that ux. I reanalyzed Giberts data with the goal of answering the following: 1) how large is the variation in organic matter ux in the epikarst, is that variation seasonal, and what factors dictate that variation? 2) how does organic mat ter transport as detritus and animals compare and are they likely regulated by the same factors? METHODS e Dorvan-Cleyzieu aquifer is a 10 km 2 low-moun tain karst (average altitude 620 m and average gradient 126 m/km) situated on the southwestern range of the Southern Jura mountains of France (see Gibert 1986 for a detailed description). It lies in middle Jurassic limestone (Bathonian and Bajocian) 50 km from the city of Lyon. At the top of the aquifer, the epikarst is drained by the Cor moran Cave stream (mean annual discharge 4.2 L/s). In

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ACTA CARSOLOGICA 42/2-3 2013 239 O RGANIC MATTER FLU X IN THE EPIKARST OF THE D ORVAN KARST F RANCE side Cormoran Cave, drips from the ceiling permit direct sampling of water exiting the epikarst. ere are no large openings or sinking streams that feed Cormoran Cave so water entering the system must travel through soils and fractures in the epikarst. W ater traveling through the aquifer ultimately exits at Pissoir Spring (mean annual discharge 76.5 L/s, range: 0 to 2,000 L/s) and from several seeps around the spring. Pissoir Spring acts as an over ow for the aquifer, owing only when the water level in the aquifer reaches a sucient height. Aer exiting the aquifer, water from the spring joins the Bief Ravinet stream and ows through an alluvial plain to the Albar ine River. e annual hydrologic cycle consists of a high-wa ter period during winter characterized by relatively high discharge and frequent oods. During summer, water levels are generally low with the Cormoran Cave stream maintaining very low ow while the Pissoir Spring ows only intermittently. Vegetation covers 8090% of the ground surface (meadows and crops 50%, forests 40%, moors and bushes 10%). Agriculture on the surface has had some inuence on groundwater quality, with chlo ride and nitrate concentrations ranging up to 8 and 14.7 mg/L, respectively, on some occasions in the Cor moran Cave stream (Gibert 1986 & 1990; Simon et al. 2001). Detritus and animals in the Cormoran Cave stream and a series of epikarst drips were collected continuously from March 1978 to February 1980. A barrier was in stalled in the Cormoran Cave stream that directed the entire stream ow into two 150m nets, one that sampled surface ow and a second that sampled transport below the surface. Data from the two nets were combined for analysis in this paper to analyze total ux in the stream. W ater level at the barrier was continuously recorded and converted to discharge with a rating curve. A series of epikarst drips in a 6 m 2 area of the cave ceiling adjacent and above the stream were directed with a waterproof tarpaulin into a 100 m net and container. Total water accumulating in the container was measured to calculate water ux between sampling periods. Contents of the nets were collected at irregular intervals ranging from 2 days to 2 weeks over the 2 year period. Invertebrates were identied and counted, but not weighed. Remaining ma terial in the nets was dried at 70C, weighed to calculate dry mass, combusted at 550C, and then reweighed to determine mineral mass. Organic matter mass was cal culated as the dierence between dry mass and the mass of mineral material remaining aer combustion. I used the high-frequency sampling data archived in Gibert (1986) to calculate uxes (mass per unit sampling period) and concentrations (mass per volume) of dead organic matter (detritus), mineral matter, and ground water animals exiting the epikarst drips and in the cave stream on a monthly basis. e aquatic fauna in the dri were dominated by amphipods ( Niphargus rhenorhod anensis ) and harpactacoid and cyclopoid copepods, so I restricted my analysis to those taxa. Because animals were counted but not weighed I converted dri density to mass using conversion factors specic to the Dorvan karst for Niphargus and values for other freshwater taxa for copepods. Niphargus dri density was converted to biomass using a value of 5.58 mg/individual, which was the average of 44 male and 50 female Niphargus collected from the Dorvan groundwater in 1980 (Gibert 1986). Harpactacoid density was converted to mass using a conversion factor of 0.0046 mg/individual, a mean of ten freshwater epibenthic harpactacoid taxa (Goodman 1979). Cyclopoid density was converted to mass using a value of 0.0014 mg/individual from surface stream cy clopoids (ODoherty 1985). Total mass of copepods was calculated as the sum of haracticoid and cyclopoid co pepods. I compared temporal patterns in organic, inorganic and animal dri using Pearson correlation analysis. Re lationships between monthly ux of detritus and animals and ux of water were examined using linear regression. e relationship between organic matter and mineral matter export was examined using standard major axis regression on the high frequency data. Monthly concen trations of detritus and animals in the epikarst and in the cave stream were compared using paired T-tests. Results were considered statistically signicant at n= 0.05. All statistical analysis were performed using R (R Core De velopment Team 2010) and I used the lmodel2 package for SMA regression. RESULTS e monthly ux of detritus, Niphargus and copepods varied by 2 orders of magnitude in the cave stream over the two year period (Fig. 1). ere was no obvious sea sonality in ux although lowest values tended to occur in late summer and autumn. Niphargus ux exceeded detri tus by about an order of magnitude in all but one month. e ux of organic matter as copepods was several orders of magnitude lower than detritus and Niphargus e

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ACTA CARSOLOGICA 42/2-3 2013 240 Fig. 1: Monthly ux of particulate organic matter as detritus, Niphargus and copepods in water seeping from the epikarst and in the Cormoran Cave stream. uxes of detritus and Niphar gus were positively correlated (p=0.017, r=0.49), but neither was correlated with copepod ux (p>0.960). In the epikarst, detritus and animal ux were more variable with detritus and, particularly, Niphargus ux absent in several months (Fig. 1). ere was over 3 or ders of magnitude variation in detritus and somewhat lower range in animal ux across months. Unlike the cave streams, ux of detritus nearly always exceeded ux of animals, largely due to low ux or absence of Niphargus As in the cave stream, there was no obvious seasonality in ux although low values were common in late summer and autumn. Overall, there was no correlation between the epikarst and the stream for ux of detritus (p=0.263), Niphargus (p=0.090) or co pepods (p=0.426). Fig. 2: Relationship between detritus or animal ux and discharge over the 2 year study period in the epikarst seepage water and the Cormoran Cave stream. Lines and statistical values are results of linear regressions of log 10 transformed data. KEVIN S. S IMON

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ACTA CARSOLOGICA 42/2-3 2013 241 Variation in discharge explained 30-69% of the monthly variation of animals in both the epikarst and the cave stream and 47% of the variation in detritus ux from the epikarst. Notably, monthly variation in detritus ux in the stream was unrelated to variation in discharge (Fig. 2) W hile total water ux did not predict detritus ux, maximum monthly discharge explained about 30% of the variation in detritus ux in the cave stream (Fig. 3). e ux of detritus and mineral matter were highly cor related in both the cave stream and the epikarst (Fig. 4) with a SMA slope close to one (1.08) in the epikarst and slightly lower (0.78) in the cave stream. In contrast, the ux of Niphargus and copepods were unrelated (p>0.23) to ux of mineral material. Concentrations of detritus, Niphargus and cope pods varied by 1 orders of magnitude over time in the epikarst and cave stream water, but displayed no particular seasonal pattern (Fig. 5). In the epikarst, con centrations of detritus exceeded that of animals. In the cave stream concentrations of Niphargus exceeded that of detritus and copepods. Concentrations of detritus and copepods were >500 (p = 0.047) and >100 (p < 0.001) times higher, respectively in the epikarst water than in the cave stream. ere was no dierence (p = 0.068) in Niphargus concentration between habitats. Fig. 3: Relationship between monthly ux of detritus and maxi mum monthly discharge in the Cormoran Cave stream over the 2 year study period. Fig. 4: Relationship between detritus ux and mineral material ux in epikarst seepage water and the Cormoran Cave stream. Statistical parameters are results of standardized major axis re gressions. Fig. 5: Monthly concentrations of particulate organic matter as detritus, Niphargus and copepods in water seeping from the epikarst and in the Cormoran Cave stream. O RGANIC MATTER FLU X IN THE EPIKARST OF THE D ORVAN KARST F RANCE

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ACTA CARSOLOGICA 42/2-3 2013 242 ere was substantial variation in organic matter ux in the Drovan epikarst with up to several orders of mag nitude variation among months for both detritus and animals. W hile the variation in ux was fairly large, there was a distinct lack of seasonality in the transport of organic matter in the epikarst drip water and the cave stream. In temperate forested surface streams as would have been typical of the Dorvan area, detritus input and transport is highly seasonal because of the temporal pat tern of leaf fall (Beneld 1997). e lack of direct ripar ian input obviously changes the magnitude and timing of organic matter ux in cave streams compared to surface streams. Indeed, the epikarst-fed Cormoran Cave stream behaved more similarly to surface streams that lack ripar ian vegetation in desert and tundra landscapes (Schade & Fisher 1997). e maximum concentration of organic matter in the cave stream (0.129 mg/m 3 ) was more than three orders of magnitude lower than the lowest value reported across a range of surface streams (210 mg/m 3 Golladay et al. 1991) demonstrating the scarcity of this energy source in the Cormoran Cave stream. Clearly, and not surprisingly, the epikarst provides an eective, if not complete, barrier to particulate matter inux to caves. is pattern is unlikely to be the case in all cave streams, particularly those with larger openings to the surface. For example, Souza-Silva (2012) found highly variable detritus input to a Brazilian cave with a seasonal pattern driven by rainfall and streamow into the cave entrance. Flux of detritus to that cave stream (up to 263 g/day) far exceed the ux of detritus observed in Dorvan. Such streams with direct connectivity to the surface likely be have more similarly to related surface streams. Temporal variation in detritus and animal ux was related to the amount of water moving through the systems, but water ux accounted for only 3069% variation in organic matter ux. In the cave stream, de tritus ux was entirely unrelated to water ux. Clearly factors other than simple mass ux of water drive or ganic matter transport in epikarst seeps and streams. In the case of detritus in the cave stream, peak monthly discharge provided a better predictor of organic matter transport. is phenomenon has been observed in sur face streams (W allace et al. 1995) where detritus reten tion can be quite high and episodically driven by large ows that mobilize material from the streambed. is is probably even more common in cave streams with direct surface connection and highly variable ow. Fu ture work on organic matter transport in caves should pay particular attention to periods of ow variation. Not only do high discharge events likely inuence de tritus transport, they appear to also inuence micro bial activity (Simon et al. 2001) providing the opportu nity to couple organic matter transport with ecological consequences. e relative pattern of detritus and animal ux in the epikarst and cave stream provides insight into how epikarst systems function. e lack of correlation in ux of detritus and animals between the epikarst and cave stream indicate dierent factors drive organic matter ux through the two habitats even though they are ultimate ly fed by the same system, i.e. surface soils and epikarst fractures. e bulk of organic matter exiting epikarst drips was detritus whereas it was predominantly amphi pods in the cave stream. In addition, concentrations of detritus were higher in the epikarst drip water than in the cave stream. Gibert (1986) attributed lower detritus concentration in the cave stream to settling of detritus in the stream where it is presumably processed. Such detritus retention is typical of surface streams (W ebster et al. 1999) and is consistent with short turnover lengths of leaf litter observed in cave streams (Simon et al. 2001). e closer relationship between detritus ux and water ux in the epikarst suggests a more passive process of bulk mobilization of presumably small particles through the epikarst matrix, although addressing this is challeng ing because of diculty in accessing the epikarst struc ture. It is unkown if detritus size distribution diered in the epikarst drip water and cave stream water. Such size partitioning has not been conducted in cave systems al though is has been shown to be critically important in organic matter transport in surface streams (omas et al. 2001). e transport of organic matter as dead detritus and living animals appears to be governed at least par tially by dierent factors. In both the epikarst and the cave stream detritus ux and mineral ux were strongly related. In the epikarst the SMA slope was very close to 1 suggesting detritus transport almost perfectly mirrored mineral material transport. e SMA slope in the cave stream was shallower but detritus and mineral ux were strongly related. is suggests detritus ux is quite pas sive and factors that drive it are closely aligned with fac tors that drive sediment ux. Of course, this likely only applies to very ne sediments as would likely have been moving through the Dorvan epikarst. In contrast, ani mal ux was unrelated to transport of mineral material. is likely represents behavioral dri by animals, which has been the subject of much study (Gibert 1986; Rouch 1986). Considering much of the organic matter ux in the epikarst was in the form of living animals, behavior may play a signicant role in organic matter distribution in epikarst habitats. D ISCUSSION KEVIN S. S IMON

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ACTA CARSOLOGICA 42/2-3 2013 243 A CKNO WLEDGEMENTS is manuscript is dedicated to Janine Gibert, a true pioneer in karst ecosystem science. Gibert (1986) used a pioneering approach of com bining dri sampling of animals in karst based on the work of Rouch (1986) with the organic matter budgeting used in surface streams. More than thirty years later her data remain the most comprehensive set of data on or ganic matter ux in a karst aquifer. W hile this budgeting process has been extensively used in surface systems with some systems having decades of ux and stocks mea sures (e.g. W allace et al. 1995) it has not been commonly applied to cave streams. Such work is labor intensive, but it provides the basis for developing mechanistic models of organic matter dynamics in aquatic systems which are sorely needed to interpret ecological patterns. A growing body of work focusing on the nature of energy limita tion and ecological interactions in cave food webs would be well served by basic data regarding how organic mat ter moves through cave systems. Recent demonstra tion of linkages among detritus, microbial respiration, CO 2 and limestone dissolution (Covington et al. 2013) show such data may also help explain physical process es in karst aquifers. Giberts work has illuminated how epikarst driven systems function but it also raises new questions. If the ux of organic matter from the epikarst as animals exceeds that of detritus what fuels that animal production? How strongly connected are the ne frac ture systems of the epikarst with the cave streams that drain them? How do cave streams with larger openings compare and what dictates energy ux through those systems? Such questions will be well served by the type of approach used by Gibert in the Dorvan karst. REFERENCES Beneld, E.F., 1997: Comparison of litterfall input to streams.Journal of the North American Bentho logical Society, 16, 104. Birdwell, J.E. & A.S. Engel, 2010: Characterization of dis solved organic matter in cave and spring waters us ing UV-Vis absorbance and uorescence spectros copy.Organic Geochemistry, 41, 3 270. Cooney, T.J. & K. S. Simon 2009: Inuence of dissolved organic matter and invertebrates on the function of microbial lms in groundwater.Microbial Ecology, 58, 599. Covington, M.D., Prelovek, M. & F. Gabrovek, 2013: Inuence of CO 2 dynamics on the longitudinal vari ation of incision rates in soluble bedrock channels: Feedback mechanisms.Geomorphology, in press. Gibert, J., 1986: Ecologie d'un systeme karstique juras sien. Hydrogologie, drive animale, transits de matires, dynamique de la population de Niphargus (Crustac Amphipode).Mmoires de Biospolo gie, 13, 1. Gibert, J., Vervier, P., Malard, F., Laurent, R. & J.L. Rey grobellet, 1994: Dynamics of communities and ecology of karst ecosystems: Example of three karsts in Eastern and Southern France.In: J. Gibert et al. (eds.) Groundwater Ecology Academic Press, pp. 425, San Diego. Goodman, K.S., 1980: e estimation of individual dry weight and standing crop of harpacticoid copep ods.Hydrobiologia, 72, 253. Golladay, S.W ., 1997: Suspended particulate organic mat ter concentration and export in streams.Journal of the North American Benthological Society, 16, 1, 122. Graening, G.O., & A.V. Brown, 2003: Ecosystem dynam ics and pollution eects in an Ozark cave stream.Journal of the American W ater Resources Associa tion 39, 497. Huntsman, B.M., Venarsky, M.P. & J.P. Benstead, 2011: Relating carrion breakdown rates to ambient re source level and community structure in four cave stream ecosystems.Journal of the North American Benthological Society, 30, 882. O RGANIC MATTER FLU X IN THE EPIKARST OF THE D ORVAN KARST F RANCE

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ACTA CARSOLOGICA 42/2-3 2013 244 Hppop, K., 2000: How do cave animals cope with the food scarcity in caves?, In: H. W ilkens, H., Culver, D.C., & Humphreys, W .F.. (eds.), Subterranean Eco systems Elsevier, pp. 159, Amsterdam. Kinsey J., Cooney, T.J. & K.S. Simon, 2007: A compari son of the leaf shredding ability and inuence on microbial lms of surface and cave forms of Gam marus minus Say.Hydrobiologia, 589, 199. ODoherty, E.C., 1985: Stream-dwelling copepods: their life history and ecological signicance.Limnology and Oceanography, 30, 3, 554. Poulson, T.L. & K.H. Lavoie, 2001: e trophic basis of subsurface ecosystems.In: W ilkens, H. et al. (eds.) Ecosystems of the World; Subterranean Ecosystems Elsevier pp. 231, New York R Development Core Team, 2010: R version 2.11.1 R Project for Statistical Computing, Vienna, Austria. www.r-project.org Rouch, R., 1986: Sur lcologie des eaux souterraines dans la karst.Stygologia, 2, 352. Sarbu, S.M., Kane, T. C. & B. F. Kinkel, 1996: A chemo autotrophically based cave ecosystem.Science, 272, 1953. Schade, J.D. & S.G. Fisher, 1997: Leaf litter in a Sono ran desert stream ecosystem.Journal of the North American Benthological Society, 16, 612. Simon, K.S., Gibert, J., Petitot, P. & R. Laurent, 2001: Spatial and temporal patterns of bacterial density and metabolic activity in a karst aquifer.Archiv fr Hydrobiologie, 151, 67. Simon, K.S., & E.F. Beneld, 2001: Leaf and wood break down in cave streams.Journal of the North Ameri can Benthological Society, 482, 31. Simon, K.S., Beneld, E.F. & S.A. Macko, 2003: Food web structure and the role of epilithic lms in cave streams.Ecology, 84, 2395. Simon, K.S., Pipan, T. & D.C. Culver, 2007: A conceptual model of the ow and distribution of organic car bon in caves.Journal of Cave and Karst Studies, 69, 2, 279. Simon, K.S., Pipan, T., Ohno, T. & D.C. Culver, 2010: Spatial and temporal patterns in abundance and composition of dissolved organic matter in two karst aquifers.Fundamental and Applied Limnol ogy, 177, 2, 81. Souza-Silva M., Ferreira de Olivera Bernardi, L., Paren toni Martins, R. & R. Lopes Ferreira, 2012: Trans port and consumption of organic detritus in a neo tropical limestone cave.Acta Carsologica, 41, 1, 139. omas, S.A., Newbold, J.D., Monaghan, M.T., Min shall, G.W ., Georgian, T. & C.E. Cushing, 2001. e inuence of particle size on seston deposition in streams.Limnology and Oceanography, 46, 1415 1424. Venarsky, M.P., 2012: e inuence of energy availability on population-, communityand ecosystem-level processes in cave stream ecosystems. Doctoral Dis sertation. e University of Alabama. 136 pp. Venarsky, M.P., Benstead, J.P. & A.D. Huryn, 2012: Ef fects of organic matter and season on leaf litter col onisation and breakdown in cave streams.Fresh water Biology 57, 4, 773. W allace, J.B., W hiles, M.R., Eggert, S., Cuney, T.F., Lugthart, G.J. & K. Chung, 1995: Long-term dynam ics of coarse particulate organic matter in three Ap palachian mountain streams.Journal of the North American Benthological Society, 14, 217. W ebster, J.R., Beneld, E.F., Ehrman, T.P., Schaeer, M.A, Tank, J.L., Hutchens, J.J. & D.J. DAngelo, 1999: W hat happens to allochthonous material that falls into streams? A synthesis of new and published information from Coweeta.Freshwater Biology, 41, 687. KEVIN S. S IMON



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ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACE SUBSURFACE AND GEOCHEMICAL BOUNDARIES IN THE ED W ARDS A QUIFER, TE X AS, USA V PLIV OKOLJA NA PRODUKCIJO IN TRANSPORT ORGANSKIH SNOVI: P RIMER VODONOSNIKA ED W ARDS, TEX AS, ZDA Benjamin T. HUTCHINS 1 Benjamin F. SCHW ARTZ 1 & Annette S. ENGEL 2 Izvleek UDK 543.38:556.33(736.4 Benjamin T. Hutchins, Benjamin F. Schwartz & Annette S. Engel: Vpliv okolja na produkcijo in transport organskih snoVpliv okolja na produkcijo in transport organskih sno vi: Primer vodonosnika Edwards, Texas, ZDA Zalite cone krakih vodonosnikov so energijsko revni habitati, ki jih oskrbuje tok organske snovi skozi zine in geokemine meje. Za jugozahodno mejo obravnavanega podroja je znailen hiter prehod med dobro prezraeno sladkimi vodami in anoksinimi slanimi vodami, kjer organska snov proizva jajo kemolitoavtotrofni mikrobi. asovno in prostorsko po razdeljenost organskih snovi ob teh mejah, smo raziskovali z izotopskimi in geokeminimi analizami. Vrednosti t 13 C za sus pendirano organsko snov (FPOM) ( ,34 do ,47), so v sunem obdobju, med jesenjo 2010 in pomladjo 2012, upadale (p<0.01) zaradi naraajoega prispevka perito na. Na meji med sladko in slano vodo, so vrednosti t 13 C FPOM (,23 do ,18) korelirane z vrednostmi t 13 C v raz topljenem anorganskem ogljiku, t 13 C DIC ki so med ,55 in ,91 (r 2 =0,33, p=0,005). Vendar je t 13 C FPOM osiromaen glede na t 13 C DIC za 28,44 %, kar ustreza izotopski frakciona ciji pri pretvorbi DIC v organski ogljik v kemolitoavtotrofnih procesih. Vrednosti t 13 C FPOM so naraale s asom (p<0.001). Tako t 13 C DIC (r 2 =0,43, p<0,001), kot tudi t 13 C FPOM (r 2 =0,35, p=0,004) ob meji med sladko in slano vodo naraata vzdol toka podzemne vode v smeri jugozahod-severovzhod. Prostor ska spremenljivost t 13 C DIC ob meji med sladko in slano vodo, je verjetno posledica spremenljivih izvorov kislosti, ki povzroa raztapljanje karbonatov. asovno spremenljivost pa povezu jemo z spremenljivostjo napajanja in vodostaja v vodonosniku, ki vplivata na prenos kemolitoavtotrofne organske snovi preko meje med sladko in slano vodo. Kljune besede: Stabilni izotopi ogljika, prostorska in asovna spremenljivost, kemolitoavtotrofna produkcija, alogeni vnos, kras. 1 Texas State University-San Marcos Department of Biology, 601 University Drive, San Marcos, T X 78666, Benjamin T. Hutching (bh1333@txstate.edu), Benjamin F. Schwartz, e-mail: bs37@txstate.edu 2 e University of Tennessee-Knoxville Department of Earth and Planetary Sciences, 1412 Circle Drive, Knoxville, TN 37996, & Annette S. Engel, e-mail: aengel1@utk.edu Received/Prejeto: 18.1.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 245, POSTOJNA 2013 Abstract UDC 543.38:556.33(736.4 Benjamin T. Hutchins, Benjamin F. Schwartz & Annette S. Engel: Environmental controls on organic matter production and transport across surface-subsurface and geochemical boundaries in the Edwards aquifer, Texas, USA Karst aquifer phreatic zones are energy limited habitats sup ported by organic matter (OM) ow across physical and geo chemical boundaries. Photosynthetic OM enters the Edwards Aquifer of Central Texas via streams sinking along its north eastern border. e southeastern boundary is marked by a rap id transition between oxygenated freshwaters and anoxic saline waters where OM is likely produced by chemolithoautotrophic microbes. Spatial and temporal heterogeneity in OM compo sition at these boundaries was investigated using isotopic and geochemical analyses. t 13 C values for stream ne particulate OM (FPOM) (.34 to .47) decreased during re gional drought between fall 2010 and spring 2012 (p<0.001), and were positively related to FPOM C:N ratios (r 2 =0.47, p<0.001), possibly due to an increasing contribution of per iphyton. Along the freshwater-saline water interface (FW SWI), t 13 C FPOM values (.23 to .18) correlated to t 13 C values for dissolved inorganic carbon (t 13 C DIC ) (.55 to .91) (r 2 =0.33, p=0.005) and were depleted relative to t 13 C DIC values by 28.44, similar to fractionation values attributed to chemo lithoautotrophic carbon xation pathways using DIC as the substrate. t 13 C FPOM values also became enriched through time (p<0.001), and t 13 C DIC values (r 2 =0.43, p<0.001) and t 13 C FPOM values (r 2 =0.35, p=0.004) at FW SWI sites increased with dis tance along the southwest-northeast owpath of the aquifer. Spatial variability in FW SWI t 13 C DIC values is likely due to vari able sources of acidity driving carbonate dissolution, and the temporal relationship is explained by changes to recharge and aquifer level that aected transport of chemolithoautotrophic OM across the FW SWI. Keywords: Carbon stable isotopes, spatial and temporal vari ability, chemolithoautotrophic production, allochthonous in put, karst.

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ACTA CARSOLOGICA 42/2-3 2013 246 e phreatic zone of karst aquifers can support diverse metazoan communities (stygobionts). In fact, some of the most diverse subterranean assemblages yet documented are recorded from extensive phreatic groundwater sys tems (Culver & Pipan 2009). However, karst aquifers are considered to be nutrient-poor, and aquifer assemblages are dependent on organic matter (OM) produced photo synthetically and imported into the subterranean realm via recharging water, gravity, animals (Poulson & Lavoie 2000; Poulson 2005), and plant root exudates (Jasinska et al. 1996), or produced in-situ through chemolithoau totrophy (Sarbu et al. 1996; Pohlman 1997). Consequently, in systems dependent on photosynthetic OM, stygobiont diversity should be predominately focused at the surfacesubsurface interface. But, the quantity and quality of OM entering karst aquifers via recharges change as a function of the seasonality of C3 and C4 plant communities on the surface, as well as benthic stream periphyton production along spatial and seasonal precipitation gradients (Bird et al. 1998; Artman et al. 2003; Silva et al. 2012). ese dif ferences can inuence stygobiont distribution, such that if surface recharge contributions diminish seasonally or over a long period of time due to aquifer evolution, then insitu OM sources become prevalent and stygobionts may be found at redox gradients between oxidizing and reducing waters in chemolithoautotrophic systems. ere has been limited research to investigate OM heterogeneity along re dox gradients in chemolithoautotrophic aquifer systems, although geochemical gradients move vertically (Hum phreys et al. 2012), and potentially laterally (Perez 1986). erefore, to understand how OM controls the distribu tion and diversity of stygobionts in karst aquifers, as well as establishes groundwater food webs, more research is need ed at the groundwater basin scale (Simon et al. 2007). e Edwards Aquifer of Central Texas is one of the most prolic karst aquifers in the world (Lindgren et al. 2004) and the sole source of drinking water for nearly two million people (Johnson et al. 2009) (Fig. 1). e regional climate is sub-tropical humid, with average an nual precipitation ranging from 610 mm in the west to 914 mm in the east (Nielson-Gammon 2008). Precipita tion primarily occurs in spring, and potentially in the fall coinciding with tree leaf drop-o (Short et al. 1984). Re charge (and input of photosynthetic OM) to the aquifer predominantly occurs by streams, fed by karstic ground water from the adjacent Trinity Aquifer, that cross ex posed limestone in the recharge zone (Fig. 1). Crossformational ow from the Trinity Aquifer (Gary et al. 2011 and references therein) is also important, but the nature of OM from this source is not known. South and west of the recharge zone, Edwards limestones are con ned below non-karstic rocks that prevent input of al lochthonous OM. In this conned portion of the aquifer, the southwestern boundary of freshwater is marked by a rapid transition between oxygenated, low total dissolved solids (TDS) waters and dysoxic to anoxic, high TDS waters that contain variably high levels of reduced com pounds, including suldes and ammonia. Six distinct geochemical facies in the saline waters (Oetting et al. 1996) correlate to changes in microbial communities (Gray & Engel 2013) and OM characteristics (Birdwell & Engel 2009). Several lines of evidence suggest that this zone is dominated by chemolithoautotrophic production (Birdwell & Engel 2009; Gray & Engel 2013). Chemo lithoautotrophic production in this part of the aquifer is independent of terrestrial inputs and the habitat is bu ered against seasonal geochemical changes (i.e. changes in water temperature, discharge, conductivity, etc.). In the last three decades, widely available and in expensive methods to analyze stable carbon isotope ratios and carbon (C): nitrogen (N) ratios in OM have contributed to studies of OM origins, OM uxes, food web structure, and the growth and tness of consum ers (Bukovinsky et al. 2012). Because of enzymatic dis crimination against the heavier isotope of carbon ( 13 C) and isotopically distinct inorganic carbon sources, dif ferent carbon xation pathways result in OM with dis tinct carbon isotope compositions (t 13 C) including t 13 C terrestrial C3 plants = to t 13 C terrestrial C4 plants = to and t 13 C chemolithoautotrophic organic matter = to < (Sarbu et al. 1996; Opsahl & Chanton 2006; Fin lay & Kendall 2007; van Dover 2007). In surface aquatic systems, carbon isotopes have been successfully used to quantify the relative contributions of C3 and C4 plants (Stribling & Cornwell 1997), and in subterranean sys tems, isotopic data have been used to dierentiate be tween photosynthetic and chemolithoautotrophic OM (Sarbu et al. 1996). As part of an ongoing investigation of food web dynamics in the Edwards Aquifer, OM at both the sur face-subsurface and freshwater-saline water interface (FW SWI) was isotopically analyzed in a geochemical and environmental framework to quantify spatial and temporal changes and to test the following hypotheses related to OM sources: (H1) the C isotope composition of OM (t 13 C OM ) in recharge streams would become progressively less nega tive along the northeast to southwest precipitation gradi ent, reecting a decrease in the relative proportion of C3 plants; (H2) FW SWI t 13 C FPOM values would be more nega tive than stream t 13 C FPOM values, reecting a greater INTRODUCTION B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 247 contribution of chemolithoautotrophic production, and values would vary across the study area, reecting regional dierences in t 13 C DIC values (the substrate for chemolithoautotrophic production); (H3) Recharge stream t 13 C OM values and C:N ratios would decrease in the summer, reecting a greater rela tive contribution of riparian C3 plants and periphyton during the dry season; and (H4) FW SWI t 13 C FPOM values and FPOM C:N val ues would remain constant over time, reecting a decou pling between surface seasonality and chemolithoauto trophic production. e results from this study provide additional evidence for both photosynthetic and chemolithoauto trophic OM in the Edwards Aquifer. More generally, this research identies potential drivers of spatial and tem poral variability in both sources. MATERIALS & METHODS FIELD SAMPLING AND GEOCHEMICAL ANALYSES Seven surface streams crossing the Edwards Aquifer re charge zone and 11 wells along the FW SWI (Fig. 1) were sampled between one and six times between 3 November 2010 and 29 March 2012 (streams) and between 16 April 2011 and 2 April 2012 (wells). e sampling period was marked by declining aquifer levels and declining stream ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ... Fig. 1: e Edwards Aquifer, formed in Cretaceous limestone, extends in a 400 km arc that varies from 4 to 56 km wide and 137 to 335 m thick. Upli of the Edwards P lateau during the late Cretaceous and downwarping of the Gulf of Mexico during the early Cenozoic led to exposure of Edwards formation limestones at the northern and western boundary of the aquifer (recharge zones) along west-east and southwest-northeast trending en echelon faults (Barker et al. 1994). Recharge stream sampling sites occur in or northwest of the recharge zone. Freshwater-saline water interface sampling sites occur along the freshwater-saline water interface. SAB = Sabinal Rv.; H on = H ondo Cr.; Med = Medina Rv.; H el = H elotes Cr.; Gua = Guadalupe Rv.; Bla = Blanco Rv.; Oni = Onion Cr.; V hm = V erstuy H ome Farm well; Tsc = Tschirhart well; Sat = San Antonio transect wells 1 & 2; Tri = Tri-County 2 well; Nbr = P aradise Alley Shallow; Girl Scout Shallow; and Girl Scout Deep wells; Aqu = Aquarena well; Kyl = Kyle transect wells 1 & 2.

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ACTA CARSOLOGICA 42/2-3 2013 248 and spring ows (Fig. 2), corresponding to a period of prolonged regional drought. Palmer Drought Severity Index (PDSI) data for the Edwards Plateau were obtained from the National Oceanic and Atmospheric Adminis tration, through the Climate Prediction Center ( www. ncdc.noaa.gov/ ) to verify regional water imbalance based on precipitation and soil moisture supply (Palmer 1965). PDSI values above zero correspond to wetter than nor mal conditions, values below zero indicate drier than normal conditions, and values below indicate extreme drought. In recharge streams, FPOM and coarse particulate organic matter (CPOM) were collected in 8 L of unl tered water in sterile carboys aer lightly disturbing the benthos by walking back and forth approximately 7 m upstream of the collection site. e benthos was dis turbed to better represent benthic OM that enters the aquifer during storm events and via downwelling. Per iphyton was collected from cobbles using the methods of Saito et al. (2007), in which three cobbles each from a ribe, run, and pool were scrubbed in the lab using a nylon brush to remove attached periphyton. At FW SWI wells (Fig. 1), two to three well volumes were purged, and physicochemistry was monitored for constituent stability before collecting 8 L of unltered water in ster ile carboys. Samples were stored in the dark on ice until ltration in the lab. FPOM, CPOM, and periphyton were ltered onto 0.7 m, pre-combusted W hatman glass ber lters for isotopic analysis. Filters were fumigated with HCl for 12 to 24 hrs and dried at ~45C. Temperature, dissolved oxygen (DO), pH, and elec trical conductivity (conductivity) were recorded with an In-Situ Inc. Troll 9500 multi-parameter probe with optical DO sensor (accuracy = .1mg/L at 0mg/L DO and .2 mg/L at > 8 mg/L DO). Sulde and am monia concentrations were measured with a CHEMet rics V Multi-analyte photometer via the methyl ene blue and salicylate methods, respectively. If sulde concentration was above the detection limit (0.2 mg/L), sulfate concentration was also measured in the eld colorimetrically using the turbidimetric method. is was done to avoid erroneously high laboratory sulfate concentration measurements (see below) resulting from abiotic sulde oxidation. Additional water samples for ion chromatography and for t 18 O and tD determination were collected and ltered through 0.45 m Fisherbrand nylon syringe lters. In the lab, dissolved ion concentra tions were measured using Dionex ICS ion chro matographs (Bannockburn, IL). Alkalinity as total titrat able bases dominated by bicarbonate was measured via end-point titration with 1.6 N sulfuric acid. t 18 O and tD in liquid water were measured on a Los Gatos Research Liquid W ater Isotope Analyzer (Mountain View, CA). W ater samples for analysis of t 13 C DIC and t 13 C of dissolved organic carbon (t 13 C DOC ) were collected and poisoned with 15 mM sodium azide and stored in glass Fig. 2: H ydrographs for Comal Springs (the largest Edwards Aquifer spring) (dark grey) and Blanco River (a major source of recharge) (light grey), and P almer Drought Severity Index (P DSI) (black) values for Divi sion 6, Edwards P lateau (as de ned by the National Climate Data Center), J anuary 2005 to J anuary, 2013. Sampling events are shown by black vertical bars. indicates stream sampling only. indicates FWSWI site sampling only. Dashed horizontal line is P DSI = 0, representing normal conditions. B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 249 ESTIMATION OF MEAN t 13 C FPOM FOR RECHARGE STREAMS e mean carbon isotope composition of FPOM enter ing the aquifer via recharging streams (t 13 FPOM ) was estimated using Bayesian modeling and isotope values weighted by discharge. is approach allows uncertainty in t 13 FPOM to be quantied by treating each FPOM iso tope measurement, c i as a sample from a separate normal distribution with separate means, i and a common pre cision, 0 (Eq. 1). Uninformative priors were given for i and 0 (Eq. 2). c i ~ N ( i 0 ) (1) i ~ N (0 1 e 6 ) (2) 0 ~gamma(0.001,0.001) (3) Each isotopic value was weighted by daily average stream discharge p i calculated as a proportion of the sum of all daily discharge measurements q i of all streams throughout the study period (Eq. 4). Discharge values were obtained from the nearest United States Geological Survey gauging (4) stations on the sampled streams. e parameter c i was estimated for all unsampled days between the rst and last sampling events by linear interpolation be tween c i values. e posterior probability distribution for t 13 FPOM (Eq. 5) was estimated using a Markov Chain Monte Carlo (MCMC) procedure. (5) Two MCMC chains were run, each with 500,000 iterations, a thinning rate of 50 and a burn-in of 1000. Plots of parameter estimates as a function of MCMC it eration were assessed for adequate burn in, and conver gence was assessed using Gelman and Ruben potential scale reduction factors (Gelman & Ruben 1992). MCMC chains were run in R v2.15 using the rjags package (Plummer 2010). vials with butyl rubber septa (Doctor et al. 2008). Car bon isotope analysis was conducted at the UC Davis Sta ble Isotope Facility using an O.I. Analytical Model 1030 TOC Analyzer (OI Analytical, College Station, T X) in terfaced to a PDZ Europa 20 isotope ratio mass spec trometer (Sercon Ltd., Cheshire, UK). STATISTICAL ANALYSIS Simple linear regressions were used to test for spatial dif ferences in t 13 C FPOM and t 13 C DIC values in streams and FW SWI sites (H1 & H2). Spatial data for sampling sites were assigned in ArcMap 10.0. A curved polyline extend ing between the southwest and northeast margins of the aquifer (approximating the general northwest-southwest direction of groundwater ow) was created using the arc tool. e polyline was converted into 806 points spaced 0.38 km apart from one another and sequentially num bered, beginning with one, at the southwest end. Sam pling sites were assigned a whole number location value corresponding to the number of the nearest point. For FW SWI sites, nested linear models were run to assess re lationships between t 13 C DIC values and location, conduc tivity, and the interaction between location and conduc tivity. Conductivity was log transformed for normality and the relative t of models was assessed using Akaike Information Criterion (AIC) for nite samples. Conduc tivity was not used as a covariate for regressions of stream t 13 C FPOM values against location. Stream t 13 C FPOM values were square-root transformed for normality. To quantify dierences in OM in streams versus FW SWI sites, analyses of covariance (ANCOVA) were used to test for dierences in t 13 C DOC and t 13 C FPOM val ues between stream and FW SWI samples, controlling for date as a confounding variable (H2). To elucidate poten tial inuences (e.g., origins and processing) on the t 13 C of OM in both streams and FW SWI sites, simple linear regressions were used to test for relationships between t 13 C FPOM values and t 13 C DIC values (H2) and between t 13 C FPOM values and FPOM C:N ratios (H3). C:N ratios were log transformed. A matrix of Pearsons productmoment correlation coecients for isotopic and physi cochemical data was visually assessed for additional, potentially signicant correlations (H2). Analysis of vari ance (ANOVA) was used to test for dierences in t 13 C of ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...

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ACTA CARSOLOGICA 42/2-3 2013 250 RESULTS During the 16-month study, 24 samples were collected from recharge streams and 32 samples were collected from FW SWI sites (Fig. 1). Stream ow varied be tween 0 m 3 s and 73.1 m 3 s over the course of the en tire study period (Fig. 2). However, during individual sampling events, streams always had detectible ows (Fig. 2). Sampling corresponded to a period of declin ing stream and spring ows during the summer of 2011 and a period of increasing stream and spring ows in the fall and winter of 2011 and 2012 (Fig. 2). PDSI val ues ranged from wetter than normal conditions prior to July 2010 through declining values indicative of drought conditions throughout 2011 and 2012. e most severe drought condition recorded was in August 2011, which corresponded to lowest discharge for the Blanco River and Comal Springs (Fig. 2). e mean of the posterior probability distribution for the estimate of t 13 C FPOM in recharging streams weighted by discharge was 21.75 (95% equal-tail credible interval = 23.39 to .13), and was similar to the unweighted analyti cal average value (24.22). FW SWI sites had 5.1 higher t 13 C DIC values, 8.76 lower t 13 C FPOM values, and 2.6 X lower FPOM C:N ratios than streams (Fig. 3). e average FPOM C:N ratio was 3.3 (range = 1.85 to 5.17) at F W S W I sites and 8.6 (range = 2.14 to 33.70) at streams. t 13 C DOC values were not signicant ly dierent between streams and F W S W I sites (Tab. 1). However, on average, DOC concentrations were 5 X lower at F W S W I sites (1.0 mg/L; range = 0.5 to 3.1 mg/L) than streams (5.0 mg/L; range = 1.2 to 13.6 mg/L), and 75% of dierent fractions of OM in recharge streams (FPOM, CPOM, DOC, and periphyton) (H3), and a two-sided unpaired Students t-test was used to test for dierenc es in t 13 C of dierent fractions of OM at FW SWI sites (FPOM and DOC) (H4). Stream OM t 13 C values were raised to the 0.3 power for normality. To test for temporal changes in 1) t 13 C FPOM values in both streams and FW SWI sites, 2) C:N ratios in recharge streams, and 3) t 13 C DIC values in FW SWI sites, linear mixed eect models were employed, grouping data by sampling site (H3 and H4). Four recharge streams and three groundwater sites that were each sampled four or more times were used in the analysis. Additional sites were sampled but excluded because of small sample size. C:N ratios were log transformed and adjusted r 2 values were calculated by treating each site-specic regression as a simple linear regression with a single covariate. All statistical analyses were conducted in R v2.15 (R Core Team 2012). Mixed eects models were run us ing the nlme package (Pinheiro et al. 2009). False dis covery rate due to multiple comparisons was controlled by adjusting n using the method of Benjamini & Hoch berg (1995). Sixteen statistical analyses were performed (Tab. 1), and signicance was set to n = 0.03. For clar ity, test statistics are not included in text but are listed in table 1. Tab. 1: Summary of statistical tests of predictions. 1: See text for hypotheses. denotes statistically signicant results. Hypotheses 1 Null predictions Statistical analysis F or t N df p r 2 1 Stream 13 CDIC is not related to location simple linear regression 1.74 30 1 & 28 0.198 0.03 1 Stream 13 CFPOM is not related to location simple linear regression 5.96 26 1 & 24 0.022* 0.17 1 & 3 Stream 13 CFPOM and 13 CDIC are not related simple linear regression 0.62 24 1 & 22 0.439 .02 1 & 2 Stream and FWSWI 13 CDOC does not dier ANCOVA 0.73 38 1 0.399 NA 1 & 2 Stream and FWSWI 13 CFPOM does not dier ANCOVA 10.16 46 1 0.003* NA 2 FWSWI 13 CDIC is not related to location simple linear regression 20.17 26 1&24 <0.001* 0.43 2 FWSWI 13 CFPOM is not related to location simple linear regression 11.19 20 1 & 18 0.004* 0.35 2 FWSWI 13 CFPOM and 13 CDIC are not related simple linear regression 10.42 20 1 & 18 0.005* 0.33 2 FWSWI 13 CFPOM and FPOM C:N are not related simple linear regression 0.41 7 1 & 5 0.552 .11 3 Stream 13 COM fractions do not dier ANOVA 2.66 89 3 & 85 0.053 NA 3 Stream 13 CFPOM does not change with time linear mixed eect model 4.02 19 14 <0.001* 0.01 to 0.98 3 Stream FPOM C:N does not change with time linear mixed eect model 2.72 17 12 0.019* 0.14 to 0.75 3 Stream 13 CFPOM and FPOM C:N are not related simple linear regression 18.70 21 1 & 19 <0.001* 0.47 4 FWSWI 13 CFPOM and 13 CDOC do not dier t test 2.44 44 28.577 0.021* NA 4 FWSWI 13 C FPOM does not change with time linear mixed eect model 4.48 12 8 0.002* .03 to 0.96 4 FWSWI 13 CDIC does not change with time linear mixed eect model 0.16 18 13 0.874 NA B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 251 F W S W I DOC concentrations were below the minimum concentrations of the analytical facilitys calibration stan dards (1.1 mg/L). In streams, t 13 C values for CPOM, FPOM, pe riphyton, and DOC were not signicantly dierent (Tab. 1; Fig. 3), and t13C FPOM values were not correlated with t 13 C DIC values (Fig. 4D). A signicant relationship between t 13 C FPOM values and C:N ratios was observed in streams (Fig. 4A), but no correlations were observed between t 13 C FPOM and ion concentrations or physico chemistry (Pearsons r < 0.2). Stream t 13 C FPOM values increased from the southwest to the northeast (Fig. 4C), but t 13 C DIC values in streams were not correlated with location (Fig. 4B). Visual examination of t 13 C periphyton data did not reveal a spatial pattern, but a relationship was not statistically assessed because, unlike allochtho nous OM, we had no reason a priori to expect the isoto pic composition of periphyton to vary spatially. At FW SWI sites, t 13 C FPOM values were signicantly more negative than t 13 C DOC values by 6.71 (Tab. 1), and a signicant positive relationship between t 13 C FPOM and t 13 C DIC values was observed (Tab. 1; Fig. 4F), with the average t 13 C FPOM value being 28.44 lower than the average t 13 C DIC value. No relationship between t 13 C FPOM values and C:N ratios was observed at F W S W I sites, although sample size was small (Tab. 1; Fig. 4E). At FW SWI sites, strong correlations were observed be tween t 13 C DIC values, conductivity and concentrations of several dissolved ions, including sulde, ammonia, chlo ride, sulfate, lithium, sodium, potassium, magnesium, and calcium (r > 0.70), but not between t 13 C DIC values/ conductivity and other physicochemistry measure ments (pH, temperature, tD, t 18 O, manganese, barium, uoride, nitrite, and nitrate concentrations) (r < 0.5). At FW SWI sites, both t 13 C DIC values (at sites with conduc tivity < 4000S/cm) and t 13 C FPOM values increased from southwest to northeast (Tab. 1; Fig. 4GH). AIC strong ly suggested that a linear model incorporating location, log conductivity, and an interaction term was substan tially more likely than nested models (AIC weight >> 1); all parameters were signicant. Temporal changes in t 13 C FPOM values were ob served at both recharge streams and FW SWI sites. Stream t 13 C FPOM values increased between 30 Septem ber 2010 and 20 March 2012, although the strength of the relationship varied greatly among streams (r 2 = 0.01 to 0.98) (Tab. 1; Fig. 5). is decrease did not corre spond directly with stream discharge or PDSI, as the last two sampling events (late January and late March, 2012) followed precipitations events that resulted in increased ow in all sampled streams (Fig. 2). Visual assessment of CPOM and periphyton isotopic compositions did not indicate temporal patterns. FPOM C:N ratios in streams exhibited a weaker, but signicant, negative re lationship with time (Tab. 1). Unexpectedly, at FW SWI sites, t 13 C FPOM values increased between 16 April 2011 Fig. 3: Boxplots for 13 C values for dierent sampled OM fractions at stream and B. FWSWI sites. ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...

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ACTA CARSOLOGICA 42/2-3 2013 252 Fig. 4: Regressions for stream sites (AD) and FWSWI sites (EH) of A and D: 13 C FPOM values against FPOM C:N ra tios; B and F: 13 C FPOM values against 13 CDIC values; C and G: 13 CFPOM values ver sus distance along the Edwards Aquifer owpath; D: 13 CDIC values versus distance; and H: 13 CDIC values versus distance and conductivity (multiple re gression surface). All isotope val ues are reported in per mil (). Trendlines are shown for signi cant regressions. 13 CDIC= .165 + 0.082 distance + 7.808 log(conductivity) 0.010 dis tance log(conductivity) B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 253 DISCUSSION e distribution and diversity of karst aquifer metazoan communities, as well as aquifer-wide food web structure, are inuenced by OM that originates from, and migrates across, physical and geochemical boundaries. Research on the factors inuencing DOM variability in karst aqui fer systems has been limited, with previous work suggest ing that OM ux into karst groundwater varies tempo rally based on precipitation and OM composition in soil and epikarst dripwaters (van Beynen et al. 2000; Datry et al. 2005; Ban et al. 2008), and the relative contributions of photosynthetic and chemolithoautotrophic OM are spatially variable (Sarbu et al. 1996; Opsahl & Chanton 2006; Birdwell & Engel 2009; Roach et al. 2011; Neisch et al. 2012). e isotopic compositions of OM in Edwards Aquifer recharge and aquifer waters, and the spatial and temporal variability in OM sources into the aquifer, have not been previously assessed, even though they support of one of the richest stygobiont communities on Earth (Longley 1981). W e hypothesized that OM sources in the Edwards Aquifer would be inuenced by 1) the relative proportion of C3 and C4 plant OM in recharging streams that changes in response to an east-west precipitation gra dient, 2) FW SWI t 13 C FPOM values that reect regional dif ferences in t 13 C DIC values due to chemolithoautotrophic production, 3) the importance of OM from periphyton and riparian C3 plants in recharge streams that increases during the dry season, and 4) the constant composition of OM at the FW SWI over time. and 2 April 2012. Again, the strength of the relation ship varied between sites (r 2 = 0.03 to 0.96) (Tab. 1; Fig. 5). t 13 C DIC values at FW SWI sites showed no tem poral trend (Tab. 1). Fig. 5: Results of linear mixed eects models for 13 C-FPOM val ues versus time for stream sites (A-D), and FWSWI sites (E-G). A. Sabinal Rv.; B. H ondo Cr.; C. H elotes Cr.; D. Blanco Rv.; E. Aquarena parking lot well; F. Girl scout deep well; G. P aradise al ley well. All isotope values are reported in per mil (). Note that mixed eect models assess relationships between 13 C-FPOM and date using site as a grouping variable. Site-specic regressions may not be signicant. ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...

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ACTA CARSOLOGICA 42/2-3 2013 254 Two of our hypotheses focus on the contribution of OM from recharge (H1 & H3). Numerous factors inuence the relative importance of allochthonous and autoch thonous OM in streams, including stream width and riparian cover (Vannote et al. 1980), land cover and the quantity of allochthonous input (Beneld 1997), and nu trient availability (Biggs 1995). e importance of these factors in Edwards Aquifer recharge streams is unknown, but the decrease in t 13 C FPOM values and FPOM C:N ratios in recharging streams during the summer of 2011, and the negative relationship between FPOM C:N ratios and t 13 C FPOM values (Figs. 4) may suggest that the observed temporal isotopic shi in FPOM results from a decrease in the relative contribution of allochthonous OM and an increase in the relative contribution of periphyton. is shi could result from decreased allochthonous input (i.e. both C3 and C4 plants), but is not likely to result from increased in-stream productivity because t 13 C FPOM minima values do not occur during spring and summer when periphyton growth is greatest (Finlay & Kendall 2007). Alternatively, the observed temporal pattern is consistent with a decrease in the relative contribution of C4 plants from beyond the riparian zone because of a decline in overland ows and a subsequent increase in allochthonous input from the C3 dominated riparian zone, as has been documented for a river in Cameroon (Bird et al. 1998). Decreasing t 13 C FPOM values in streams continued through increased ows in winter and spring of 2012, although the pattern was inconsistent among streams (Figs. 2 & 5). In particular, t 13 C FPOM values from the Sabinal River changed minimally aer an ini tial decrease aer September 2010. A small increase in t 13 C FPOM values in the Sabinal and Blanco Rivers corre sponded to increased ow in winter and spring of 2012, but this was not observed in Helotes Creek. Spring and stream hydrographs and the PDSI show 2 year oscilla tions with wetter than normal periods corresponding to El Nio periods (Fig. 2), and the general trend of declin ing stream t 13 C FPOM values may be linked to these longer ENSO time-scale trends in stream discharge. e rela tively small increase in discharge in winter and spring of 2012, and the negative trend in t 13 C FPOM values, were embedded within a longer drying trend, as illustrated by the PDSI values from spring of 2011 through December 2012 (Fig. 2). Although the relationship was weak, stream t 13 C FPOM values became more enriched from southwest to northeast (Tab. 1, Fig. 4), which does not support our hypothesis of increasing contributions of C3 plant ma terial in the northeast. Furthermore, the lack of spatial gradients in t 13 C periphyton and t 13 C DIC values, and of a sig nicant regression between t 13 C FPOM and t 13 C DIC values (Tab. 1, Fig. 4), indicates that the observed spatial gra dient in t 13 C FPOM values is not the result of spatial dif ferences in t 13 C periphyton values that would result from re gional dierences in t 13 C DIC values. Rather, the observed gradient may indicate decreasing contributions of per iphyton and increased contributions of terrestrial plant OM in the northeast, although our data do not allow for estimated proportions because of the large degree of overlap in t 13 C CPOM and t 13 C periphyton values. e estimated discharge-weighted average value for t 13 C FPOM entering streams (.75) was similar to the unweighted average, yet the employed Bayesian method of estimation has several advantages to an unweighted analytical average. Most obviously, an unweighted ana lytical average can over-emphasize values collected dur ing low recharge periods and under-emphasize values collected during high recharge periods. Secondly, this method incorporates uncertainty associated with indi vidual isotopic measurements, allowing for calculation of 95% equal tail credible intervals. Finally, although not investigated here, the model has potential for incorpo ration of increased complexity. Specically, the relation ship between the amount of OM entering the aquifer and discharge could be modeled non-linearly (e.g., it may reach an asymptote at some discharge threshold), and the relationship between the amount of OM enter ing the aquifer and discharge could be modelled to vary among streams. CONTRIBUTIONS OF ORGANIC MATTER FROM SURFACE RECHARGE CONTRIBUTIONS OF ORGANIC MATTER FROM THE FW SWI e juxtaposition of reduced electron donors (e.g., H 2 S) and electron acceptors (e.g., O 2 NO 3 ) at the FW SWI, cou pled with a plentiful source of inorganic carbon source (DIC as HCO 3 and CO 2 ) from carbonate dissolution sup port chemolithoautotrophic metabolic processes. Pro nounced dierences in OM dynamics between recharge stream and FW SWI well waters were revealed through isotope analysis. t 13 C FPOM values were signicantly more B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 255 negative at FW SWI sites than in recharging streams (Tab. 1, Fig. 3), which suggests strong isotopic discrimination against 13 C during autotrophic C xation. W e hypoth esized that chemolithoautotrophic production occurred along the FW SWI, based on identication of putative sulfur-oxidizing microbial groups from the FW SWI (e.g., Epsilonproteobacteria, iothrix spp., iobacillus spp.) (Engel & Randall 2011; Gray & Engel 2013). e results support our hypothesis, and also corroborate previous ndings that microbial, rather than surface (i.e. plant), humic-like, OM is present at the FW SWI (Birdwell & Engel 2009). However, the positive relationship between t 13 C FPOM and t 13 C DIC at FW SWI sites (Fig. 4) illustrates that C isotope data alone are insucient to quantify the relative proportions of photosynthetic and chemolitho autotrophic OM in a sample. Several factors inuence the isotopic composition of chemolithoautotrophic OM, in cluding the isotopic signature of the C substrate, C limita tion (Cowie et al. 2009), C xation rate (Laws et al. 1995), and the C xation pathway utilized (Berg et al. 2010). e 6.71 dierence between t 13 C FPOM and t 13 C DOC values at FW SWI sites could be the result of several processes. Relative to DOM, POM is not trans ported great distances into groundwater systems (Si mon et al. 2003), so FW SWI DOC may be comprised of a greater proportion of surface derived, photosynthetic OM. Alternatively, DOC may represent more processed or recalcitrant OM. In soils, preferential metabolism of 12 C in OM during decomposition can increase t 13 C OM by 1 (Bostrm et al. 2007). To our knowledge, howev er, this has not been documented for groundwater. Last ly, the values may suggest additional C assimilation due to methanotrophy (W hiticar 1999). Although analysis of the spatial distribution of CH 4 in the Edwards Aquifer saline zone has not been studied in detail, Zhang et al. (1998) report a positive relationship between CH 4 and SO 4 2 concentrations in the saline zone, and we cannot rule out regional dierences in CH 4 concentration that could inuence regional variability in FW SWI t 13 C OM Reasons for the enrichment in FW SWI t 13 C FPOM values between April 2011 and March 2012 (Fig. 5) are unclear, but heterotrophic processing of OM is insuf cient to account for the observed isotopic dierences, as much as 18. e observed temporal changes could be the result of changing contributions of OM produced in geochemically distinct portions of the aquifer. ese changes may be the result of declining aquifer levels and variability in ow along the FW SWI due to drought; however, there are currently no data to support this hy pothesis. e signicant positive relationship between t 13 C DIC and t 13 C FPOM values (Tab. 1, Figs. 4) supports our prediction that regional dierences in t 13 C DIC values, the substrate for chemolithoautotrophic production, aect t 13 C FPOM values. t 13 C DIC values at FW SWI sites increased from southwest to northeast for sites with conductiv ity < 4000 S/cm (Fig. 4H) and showed no signicant temporal variation. is trend mirrors patterns in stable isotopes of helium (Hunt et al. 2010), which were inter preted as evidence of increasing groundwater residence times from the southwest to northeast. Increased resi dence times, and subsequent increased time for rock-wa ter interaction, can shi t 13 C DIC values towards the isoto pic composition of the host rock (~ for Edwards carbonates) (Ellis 1985; Gonantini & Zuppi 2003). Variable sources of acidity may also have an impor tant role in the isotopic composition of FW SWI t 13 C DIC values. Dissolution of calcite by carbonic acid (derived from CO 2 respired during decomposition of plant mat ter in soils and/or hyporheic zones) results in DIC with a t 13 C value intermediate between that of the calcite and carbonic acid (Finlay 2003; Breeker et al. 2012). Disso lution of Edwards limestones by carbonic acid derived from respired CO 2 will produce DIC with t 13 C values ~ for calcite-saturated water, as well as alkalini ties and Ca 2+ concentrations similar to those observed in surface streams (appendix 1). However, DIC sourced from dissolution of calcite by an acid other than carbon ic acid (e.g., sulfuric acid) will have an isotopic compo sition closer to that of the host rock, as is observed in FW SWI sites. e strong positive relationship between FW SWI t 13 C DIC conductivity (Fig. 4H), and sulde, sup ports the hypothesis that dissolution in low conductiv ity freshwaters is driven by carbonic acid, and in high conductivity saline waters (with locally high levels of sulde ~100 mg/L), dissolution is driven by sulfuric acid derived from microbially mediated oxidation of reduced sulfur compounds (Gray & Engel 2013). CONCLUSION Groundwater ecosystems can be supported, at least in part, by allochthonous OM input from the surface, whereby the composition and supply rate is temporally variable and dependent on water balance conditions at ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...

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ACTA CARSOLOGICA 42/2-3 2013 256 the surface. In streams supplying allochthonous OM to the Edwards Aquifer, t 13 C FPOM values and FPOM C:N ra tios decreased during and aer a severe drought in 2011, suggesting a diminished contribution of terrestrial plant material (especially C4 material from beyond the ripar ian zone) and an increasing contribution of in-stream production. A spatial gradient in stream t 13 C FPOM values due to changes in the relative abundance of C3 and C4 plants was not apparent. W eighting t 13 C FPOM values for FPOM input into aquifers by recharge provides a more re alistic estimate, and quantication of uncertainty around estimates is both important and straightforward using a Bayesian approach. In addition to allochthonous inputs, chemolithoautotrophy along the FW SWI is an important source of autochthonous OM, based on geochemical, microbial, and isotopic evidence. For the Edwards Aqui fer, allochthonous and autochthonous OM sources were, on average, isotopically distinct, although the isotopic composition of chemolithoautotrophic OM was spatially variable and dependent on the isotopic composition of DIC. Additional research is needed to understand the degree to which the distinct OM sources are utilized by the diverse microbial and metazoan community in the Edwards Aquifer, as well as to characterize the OM geo chemically (e.g., through high-resolution spatial analyses of OM isotopic composition and degree of humication). ese analyses would allow for better quantication of the relative proportions of allochthonous and autoch thonous OM throughout the aquifer. ACKNO WLEDGEMENTS is research was funded by the National Science Foun dation (#0742306; 1110503) the U.S. Geological Survey (#9658), and the Jones Endowment for Aqueous Ge ochemistry at the University of Tennessee. e San Anto nio W ater System, e Edwards Aquifer Authority, Zara Environmental, and several private landowners provided or facilitated access to sampling sites. Kevin Simon and an anonymous reviewer provided valuable comments during the revision of this manuscript. REFERENCES Artman, U., W aringer, J.A. & M. Schagerl, 2003: Seasonal dynamics of algal biomass and allochthonous input of coarse particulate organic matter in a low-order sandstone stream (W eidlingbach, Lower Austria).Limnologica, 33, 77. Ban, F., Pan, G., Zhu, J., Cai, B. & M. Tan, 2008: Tem poral and spatial variations in the discharge and dissolved organic carbon of drip waters in Beijing Shihua Cave, China.Hydrological Processes, 22, 3749. Barker, R.A., Bush, P.W & E. T. Baker, Jr., 1994: Geologic History and Hydrogeologic Setting of the EdwardsTrinity Aquifer System, W est-Central Texas.Unit ed States Geological Survey, Report number: W ater Resources Investigations Report 94. Beneld, E.F., 1997: Comparison of litterfall input to streams.Journal of the North American Bentho logical Society, 16, 104. Benjamini, Y. & Y. Hochberg, 1995: Controlling the False Discovery Rate: A Practical and Powerful Approach to Multiple Testing.Journal of the Royal Statistical Society, 57, 289. Berg, I. A., Kockelkorn, D., Ramos-Vera, W .H., Say, R. F., Zarzycki, J., Hgler, M., Alber, B.E. & G. Fuchs, 2010: Autotrophic carbon xation in Archaea.Na ture Reviews Microbiology, 8: 447. Biggs, B.J.F., 1995: e Contribution of Flood Distur bance, Catchment Geology and Land Use to the Habitat Template of Periphyton in Stream Ecosys tems.Freshwater Biology, 33, 419. Bird, M.I., Giresse, P. & S. Ngos, 1998: A seasonal cycle in the carbon-isotope composition of organic car bon in the Sanaga River, Cameroon.Limnology and Oceanography, 43, 143. B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 257 Birdwell, J.E. & A.S. Engel, 2009: Variability in Terres trial and Microbial Contributions to Dissolved Or ganic Matter Fluorescence in the Edwards Aquifer, Central Texas.Journal of Cave and Karst Studies, 71, 144. Bostrm, B., Comstedt, D. & A. Ekblad, 2007: Isotope Fractionation and 13 C Enrichment in Soil Proles During the Decomposition of Soil Organic Matter.Oecologia, 153, 89. Breeker, D.O., Payne, A.E., Q uade, J., Banner, J.L., Ball, C.E., Meyer, K.W & B.D. Cowan, 2012: e sources and sinks of CO 2 in caves under mixed woodland and grassland vegetation Geochimica et Cos mochimica Acta.96, 230. Bukovinskey, T., Verschoor, A.M., Helmsing, N.R., Beze mer, T.M. Bakker, E.S., Vos, M. & L.N. de Senerpont Domis, 2012: e Good, the Bad and the Plenty: In teractive Eects of Food Q uality and Q uantity on the Growth of Dierent Daphnia Species.PLOS ONE, 7, e42966. Cowie, B.R., Slater, G.F., Bernier, L. & L.A. W arren, 2009: Carbon Isotope Fractionation in Phospholipid Fatty Acid Biomarkers of Bacteria and Fungi Native to an Acid Mine Drainage Lake.Organic Geochemistry, 40, 956. Culver, D.C. & T. Pipan, 2009: e Biology of Caves and Other Subterranean H abitats. e Biology of Habi tats, Oxford University Press, pp. 254, Oxford. Datry, T., Malard, F. & J. Gibert, 2005: Response of in vertebrate assemblages to increased groundwa ter recharge rates in a phreatic aquifer.Journal of the North American Benthological Society, 24: 461. Doctor, D.H., Kendall, C., Sebestyen, S.D., Shanley, J.B., Ohte, N. & E.W Boyer, 2008: Carbon Isotope Frac tionation of dissolved Inorganic Carbon (DIC) Due to Outgassing of Carbon Dioxide from a Headwater Stream.Hydrological Processes, 22, 2410. Ellis, P. M., 1985: Diagenesis of the Lower Cretaceous Ed wards Group in the Balcones Fault Z one Area, SouthCentral Texas.PhD thesis, University of Texas at Austin, pp. 326. Engel, A.S. & K.W Randall, 2011: Experimental Evi dence for Microbially Mediated Carbonate Disso lution from the Saline W ater Zone of the Edwards Aquifer, Central Texas.Geomicrobiology Journal, 28: 313. Finlay, J.C., 2003: Controls of Streamwater Dissolved Inorganic Carbon Dynamics in a Forested W ater shed.Biogeochemistry, 62, 231. Finlay, J.C. & C. Kendall, 2007: Stable Isotope Tracing of Temporal and Spatial Variability in Organic Matter Sources to Freshwater Ecosystems.In: Michener, R. & K. Lajtha (eds.) Stable Isotopes in Ecology and Environmental Science, Blackwell Publishing, pp. 283, Malden. Gary, M.O., Gary, R.H. & B.B. Hunt (eds.), 2011: Inter connection of the Trinity (Glen Rose) and Edwards Aquifers Along the Balcones Fault Z one and Related Topics, Karst Conservation Initiative February 17, 2011 Meeting P roceedings, pp. 46, Austin. Gelman, A. & D.B. Rubin, 1992: Inference from Iterative Simulation Using Multiple Sequences.Statistical Science, 7, 457. Gonantini, R. & G.M. Zuppi, 2003: Carbon Isotope Ex change Rate of DIC in Karst Groundwater.Chemi cal Geology, 197, 319. Gray, C.J. & A.S. Engel, 2012: Microbial Diversity and Impact on Carbonate Geochemistry Across a Changing Geochemical Gradient in a Karst Aqui fer.e ISME Journal, advance online publication: 1. Humphreys, W ., Tetu, S., Elbourne, L., Gillings, M., Sey mour, J., Mitchell, J. & I. Paulsen, 2012: Geochemi cal and Microbial Diversity of Bundera Sinkhole, an Anchialine System in the Eastern Indian Ocean.Natura Croatica, 21, 59. Hunt, A.G., Lambert, R.B., Fahlquist, L. 2010: Sources of Groundwater Based on H elium Analyses in and Near the Freshwater/ Saline-Water Transition Z one of the San Antonio Segment of the Edwards Aquifer, SouthCentral Texas, 2002.United States Geological Survey, Report Number: Scientic Investigations Report 2010 Jasinska, E.J., Knott, B., McComb, A.J. 1996: Root Mats in Ground W ater: A Fauna-Rich Cave Habitat.Journal of the North American Benthological Soci ety, 15, 508. Johnson, S., Schindel, G., & J. Hoyt, 2009: Water Qual ity Trends Analysis of the San Antonio Segment, Bal cones Fault Z one Edwards Aquifer, Texas. Edwards Aquifer Authority, Report Number: 09. Laws, E.A., Popp, B.N., Bidigare, R.R., Kennicutt, M.C. & S.A. Macko, 1995: Dependence of Phytoplank ton Carbon Isotopic Composition on Growth Rate and [CO 2 ] aq : eoretical Considerations and Ex perimental Results.Geochimica et Cosmochimica Acta, 59, 1131. ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...

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ACTA CARSOLOGICA 42/2-3 2013 258 Lindgren, R. J., Dutton, A. R., Hovorka, S. D., W orthing ton, S. R. H. & S. Painter, 2004: Conceptualization and Simulation of the Edwards Aquifer, San Antonio Region, Texas.United States Geological Survey, Report Number: Scientic Investigations Report 2004. Neisch, J.A., Pohlman, J.W & T.M. Ilie, 2012: e use of stable and radiocarbon isotopes as a method for de lineating sources of organic material in anchialine systems.Natura Croatica, 21, 83. Nielsen-Gammon, J.W ., 2008: e Changing Climate of Texas.In: Schmandt, J., Clarkson, J. & G.R. North (eds.) e Impact of Global Warming on Texas, 2 nd edition. University of Texas Press, pp. 39, Aus tin. Oetting, G.C., Banner, J.L. & J.M. Sharp, Jr., 1996: Re gional Controls on the Geochemical Evolution of Saline Groundwaters in the Edwards Aquifer, Cen tral Texas.Journal of Hydrology, 181, 251. Opsahl, S.P. & J.P. Chanton, 2006: Isotopic evidence for methane-based chemosynthesis in the Upper Flori dan aquifer food web.Oecologia, 150, 89. Palmer, W .C., 1965: Meteorologic drought.U.S. W eath er Bureau, Resource Paper 45. Perez, R., 1986: P otential for Updip Movement of Saline water in the Edwards Aquifer, San Antonio, Texas.United States Geological Survey, Report Number: W ater-Resources Investigations Report 86. Pinheiro, J., Bates, D., DebRoy, S. & D. Sarkar, 2009: Lin ear and Nonlinear Mixed Eects Models. R package version 3,1. R Foundation for Statistical Com puting, Vienna. Plummer, M., 2010: Bayesian Graphical Models Using MCMC.R package version 2.1, 0. Pohlman, J.W ., Ilie, T.M. & L.A. Cifuentes, 1997: A Stable Isotope Study of Organic Cycling and the Ecology of an Anchialine Cave Ecosystem.Marine Ecology Progress Series, 155, 17. Poulson, T.L., 2005: Food Sources.In: Culver, D.C. & W .B. W hite (eds.) Encyclopedia of Caves Elsevier Academic Press, pp. 255, Amsterdam. Poulson, T.L. & K.H. Lavoie, 2000: e Trophic Basis of Subsurface Ecosystems.In: W ilkens, H., D. C. Cul ver & W F. Humphreys (eds.) Subterranean Ecosys tems. Ecosystems of the World 30 Elsevier Academic Press, pp. 231, Amsterdam. R Core Team, 2012: R: A Language and Environment for Statistical Computing.R Foundation for Statistical Computing, Vienna. Roach, K.A., Tobler, M. & K.O. W inemiller, 2011: Hy drogen sulde, bacteria, and sh: a unique, subter ranean food chain.Ecology, 92, 2056. Saito, L., Redd, C., Chandra, S., Atwell, L., Fritsen, C. H. & M. R. Rosen, 2007: Q uantifying Foodweb Inter actions with Simultaneous Linear Equations: Stable Isotope Models of the Truckee River, USA.Journal of the North American Benthological Society, 26, 642. Sarbu, S.M., Kane, T.C. & B. K. Kinkle, 1996: A Chemoau totrophically Based Cave Ecosystem.Science, 272, 1953. Short, R.A., Smith, S.L., Guthrie, D.W ., Stanford, J.A. 1984: Leaf litter processing rates in four Tex as streams.Journal of Freshwater Ecology, 2, 469. Silva, M.S., de Oliveira Bernardi, L.F., Martins, R.P. & R.L. Ferreira, 2012: Transport and Consumption of Organic Detritus in a Neotropical Limestone Cave.Acta Carsologica, 41, 139. Simon, K.S., Beneld, E.F. & S.A. Macko, 2003: Food W eb Structure and the Role of Epilithic Biolms in Cave Streams.Ecology, 84, 2395. Simon, K.S., Pipan, T. & D.C. Culver, 2007: A Concep tual Model of the Flow and Distribution of Organic Carbon in Caves.Journal of Cave and Karst Stud ies, 69, 279. Stribling, J.M. & J.C. Cornwell, 1997: Identication of Important Primary Producers in a Chesapeake Bay Tidal Creek System Using Stable Isotopes of Carbon and Sulfur.Estuaries, 20, 77. van Beynen, P. v., Ford, D. & H. Schwarcz, 2000: Seasonal variability in organic substances in surface and cave waters at Marengo Cave, Indiana.Hydrological Processes, 14, 1177. van Dover, C.L, 2007: Stable isotope studies in marine chemoautotrophically based ecosystems: an up date.In: Michener, R. & K. Lajtha (eds.) Stable Iso topes in Ecology and Environmental Science, Black well P ublishing, pp. 202, Malden. Vannote, R.L., Minshall, G.W ., Cummins, K.W ., Sedell, J.R. & C.E. Cushing, 1980: e River Continuum Concept.Canadian Journal of Fisheries and Aquat ic Science, 37, 130. W hiticar, M.J., 1999: Carbon and Hydrogen Isotope Sys tematics of Bacterial Formation and Oxidation of Methane.Chemical Geology, 161, 291. Zhang, C., Grossman, E.L. & J.W Ammerman, 1998: Factors Inuencing Methane Distribution in Texas Ground W ater.Ground W ater, 36, 58. B ENJAMIN T. HUTCHINS, B ENJAMIN F. SCHW ARTZ & ANNETTE S. ENGEL

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ACTA CARSOLOGICA 42/2-3 2013 259 Appendix 1: P hysicochemical data average values (min/max) for recharge stream and FWSWI sites. Recharge Streams Sab Hon Med Hel Gua Bla Oni DO (mg/L) 8.29 (5.50/10.65) 8.98 (8.22/9.59) 8.00 (6.58/9.26) 8.41 (6.71/10.90) 9.14 (na) 9.76 (7.87/12.45) 9.19 (7.25/10.97) pH 7.81 (7.05/8.37) 7.86 (7.46/8.11) 7.51 (7.13/8.05) 7.68 (6.95/8.14) 7.52 (na) 8.16 (7.40/9.44) 7.33 (7.23/7.54) Cond (S/cm) 405 (234/548) 312 (297/320) 411 (348/484) 538 (434/693) 349 (na) 432 (358/504) 445 (350/557) T (C) 16.42 (10.02/26.21) 10.69 (7.88/15.99) 18.20 (13.24/24.46) 16.69 (13.12/22.35) 8.17 (na) 14.74 (3.68/21.16) 13.26 (6.45/21.81) 13 C DIC () .92 (.56/.44) .55 (.89/.21) .67 (.34/.87) .25 (.95/.03) .55 (na) .00 (.98/.20) .88 (.40/.15) 13 C DOC () .0 (.0/.4) .7 (na) nm (nm) .9 (.6/.2) nm (nm) .1 (.1/.4) nm (nm) D () .59 (.76/.63) .23 (.40/.83) .34 (.38/.39) .04 (.43/.71) .11 (na) .66 (.25/.54) .63 (.98/.68) 18 O () .86 (.41/.79) .26 (.87/.60) .13 (.73/.53) .03 (.95/0.20) .81 (na) .97 (.32/.38) .73 (.47/.79) Sulde (ppm S) 0.05 (0.04/0.06) 0.05 (na) nm (nm) 0.05 (0.05/0.07) nm (nm) 0.08 (0.05/0.16) nm (nm) Fl (ppm) 0.16 (0.15/0.20) 0.20 (0.19/0.23) 0.18 (0.17/0.19) 0.11 (0.09/0.12) 0.24 (na) 0.21 (0.16/0.30) 0.20 (0.19/0.22) Cl (ppm) 15.38 (6.09/20.54) 14.44 (12.40/15.87) 12.97 (11.90/13.62) 46.80 (15.07/76.14) 26.70 (na) 14.93 (13.78/15.45) 24.92 (21.07/27.66) NO 2 (ppm) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (na) 0.00 (0.00/0.00) 0.00 (0.00/0.00) Br (ppm) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (na) 0.00 (0.00/0.00) 0.00 (0.00/0.00) NO 3 (ppm) 0.89 (0.00/3.12) 0.76 (0.43/1.27) 1.52 (1.14/2.12) 2.10 (0.00/8.04) 2.67 (na) 2.72 (0.00/8.58) 0.44 (0.00/1.17) PO 4 3 (ppm) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (na) 0.00 (0.00/0.00) 0.00 (0.00/0.00) SO 4 2 (ppm) 51.53 (29.21/107.6) 65.24 (52.95/76.78) 57.58 (53.80/60.20) 63.08 (21.87/115.9) 29.61 (na) 37.20 (26.07/56.65) 49.65 (40.34/55.73) Alkalinity (ppm) 233.29 (152.79/295.72) 88.71 (na) nm (nm) 194.68 (192.22/197.15) 256.29 (na) 156.07 (88.72/202.08) nm (nm) Li + (ppm) 0.00 (0.00/0.01) 0.00 (0.00/0.01) 0.00 (0.00/0.00) 0.00 (0.00/0.01) 0.00 (na) 0.00 (0.00/0.01) 0.00 (0.00/0.00) Na + (ppm) 8.79 (5.51/10.77) 10.55 (9.97/11.56) 7.77 (7.37/8.15) 19.17 (8.16/26.14) 15.03 (na) 7.48 (5.51/8.74) 10.08 (9.17/11.33) NH 4 + (ppm) 0.01 (0.00/0.09) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (na) 0.00 (0.00/0.03) 0.00 (0.00/0.00) Ka + (ppm) 1.29 (0.77/2.89) 1.45 (1.04/2.20) 1.46 (1.36/1.52) 1.42 (0.72/2.10) 1.54 (na) 1.34 (1.06/1.47) 1.11 (0.86/1.36) Mg 2+ (ppm) 12.00 (5.56/18.32) 8.77 (7.57/10.00) 12.01 (11.38/13.10) 15.40 (10.05/19.63) 27.53 (na) 16.92 (13.01/22.65) 15.95 (12.26/18.73) Mn 2+ (ppm) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.00 (na) 0.00 (0.00/0.00) 0.00 (0.00/0.00) Ca 2+ (ppm) 62.28 (37.27/79.76) 55.77 (46.38/60.60) 60.58 (54.22/63.83) 63.41 (50.99/81.83) 61.47 (na) 51.87 (35.21/69.78) 73.08 (69.24/76.76) Sr 2+ (ppm) 0.58 (0.00/2.04) 0.25 (0.00/0.75) 0.00 (0.00/0.00) 0.33 (0.00/0.99) 0.00 (na) 0.67 (0.00/2.26) 0.00 (0.00/0.00) Ba 2+ (ppm) 0.07 (0.00/0.36) 0.00 (0.00/0.00) 0.00 (0.00/0.00) 0.01 (0.00/0.07) 0.00 (na) 0.09 (0.00/0.31) 0.00 (0.00/0.00) ENVIRONMENTAL CONTROLS ON ORGANIC MATTER PRODUCTION AND TRANSPORT ACROSS SURFACESUBSURFACE ...



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USING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN A CAVE SPRING NORTH W ESTERN A RKANSAS USA UPORABA IZOTOPOV RAZTOPLJENEGA ANORGANSKEGA OGLJIKA V VODI ZA LO EVANJE VIROV NAPAJANJA IZVIRNE JAME, SEVEROZAHODNI ARKANSAS, ZDA Katherine J. KNIERIM 1* Erik POLLOCK 2 & Phillip D. H AYS 3 1 Environmental Dynamics Program, University of Arkansas, Fayetteville, AR, (p) 479, (f) 479, e-mal: kknierim@uark.edu 2 University of Arkansas Stable Isotope Lab, University of Arkansas, Fayetteville, AR, (f) 479, e-mal: epolloc@uark.edu 3 U.S. Geological Survey, Arkansas W ater Science Center, Fayetteville, AR; Department of Geosciences, University of Arkansas, Fayetteville, AR, (f) 479, e-mal: pdhays@usgs.gov corresponding author Received/Prejeto: 15.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 261, POSTOJNA 2013 Abstract UDC 551.44:546.26.027*14(736.7) Katherine J. Knierim, Erik Pollock & Phillip D. Hays : Using isotopes of dissolved inorganic carbon species and water to separate sources of recharge in a cave spring, northwestern Arkansas, USA Blowing Spring Cave in northwestern Arkansas is represen tative of cave systems in the karst of the Ozark Plateaus, and stable isotopes of water (t 18 O and t 2 H) and inorganic carbon (t 13 C) were used to quantify soil-water, bedrock-matrix wa ter, and precipitation contributions to cave-spring ow during storm events to understand controls on cave water quality. W a ter samples from recharge-zone soils and the cave were collect ed from March to May 2012 to implement a multicomponent hydrograph separation approach using t 18 O and t 2 H of water and dissolved inorganic carbon (t 13 CDIC). During baseow, median t 2 H and t 18 O compositions were .6 and .2 for soil water and were .2 and .9 for cave water, re spectively. Median DIC concentrations for soil and cave wa ters were 1.8 mg/L and 25.0 mg/L, respectively, and median t 13 CDIC compositions were .9 and .3, respec tively. During a March storm event, 12.2 cm of precipitation fell over 82 h and discharge increased from 0.01 to 0.59 m 3 /s. e isotopic composition of precipitation varied through out the storm event because of rainout, a change of 50 and 10 for t 2 H and t 18 O was observed, respectively. Although, at the spring, t 2 H and t 18 O only changed by approximately 3 and 1, respectively. e isotopic compositions of pre cipitation and pre-event (i.e., soil and bedrock matrix) water were isotopically similar and the two-component hydrograph separation was inaccurate, either overestimating (>100%) or underestimating (<0%) the precipitation contribution to the spring. During the storm event, spring DIC and t 13 CDIC de Izvleek UDK 551.44:546.26.027*14(736.7) Katherine J. Knierim, Erik Pollock & Phillip D. Hays : Upo raba izotopov raztopljenega anorganskega ogljika v vodi za loevanje virov napajanja izvirne jame, severozahodni Ar kansas, ZDA Jama Blowing Spring Cave v severozahodnem Arkansasu je tipini jamski sistem krasa na planoti Ozark. Stabilni izotopi vode (t 18 O in t 2 H) in anorganskega ogljika (t 13 C) so bili upo rabljeni za ovrednotenje, koliko med padavinskimi dogodki k pretoku izvira prispevajo voda iz prsti, kamnine in neposredne padavine, in za bolje razumevanje kakovosti izvirske vode. Za vekomponentni pristop loevanja hidrograma z t 18 O in t2H vode in raztopljenega anorganskega ogljika (t 13 CDIC) so bili med marcem in majem 2012 vzeti vzorci vode iz prsti in jame. Med bazinim tokom je bila mediana t2H in t18O kom pozicije ,6 in ,2 za vodo iz prsti ter ,2 in ,9 za jamsko vodo. Mediani DIC koncentracij sta bili 1,8 mg/l in 25,0 mg/L ter mediani kompozicije t 13 CDIC ,9 in ,3 Med nevihtnim dogodkom v marcu 2012, ko je v 82 urah padlo 12,2 cm padavin in se je pretok na izviru poveal z 0,01 na 0,59 m 3 /s, se je izotopska sestava t 2 H in t 18 O v pa davinah spreminjala za 50 oziroma 10 Kljub temu pa se je na izviru izotopska sestava t 2 H in t 18 O spreminjala le za priblino 3 oziroma 1 Izotopske sestava padavin in vode iz prsti in kamnine pred nevihtnim dogodkom je bila podobna, zato je bilo izotopsko loevanje dveh komponent hidrograma netono in je bodisi precenjevalo (> 100 %) ali podcenjevalo (< 0%) prispevek padavine k izvirski vodi. Med nevihtnim do godkom so se vrednosti DIC in t 13 CDIC na izviru zmanjale na najmanj 8,6 mg/l in ,2 Ob predpostavki, da je prispe vek padavin nien, je bilo ugotovljeno, da je prispevek vode iz prsti med 23 % pretoka na izviru. eprav je predpostavka

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ACTA CARSOLOGICA 42/2-3 2013 262 K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS creased to a minimum of 8.6 mg/L and .2, respectively. If the contribution from precipitation was assumed to be zero, soil water was found to contribute between 23 to 72% of the total volume of discharge. Although the assumption of negli gible contributions from precipitation is unrealistic, especially in karst systems where rapid ow through conduits occurs, the hydrograph separation using inorganic carbon highlights the importance of considering vadose-zone soil water when ana lyzing storm chemohydrographs. Keywords : carbon, stable isotopes, cave, hydrograph, Arkansas. o zanemarljivih prispevkih padavin nerealna, zlasti v krakih sistemih, kjer prihaja do hitrega pretoka skozi kanale, loitev hidrograma z anorganskim ogljikom poudarja pomembnost upotevanja vode iz vadozne cone in prsti pri analizi kemo hidrogramov med padavinskimi dogodki. Kljune besede: ogljik, stabilni izotopi, jama, hidrogram, Ar kansas. I NTRODUCTION Blowing Spring is the focal point of a park located in Bella Vista, Arkansas, which lies on the Springeld Plateau in the Ozark Plateaus (Fig. 1). e spring has experienced de graded water quality since the 1990s, including transient, elevated E. coli levels with nitrate and chloride concentra tions increasing over time. e spring discharges from a cave in the Boone Formation, a Mississippian-aged lime stone with up to 50% chert that hosts abundant karst fea tures including caves, springs, and dissolution-enlarged fractures and conduits (Adamski et al. 1995). e pro posed research aimed to quantify sources of water to the cave stream across the range of hydrologic conditions to assess contaminant eects (bacteria, organic carbon, nu trients) on cave and spring water quality. Much research has focused on water quality at springs, but for this study sampling soil above the cave provided access to vadose zone groundwater and sampling within the cave allowed for direct sampling of bedrock matrix waters, enabling a more complete understanding of water-quality controls. is labor-intensive method monitored karst recharge as the water traveled from the atmosphere, through the soil and epikarst zones, into karst conduits, and discharged at the spring providing a more thorough assessment of geochemical evolution along groundwater ow paths. Stable isotopes are a valuable tool for characterizing ow paths and biogeochemical processing in anisotropic karst Fig. 1: Blowing Springs Cave (star) is located in northwestern Arkansas on the Springeld P la teau, which is one of the Ozark P lateaus. e cave stream begins at a sump near sampling site BS07 and ows to the spring at BS01. Site BS06 is in a side pas sage where drip water collects in a pool (approximately 3 m by 3 m wide and 0.2 m deep). Land use above the cave is mixed forest and suburban development. Cave mapping completed by the Boston Mountain Grotto.

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ACTA CARSOLOGICA 42/2-3 2013 263 STUDY S ITE e Ozark Plateaus, which include the Boston Mountains, Springeld Plateau, and Salem Plateau, extend through Missouri, Kansas, Oklahoma, and Arkansas and are one of the major karst terrains in North America (Adamski et al. 1995; Fig. 1). e karst landscape in the Ozarks of northwestern Arkansas is characterized by chert regolith mantle overlying Paleozoic limestone, dolomite, sand stone, and shale. Orthogonal fracture sets in the carbon ate rocks provide pathways for water migration, causing dissolution-enlarged fractures, conduits, caves, sinking streams, and sinkholes and creating a landscape with di rect connections between surface water and groundwater U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ... systems (Panno et al. 2001; Trek et al. 2006), and this research applied stable isotope techniques to caves. Stable isotopes of water (t 2 H and t 18 O) are com monly used as natural tracers because (in the absence of evaporation) the isotopes behave conservatively, meaning that changes in the isotopic composition of water are due to mixing of water sources, as opposed to biogeochemi cal reactions (Sklash et al. 1976). Marked dierences in individual storm isotopic compositions imparted by seasonal variation in temperature, humidity, vapor sources, storm tracks, rainout, and other factors (Clark & Fritz 1997; Gibson et al. 2008; Harvey 2001) oen cause storm-event water to be isotopically distinct from water stored in the recharge zone, because mixing in the phre atic zone dampens the isotopic signal of individual storm events providing a time-averaged composition for the stored water (Buttle 1994). Characterizing local precipita tion is important in groundwater studies because the iso topic composition of precipitation varies over space and time, and precipitation is ultimately the water source for most aquifers (Harvey 2001; Simpkins 1995). Hydrograph separations are mixing models used to quantify source-water contributions to stream and spring ow during storm events and generally include three components; precipitation, soil water, and ground water (Lee & Krothe 2003; Ogunkoya & Jenkins 1993; Rice & Hornberger 1998). Hydrograph separations that use a combination of conservative (t 2 H and t 18 O) and non-conservative (t 13 C) tracers can separate water stored in the recharge zone prior to a storm event (preevent water) from water delivered during a storm event (event water) and account for changes in water chemistry along ow paths (Kendall et al. 2001; Sklash & Farvolden 1979). Most mixing models do not adequately account for hydrodynamic dispersion during water transport, but some amount of mixing between sources must occur along the ow path (Jones et al. 2006); therefore, hydro graph separations should not be considered to discretely separate water sources that are conservatively partitioned between zones, but reect the concomitant eects of wa ter movement from and through those zones. In karst settings, hydrograph separations have been applied to quantify the proportion of quick ow (repre sented by precipitation) entering springs during storm events (Lakey & Krothe 1996; Lee & Krothe 2001; Lee & Krothe 2003; Long 2009). Surface features such as sink holes and losing stream segments and bedrock charac teristics such as fractures and dissolution-enlarged con duits allow surface water to rapidly enter the subsurface. Although the preferential pathways allow for rapid inl tration, vadose water (either soil or epikarst) still contrib utes substantial volumes (>50%) to storm ow discharge (Doctor et al. 2006; Lakey & Krothe 1996; Lee & Krothe 2001; Lee & Krothe 2003; Trek et al. 2006). According to Lakey and Krothe (1996), the delivery mechanism for this water is rapid displacement of water within or in direct contact with conduits, so that pre-event water is quickly transported to karst springs. ese techniques have focused on storm-ow hydrographs from springs (Lee & Krothe 2001) and only limited work interpret ing the eects of conduit geometry on water geochemis try has been completed either at springs (Luhman et al. 2012) or within the conduit of a cave (Raeisi et al. 2007). In this research, a three-component hydrograph separation (Lee & Krothe 2001) was completed using t 18 O and t 2 H of water and the concentration and isotopic composition of dissolved inorganic carbon (t 13 CDIC) to separate the contributions of precipitation (Q R ), soil or vadose water (Q S ), and bedrock matrix or groundwa ter (Q B ) to the cave stream during storm events. In this conceptual model, soil water represents the vadose zone (not separately accounting for epikarst water) and the cave stream at baseow represents bedrock matrix water. e epikarst is a very important zone for water storage and biogeochemical processing in karst settings (Peter son et al. 2002; Laincz 2011), but because n tracers can separate n + 1 sources, the addition of epikarst source water would require an additional tracer. By quantifying the discharge of the cave stream, the ux of contaminants can be calculated and related directly to the proportions of quick-ow (Q R ) and diuse ow (Q S and Q B ).

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ACTA CARSOLOGICA 42/2-3 2013 264 METHODS SAMPLE C OLLECTION AND ANALYSIS A stage-discharge relation was developed using stage readings from a 90 V-notch weir constructed 50 m downstream from the spring and accuracy checked us ing the cross-section method (Rantz 1982) and a MarchMcBirney Flo-Mate 2000. Continuous stage and tem perature readings were recorded using HOBO U20 transducers located in the weir pool and at the spring. W eather conditions were recorded using a HOBO Micro Station with barometric pressure and temperature sen sors and a tipping-bucket rain gauge, located 1 km west of the spring. Collection of stage, discharge, and weather data began in February 2012. W ater samples were collected from two lysimeters (L1 and L2) installed in the soil zone above the cave (ap proximately 0.75 m depth), three locations in the cave (BS03, BS06, BS07), and the spring (BS01). e bound aries of the recharge area for the cave stream are not dened, so the dry, headwater valleys above the cave provided the best means to sample soil water that inl trates through the regolith and enters Blowing Spring Cave as dripwater. Sampling goals included bi-monthly sampling (i.e., every other week) to monitor background (baseow) conditions for soil and cave waters and more frequent storm-event sampling to quantify dierences in baseow versus storm-event geochemistry. Preliminary, intermittent cave sampling was initiated June 2011 and consistent bi-monthly sampling began in March 2012. W ater samples pumped from the soil lysimeters or collected from the cave were analyzed for stable isotopes of water (t 2 H/t 18 O, collected in 60-mL HDPE bottles), DIC (t 13 CDIC, ltered through Supor 0.45-m lters (Brahana 1997). e primary threat to water quality is nutrients and bacteria because of the karst topography and agriculture (Adamski et al. 1995), which is domi nated by poultry and cattle production (U.S. Department of Agriculture 2007). Best Management Practices for re ducing contaminant transport into the vulnerable karst waters have been employed, although the long-term ef fects of these practices are not fully understood (Davis et al. 2000). erefore, research that quanties contami nant transport along groundwater ow paths is vital in the Ozarks to better protect karst waters. Northwestern Arkansas has a temperate climate; average annual air temperature is 15.6C and precipita tion is approximately 109 cm per year (Adamski et al. 1995). During 2012, northwestern Arkansas experi enced a drought; especially during the warmer summer months (Simeral 2013). For example, 16.4 cm of rain fell during June, July, and August, compared to histori cal averages of 30 cm for the summer months (National Oceanic and Atmospheric Administration 2009). Re charge to the Springeld Plateau includes precipita tion and stream piracy (Brahana 1997). Groundwater levels generally reect surface topography (Adamski et al. 1995), but, as is characteristic in karst aquifers, groundand surface-water divides do not always coin cide (Brahana 1997). Karst features are more common in the pure carbonate lithologies, compared to units with higher proportions of chert and insoluble clays (Adamski et al. 1995; Brahana 1997). Caves are typi cally less than 150 m long and less than 30 m deep in the Ozarks of Arkansas because of the nearly horizontal bedding and the insoluble nature of the clay-rich rego lith mantle (Taylor et al. 2009). Blowing Spring Cave includes 2.4 km of mapped passage in the St. Joe Limestone Member of the Boone Formation (Fig. 1). e branching cave passage is domi nated by a cave stream passage, which originates at a sump (BS07). e recharge-area boundaries for the cave stream have not yet been identied. Using the normal ized baseow method (Brahana 1997) and a baseow at the spring of 0.009 m 3 /s, the recharge area is estimated to be between 2.9 to 6.1 km 2 Meteoric water is recharged through the chert regolith mantle and into the shallow Springeld Plateau Aquifer, which includes the Boone Formation (Adamski et al. 1995), and precipitation en ters Blowing Spring Cave by either percolation through the soil/regolith mantle above the cave or via the unde ned ow paths of the cave stream sump. Soils above the cave are predominantly extremely gravelly silt loam, 1.8 m thick, with a high capacity to transmit water, from approximately 5 to 15 cm/hr (Natural Resources Con servation Service 2012). Chert beds in the Boone For mation can be observed in the cave ceiling and cause lateral groundwater ow due to local perching. Discrete points of drip water enter the cave passage where these chert layers are breached by fractures; discharge at these drip-water points increases following storm events. Site BS06 is one example where drip water enters a domed, side passage and collects in a pool, before owing into the main cave stream. Site BS01 is located where the cave stream discharges at the surface as a spring and tributary to Little Sugar Creek, which denes local base level. K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 265 into 40-mL total organic carbon vials without headspace and preserved with 40 L of 3.6 M sodium azide to stop biologic activity), and major anions (collected in 125-mL HDPE bottles). e limited volumes of soil water (espe cially during dry surface conditions) required that water samples for isotopic analysis be immediately transferred to the smallest practical vial size to minimize headspace and decrease the potential for evaporation. Physical pa rameters (pH, specic conductance, and temperature) were monitored in the cave stream during sample col lection (small water volumes prevented measurement of physical parameters in soil water). Samples were kept on ice until transported to an environmental chamber (4C) at the University of Arkansas Stable Isotope Laboratory (UASIL) in Fayetteville, Arkansas. Precipitation samples were collected at the HOBO Micro Station for t 2 H/t 18 O analysis. A funnel directed precipitation through looped tubing and into 1-L HDPE sample bottle, which was also connected to an overow bottle lled with deionized water to prevent evapora tion. Samples were collected in 60-mL to 250-mL HDPE bottles (depending on precipitation volume, to minimize headspace). Precipitation samples were collected daily for most rain events (composite samples) or, during larger rain events, samples were collected periodically throughout the storm. Hydrogen and oxygen stable isotope ratios (t 2 H/ t 18 O) were measured using a high-temperature reduc tion unit interfaced to a Delta plus XP isotope ratio mass spectrometer (IRMS TCEA, ermo Scientic) at UA SIL. Samples were loaded into 1.5-mL auto-sampler vi als and 1 L of sample was injected into the TCEA. e furnace on the TCEA was operated at 1,425C with a glassy carbon reactor. A 5a-mol-sieve gas chromatogra phy column (GCC) separated the resulting H 2 and CO gases, which were admitted to the IRMS via a con Flo III interface (Gehre et al. 2004). Samples were normalized to the Vienna Standard Mean Ocean W ater (VSMO W ) scale following Nelson (2000) using three isotopically distinct standards analyzed multiple times throughout the run. e precision of the hydrogen measurement was .0 and the precision for oxygen was .2. DIC samples were analyzed for concentration and isotopic composition (t 13 CDIC) at the Colorado Pla teau Stable Isotope Laboratory in Flagsta, Arizona on a Total Organic Carbon Analyzer (Aurora OI 1010 Col lege Station, Texas) interfaced to an IRMS (Delta plus XL ermoQ uest Finnigan Bremen, Germany) following a procedure modied from St-Jean (2003). DIC was acidi ed with phosphoric acid to form carbon dioxide (CO 2 ), which was then carried via helium through a scrubber unit and GCC to remove nitrogen interferences and into the IRMS (Knierim 2009). Isotopic compositions were reported using t nota tion: Eqn. 1 where t represents the isotopic system and R is the ratio of the heavy to light isotope ( 13 C/ 12 C, 2 H/ 1 H, or 18 O/ 16 O) for the sample relative to a standard. e Vienna Peedee Belemnite was used as the standard for t 13 CDIC and VSMO W was used for t 2 H/t 18 O (Coplen 1996). Major anion geochemistry was analyzed at the Ar kansas W ater Resources Center (A WRC) W ater Q uality Laboratory in Fayetteville, Arkansas using ion chroma tography; a Dionex D X with an IonPac AS4ASC analytical column measured uoride, bromide, chloride, nitrate, and sulfate (A WRC 2008). Only the chloride (Cl) data will be discussed. D ATA ANALYSIS Following methods by Lee and Krothe (2001), a threecomponent hydrograph separation was completed to quantify precipitation (Q R ), soil water (Q S ), and bedrock matrix water (Q B ) contributions to the cave stream dur ing storm events (Q M ). Soil and bedrock-matrix waters together represent the pre-event component of storm ow (Q P ) and Q R represents storm-event water that has traveled rapidly along macropores and dissolution con duits. To separate Q P and Q R from Q M a two-component mixing model using t 2 H and t 18 O of water was solved rst (Lakey & Krothe 1996): Eqn. 2 where Q M is the total discharge of the cave stream at the spring (BS01). e isotopic composition of storm event water at BS01 (t M ), pre-event water from the soil and cave during baseow conditions (t P ), and precipitation (t R ) were determined using either hydrogen or oxygen, to check for consistency in the technique (Lakey & Krothe 1996). Cl was used as a secondary tracer and Equation 2 was completed using Cl concentration, not the stable isotope ratio of Cl. Once the proportions of Q P and Q R were deter mined, Q P was separated into Q S and Q B components using DIC concentration (C) and isotopic composition (t) in a three-component mixing model (Lee & Krothe 2001): Eqn. 3 U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 266 RESULTS METEORIC WATER LINE AND ISOTOPES OF WATER r t 2 H AND t 18 O f Median compositions for t 2 H and t 18 O of precipitation were .1 and .7, respectively (n = 41), for the time period between October 2011 and September 2012. A local meteoric water line (LMWL) for precipitation (not normalized for amount) was dened by rst-order regression as (Fig. 2): 2 H = 6.7 18 O + 1.5 Eqn. 5 Median soil water t 2 H and t 18 O compositions were .6 and .2, respectively (n = 11). Median cave water t 2 H and t 18 O compositions were .2 and .7, respectively, for all samples (baseow and storm ow, n = 83). During baseow conditions, median cave water t 2 H and t 18 O compositions were .2 and .9, respectively (n = 59). Collectively, the median t 2 H and t 18 O compositions for pre-event water (i.e., combing soil and cave baseow samples) were .4 and .9, respectively, (t P in Eqn. 2). D ISSOLVED INORGANIC CARBON rDICf Soil-water DIC had a median concentration of 1.8 mg/L (C S in Eqn. 3, n = 8) and median t 13 CDIC composition of .9 (t S in Eqn. 3, n = 8). Soil DIC concentration and isotopic composition were generally lower and lighter, respectively, than DIC in the cave (Fig. 3). Cave-water DIC had a median concentration of 24.6 mg/L (n = 56) and median t 13 CDIC composition of .5 (n = 56) for all samples (including baseow and storm ow). During base ow conditions, median DIC concentration and t 13 CDIC composition were 25.0 mg/L (n = 43) and .3 (n = 43), respectively, in the cave. Site BS06 is a drip-water pool (sep arate from the cave stream), so DIC concentrations and compositions were also calculated for the cave stream sites only (BS01, BS03, and BS07); DIC in the cave stream had a median concentration of 25.2 mg/L (n = 35) and median t 13 CDIC composition of 14.5 (n = 35). Additionally, site BS07 is closest to the sump (i.e., the source of water for the cave stream), and median DIC concentration for BS07 (during baseow) was 27.8 mg/L (n = 8) and median t 13 CDIC composition was .3 (n = 8). Q B = Q M Q R Q S Eqn. 4 where the isotopic composition of storm event water at BS01 (t M ), soil water (t S ), the cave stream at baseow (t B ), and precipitation (t R ) were determined using t 13 C DIC. For the full derivation of the previous equations, see Lakey and Krothe (1996) and Lee and Krothe (2001). Fig. 2: A Local Meteoric Water Line (LMWL) for northwest ern Arkansas was developed for comparison to the Global Mete oric Water Line (GMWL; Craig 1961). K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 267 HYDROGRAPH S EPARATIONS During a March 2012 storm event, 12.2 cm of precipi tation fell over 82 h and discharge increased from 0.01 to 0.59 m 3 /s in 59 h (Fig. 4). e isotopic composition of precipitation varied throughout the storm event be cause of rainout; for example, t 2 H decreased from .7 to .2 over 28 h and a similar pattern occurred for t 18 O. Although, at the spring, t 2 H and t 18 O only changed by approximately 3 and 1, respectively (Fig. 4). Precipitation from the storm event was normalized for amount because throughout the event the isotopic com position of precipitation became more depleted in 2 H and 18 O (Fig. 5). erefore, the values for t R in Equation 2 changed throughout the storm event (Fig. 5) and ranged between .2 and .4 for t 2 H and .7 and .6 for t 18 O (Tab. 1). Solving for Equation 2, the two-compo nent mixing model either overestimated (>100%) or un derestimated (<0%) the contribution from Q R (Tab. 1). e two-component mixing model was completed a second time using Cl concentration to provide an es timate of contribution from precipitation, as the mixing model using stable isotopes was inaccurate. Median Cl was 3.2 mg/L (n = 5) in the soil and 6.3 mg/L (n = 23) in the cave during baseow. Collectively, median Cl dur ing baseow for all sites (soil and cave) was 6.2 mg/L (t P in Eqn. 2). Precipitation samples were only analyzed for t 2 H and t 18 O, so historical, annual Cl concentration from the National Atmospheric Deposition Programs (NADP) National Trends Network were used for nearby sites in northern Arkansas (AR27), southeastern Kansas (KS07), southern Missouri (MO50), and eastern Okla homa (OK08 and OK99). Precipitation from the NADP samples had median Cl of 0.138 mg/L (t R in Eqn. 2) for the time period between 1980 and 2011, depending on the individual site (NADP 2007). Using Cl, Q R was found to vary with time and contribute between 6 and 43% of the total discharge throughout the March storm event, depending on the point along the hydrograph (Tab. 1). For example, Q R contributed the maximum amount to Q M on March 21, 2012 at 3:35 p.m. (Tab. 1), approxi mately 43 h aer precipitation began. During the storm event, DIC decreased to a mini mum of 8.6 mg/L from a baseow concentration of 29.9 mg/L and t 13 CDIC decreased to .2 from a base ow composition of .0 (Fig. 4). If Q R is ignored because the two-component mixing model using t 2 H/ t 18 O was inaccurate, the three-component mixing mod el (Eqns. 3 and 4) becomes a two-component mixing model, separating Q P into Q S and Q B using the following equations: Eqn. 6 Q M = Q P = Q B + Q S Eqn. 7 e precipitation terms (C R and t R ) are eliminat ed from Equation 3. Solving for Equations 6 and 7 us ing DIC concentrations and isotopic composition at all cave water sites during baseow (BS01, BS03, BS06, and BS07), Q S contributed 23 to 69% of the total discharge (Tab. 2, Fig. 6). If site BS06 is not included, Q S contrib Fig. 3: e cave stream tended to have higher DIC concentrations and heavier 13 CDIC composi tions than the soil. Storm samples at BS01 represent mixed ow ( M ) from a combination of soil ( S ) and bedrock-matrix ( B ) wa ters. U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 268 uted 24 to 69% of the total discharge (Tab. 2). If only site BS07 is used, then Q S contributed 32 to 72% of the total discharge (Tab. 2). If Q R was found to contribute be tween 6 and 43% of the total ow (based on the two-component mixing model using Cl), then the remaining portion of storm ow (i.e., the contribution from Q P ) can be separated into Q S and Q B to solve Equation 4 (using the percentages calculated previously from Eqns. 6 and 7, Tab. 2). Similar to the previous discus sion, dierent values for C B and t B were used, depending on the location in the cave (Fig. 3). Using DIC concentration and isotopic composition at all cave wa ter sites during baseow, Q S contributed 20 to 50% of the total discharge (Tab. 2, Fig. 6). If site BS06 is not included, Q S contributed 22 to 50% of the total dis charge (Tab. 2). If only site BS07 is used, then Q S contributed 28 to 53% of the to tal discharge (Tab. 2). Tab. 1: Results from the two-component hydrograph separation separating pre-event water (Q P ) and event water (Q R 2 H/ 18 O) or chloride (Cl). Date (GMT -6) BS01 ( M ) Discharge (Q M ) Precipitation ( R )* Precipitation (Q R ) 2 H 18 O Cl 2 H 18 O using 2 H using 2 H using 18 O using 18 O using Cl using Cl () () (mg/L) (m 3 /s) () () (m 3 /s) (%) (m 3 /s) (%) (m 3 /s) (%) 3/20/12 7:50 .5 .9 5.2 0.15 .2 .8 .12 .17 0.03 17 3/20/12 9:15 .8 .8 5.0 0.25 .2 .8 2.07 825 .03 0.05 19 3/20/12 11:20 .3 .4 4.8 0.42 .7 .7 .60 .29 0.10 23 3/20/12 12:55 .5 .5 4.6 0.44 .7 .7 .78 .26 0.11 25 3/20/12 14:35 .6 .6 4.3 0.42 .7 .7 .16 .18 0.13 31 3/21/12 8:15 .4 .1 3.8 0.42 .8 .0 0.13 30 .32 0.17 40 3/21/12 12:10 .3 .5 3.7 0.37 .8 .4 .06 .10 0.15 42 3/21/12 15:35 .3 .5 3.6 0.50 .8 .5 .07 .13 0.22 43 3/22/12 8:40 .7 .5 3.6 0.41 .3 .6 0.02 4 .10 0.17 42 3/23/12 8:00 .5 .8 3.7 0.19 .4 .6 0.00 1 .01 0.08 41 3/24/12 9:00 .6 .5 3.7 0.11 .4 .6 .01 .03 0.04 41 3/26/12 15:45 .0 .4 4.4 0.04 .4 .6 0.01 20 .01 0.01 29 3/28/12 15:55 .2 .8 4.7 0.02 .4 .6 0.00 10 0.00 0.01 25 4/6/12 8:30 .0 .6 5.8 0.02 .4 .6 0.00 21 0.00 0.00 6 Note: Percentages for discharge vary because of rounding. 2 18 O, and Cl was 0.138 mg/L (NADP 2007). K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 269 Tab. 2: Results from the three-component hydrograph separation (reduced to a two-component hydrograph separation) quantifying the contributions from soil water (Q S ) assuming either contributions from precipitation (Q R ) are zero or using Q R from Table 1. Date (GMT -6) BS01 ( M ) Discharge using Q R = 0 using Q R from Table 1 DIC 13 CDIC (Q M ) Q S Q S Q S ** Q S ** Q S *** Q S *** Q S Q S Q S ** Q S ** Q S *** Q S *** (ppm) () (m 3 /s) (m 3 /s) (%) (m 3 /s) (%) (m 3 /s) (%) (m 3 /s) (%) (m 3 /s) (%) (m 3 /s) (%) 3/20/12 7:50 13.6 .2 0.15 0.08 51 0.08 52 0.08 57 0.06 43 0.06 43 0.07 47 3/20/12 9:15 11.4 .9 0.25 0.15 62 0.16 62 0.17 66 0.13 50 0.13 50 0.13 53 3/20/12 11:20 10.3 .5 0.42 0.27 65 0.27 66 0.29 69 0.21 50 0.21 50 0.22 53 3/20/12 12:55 11.3 .0 0.44 0.26 58 0.26 59 0.28 63 0.19 44 0.20 44 0.21 47 3/20/12 14:35 10.8 .3 0.42 0.25 60 0.25 61 0.27 65 0.17 41 0.18 42 0.19 45 3/21/12 8:15 9.1 .0 0.42 0.28 66 0.28 67 0.30 70 0.17 40 0.17 40 0.18 42 3/21/12 12:10 9.3 .2 0.37 0.24 64 0.24 65 0.25 68 0.14 38 0.14 38 0.15 40 3/21/12 15:35 8.6 .9 0.50 0.35 69 0.35 69 0.36 72 0.20 39 0.20 40 0.21 41 3/22/12 8:40 8.6 .0 0.41 0.28 69 0.28 69 0.29 72 0.16 40 0.16 40 0.17 42 3/23/12 8:00 9.8 .2 0.19 0.12 62 0.12 62 0.12 66 0.07 36 0.07 37 0.07 39 3/24/12 9:00 10.8 .1 0.11 0.06 57 0.06 58 0.07 62 0.04 34 0.04 35 0.04 37 3/26/12 15:45 12.9 .7 0.04 0.02 48 0.02 49 0.02 54 0.01 34 0.01 35 0.02 39 3/28/12 15:55 16.5 .9 0.02 0.01 40 0.01 41 0.01 47 0.01 30 0.01 31 0.01 35 4/6/12 8:30 21.3 .3 0.02 0.01 23 0.01 24 0.01 32 0.00 20 0.00 22 0.01 28 Note: Percentages for discharge vary because of rounding. using DIC/t 13 CDIC values (C B /t B ) from all cave sites (BS01, BS03, BS06, and BS07) ** using DIC/t 13 CDIC values (C B /t B ) from cave stream sites only (BS01, BS03, and BS07) *** using DIC/t13CDIC values (C B /t B ) from site BS07 only U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 270 Fig. 4: During a storm event in March 2012, 12.2 cm of precipitation fell over 82 hrs. Baseow conditions (from March 16 th 2012) are shown with a dashed line for 2 H and 18 O. DIC was 29.9 mg/L and 13 CDIC was .0 on March 16 th 2012. K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 271 Fig. 6: H ydrograph separations for the March 2012 storm event quantifying the contributions from rain (Q R ), soil water (Q S ), and bedrock matrix wa ter from the cave (Q B ). Fig. 5: e isotopic composition of precipitation ( R ) was normalized for precipitation amount (arrow). e isotopic composition of pre-event water ( P ) includes soil water ( S ) and bedrock matrix water from the cave ( B ). U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 272 e LMWL for northwestern Arkansas (Fig. 2) had a lower slope and y-intercept than the global meteoric water line (Craig 1961) similar to meteoric water lines for the northern Great Plains (Harvey 2001; Harvey & W elker 2000; Simpkins 1995). Values for deuterium ex cess (d dened as t 2 H 8 t 18 O; Dansgaard 1964) are primarily controlled by the relative humidity at the time of vapor formation (Merlivat & Jouzel 1979) and would be between 3 and 15 for transported water vapor that has not undergone secondary processes (Harvey 2001). Smaller d values (d < 3) generally correspond to water that has experienced secondary evaporation and larger d values (d > 15) with water vapor that has second ary moisture added to it (Harvey 2001). For precipita tion from northwestern Arkansas, d varied between and 35, with a median value of 8. Most of the d values less than 3 occurred during the warmer, summer months and secondary evaporation may have occurred as the precipitation fell through the warm, dry atmos phere. e LMWL was developed during a drought and period of above-average temperatures for northwestern Arkansas (Simeral 2013), which may cause a lower slope and y-intercept than for a LMWL generated during times more representative of long-term local climatic condi tions. Moisture decits during the spring and summer months have been observed during historical, decadelong droughts in the mid-continent (Burnette & Stahle 2013) and future projections predict drier conditions for the southern Great Plains during the summer months (Patricola & Cook 2013). erefore, understanding karst hydrologic processes during periods of moisture decit will likely be important for future protection of karst wa ter resources. Both soil and cave water tended to cluster around t 18 O of (Fig. 2), representative of the regional av erage for precipitation (Clark & Fritz 1997). Generally, soil water is more enriched in 2 H and 18 O than precipi tation because of evaporative enrichment of heavier iso topes in the residual soil water (Gibson et al. 2008). In the soil above Blowing Springs Cave, soil water tended to be depleted in 2 H and 18 O compared to cave water (Fig. 2). Note that the two heavier soil-water composi tions (~ for t 18 O) were collected immediately fol lowing a precipitation event and reect the composition of meteoric water that has mixed with soil water. ere fore, water entering the soil zone, which is ultimately me teoric water, must have some nite residence time in the soil zone to be considered pre-event water. Residence times (or, in comparison, transit times) for soil water vary based on the type of soil, intensity of precipitation, and topography (Tetzla et al. 2011). e soils above Blowing Springs Cave have a high capacity to transmit water (Natural Resources Conservation Service 2012), which would imply a short residence time and fast tran sit time in or through the soil zone, although exact times have not yet been constrained for this study site. e isotopic dierence between median soil water and median cave water (during baseow) was approxi mately 4.4 and 0.3 for t 2 H and t 18 O, respectively, and cave water exerts a stronger control on the isotopic composition of t P (Fig. 5), likely because of sampling bias as more cave water samples were collected (n = 59 for cave at baseow) compared to soil water (n = 11). e isotopic variation between precipitation from the March 2012 storm event and the pre-event water was ap proximately 8.0 and 1.7 for t 2 H and t 18 O, respec tively, at the end of the storm event (Fig. 5). e isotopic composition of precipitation became lighter throughout the storm event because of rainout (Fig. 5); the process of the vapor mass becoming progressively lighter dur ing the storm event as the heavier isotopes ( 2 H/ 18 O) are partitioned into the liquid phase with a progressive, con comitant eect on continued precipitation (Clark & Fritz 1997). Most of the precipitation fell overnight on March 19 th (Fig. 4) and this composite sample (.7 and .9 for t 2 H and t 18 O, respectively) was isotopically similar to the pre-event water (Fig. 5). Additionally, the storm-ow samples at BS01 (t M ) were enriched in 18 O compared to pre-event water (t P ). ese two factors con tributed to the two-component hydrograph separation using t 2 H/t 18 O either overor under-estimating the con tribution from precipitation to the spring and disagree ment between solving Equation 2 for t 2 H versus t 18 O (Tab. 1). Other studies have found that the quick-ow component (or precipitation) to karst springs accounted for 10 to 40% of total discharge (Lakey & Krothe 1996; Lee & Krothe 2001; Long 2009; Mahler & Garner 2009; Trek et al. 2006). If the contributions from Q R were ignored, which is unrealistic in karst systems where rapid ow through soil macropores (Iqbal & Krothe 1995) and conduits occur (Lee & Krothe 2003), the hydrograph separation using DIC (Eqns. 3 and 4) was completed as a two-component mixing model (Eqns. 6 and 7) to separate Q P into Q S and Q B which provides an initial estimate for contributions from water stored in the recharge zone. Q uantifying the proportions of pre-event water is useful in karst settings because changing contributions from the unsaturated versus saturated zones can control spring geochemistry during baseow and storm events (Peterson et al. 2002). Storm-ow samples from BS01 (C M /t M ) plotted between soil water (C S /t S ) and cave water (C B /t B ), meaning that D ISCUSSION K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 273 C ONCLUSIONS e two-component hydrograph separation using t 2 H/ t 18 O of water to separate pre-event water (Q P soil and cave-stream water at baseow) and event water (Q R precipitation) either overestimated (>100%) or under estimated (<0%) the contribution from Q R because (1) precipitation (when normalized for amount) was isotopi cally similar to the pre-event water and (2) storm ow at BS01 (t M ) was outside the range of the two end-members (Fig. 5). erefore, an additional source of water enriched in 18 O may need to be considered. If Q R was ignored, Q S was found to contribute 23 to 72% of the total discharge, depending on the values used for C B /t B (Tab. 2). If Q R was calculated using Cl, Q S was found to contribute 20 to 53% of the total discharge, again depending on the values used for C B /t B (Tab. 2). Is the variation in contributions from Q S or Q B important (Fig. 6), when using dierent DIC concentrations and isotopic compositions (C B /t B ) from along the cave stream reach? At each time-step (or sampling point along the hydrograph), the dierence in contributions from Q S or Q B between (1) all cave water at baseow (BS01, BS03, BS06, and BS07) and (2) only cave water at BS07 (closest to the sump source of the cave stream) averaged 5% when Q R was assumed to be zero or 4% when Q R was calculated using Cl (Tab. 2). is ques tion of importance will be addressed in future research by solving the three-component mixing model (Eqn. 3) to assess the sensitivity of model to variations in DIC. At this point, the variation in contributions from Q S or Q B using dierent C B /t B values may or may not be im portant because of solving Equations 6 and 7 (where Q R Q S and Q B are not calculated independently using DIC), and possibly over-estimating the contributions from Q S compared to Q B Of the stored-water component (Q P ), soil water ex erted important control on storm-event geochemistry at Blowing Spring (Figs. 3 and 6). e importance of va dose-zone water (either soil or epikarst) has been dem onstrated in other karst systems (Doctor et al. 2006; Lee & Krothe 2001; Lee & Krothe 2003; Trek et al. 2006). Vadose-zone water is inherently dicult to sample (small volumes, heterogenous, etc.), but this research highlights the importance of considering soil water in mantled karst settings. is is especially important in northwestern Arkansas where the soil zone is impacted by anthropogenic changes in land use, including ur banization and agriculture. Further research at Blowing the two-component mixing model should reect real istic mixing conditions during the storm event (Fig. 3). Using dierent values for C B/ t B depending on the loca tion in the cave stream (BS01, BS03, or BS07) or drip pool (BS06), changed the contribution from Q S by 4 to 9% throughout the storm event (Tab. 2). Cave-stream water at BS07 during baseow had similar t 13 CDIC compositions to all cave stream samples (~ .3), but higher DIC concentrations (Fig. 3). erefore, solving Equations 6 and 7 using C B from BS07 caused the lowest estimations for the contributions from Q B or increased the contribution from Q S (Fig. 6; Tab. 2). e two-component mixing model using Cl esti mated that Q R contributed up to 43% of the total storm ow, similar to other karst spring hydrograph separations (Mahler & Garner 2009), which decreased the estimated contribution from Q S and Q B For example, if C B /t B val ues for only site BS07 at baseow were used, then Q B contributed between 16 and 61% of the total storm ow (down from 28 to 68% when Q R was assumed to be zero). Similar to the previous discussion, using DIC concentra tion and t 13 CDIC composition from only site BS07 to complete Equations 6 and 7 caused the greatest estima tions for contributions from Q S (Fig. 6; Tab. 2). Noting that Equation 3 was not completed for this discussion, as the two-component hydrograph separation using t 2 H versus t 18 O was inaccurate, the two-component hydro graph separation using Cl was completed to provide constraints on the pre-event contribution to storm ow. erefore, including Q R as a third component in Equa tion 3 would provide a third source of water (i.e., precip itation) to mixed ow that has lower DIC concentrations and t 13 CDIC compositions enriched in 13 C compared to the other sources, based on equilibrium exchange and fractionation between atmospheric CO 2 and precipita tion (Clark & Fritz 1997; Lee & Krothe 2001). For exam ple, Lee and Krothe (2001) found that DIC concentration in precipitation was 2 mg/L and t 13 CDIC composition was calculated to be erefore, completion of the mixing model using Equations 6 and 7 likely over estimated the contribution from Q S because C S and C R would have similar DIC concentrations compared to C B Future expansion of the mixing model, including com pleting Equation 3 using temperatureand pH-depen dent values for C R and t R at Blowing Spring, may nd that the bedrock-matrix contributes greater volumes of water during storm events. U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 274 REFERENCES Adamski, J., Petersen, J., Freiwald, D. & J. Davis, 1995: Environmental and hydrologic setting of the Ozark P lateaus study unit, Arkansas, Kansas, Missouri, and Oklahoma.U.S. Geological Survey, Report number: W ater Resources Investigation Report 94. Arkansas W ater Resources Center, 2008: A WRC W a ter Q uality Laboratory.[Online] Available from: http://www.uark.edu/depts/awrc/waterqualitylab. html [Accessed 6th February 2013]. Brahana, J., 1997: Rationale and methodology for ap proximating spring-basin boundaries in the man tled karst terrane of the Springeld Plateau, north western Arkansas.In: Beck & Stephenson (eds.) e Engineering Geology and H ydrogeology of Karst Terranes, Balkema, pp. 77, Rotterdam. Burnette, D. & D. Stahle, 2013: Historical perspective on the dust bowl drought in the central United States.Climatic Change, 116, 479. Buttle, J., 1994: Isotope hydrograph separations and rapid delivery of pre-event water from drainage basins.Progress in Physical Geography, 18, 1, 16. Clark, I. & P. Fritz, 1997: Environmental Isotopes in H y drogeology. Lewis Publishers, pp. 328, Boca Raton. Coplen, T., 1996: New guidelines for reporting stable hydrogen, carbon, and oxygen isotope-ratio data.Geochimica et Cosmochimica Acta, 60, 17, 3359 3360. Craig, H., 1961: Variations in meteoric waters.Science, 133, 1702. Daansgard W ., 1964: Stable isotopes in precipitation.Tellus, 16, 4, 436. Spring will include analysis of major cation/anion geo chemistry in the soil zone and comparison to geochem istry at site BS06, which receives drip water from the soil and epikarst zones. Additionally, the residence and tran sit times of soil water will be better constrained, to help address questions of mixing water sources along ground water ow paths. e relation between the isotopic composition of the gaseous cave atmosphere (CO 2 ) and the aqueous cave stream (DIC) is being investigated at Blowing Spring to further understand the variation in DIC concentration and isotopic composition along the cave stream, and how diering values for C B /t B change the nal results of the two-component and three-component mixing mod els. Research has shown that the concentration and iso topic composition of cave-CO 2 varies over time (Kow alczk & Froelich 2010; Pollock et al. 2011; Sptl et al. 2005) so the potential eect of the cave atmosphere on the cave stream will be considered when analyzing these antecedent conditions prior to storm events. Although the two-component hydrograph separation could not be completed using stable isotopes of water (t 2 H and t 18 O), the variation in DIC concentration and isotopic compo sition between the soil and cave highlights the useful ness of carbon to characterize the geochemistry of karst systems. Future work will expand the mixing models discussed, compare multiple storm events to assess vari ability in spring response to storm events in the Ozarks, and relate the source-water contributions to changes in water quality. A CKNO WLEDGMENTS is manuscript was greatly improved by the comments from Natasa Ravbar, Tim Kresse, James Petersen, Keith Lucey, and an anonymous reviewer. Funding for this re search was provided by the Geological Society of America and the National Speleological Society. is manuscript was developed under STAR Fellowship Grant Number FP917347 awarded by the U.S. Environmental Protection Agency (EPA). It has not been formally reviewed by the EPA. e views expressed in this manuscript are solely those of Katherine J. Knierim, and EPA does not endorse any products or commercial services mentioned in this manuscript. Any use of trade, rm, or product names is for descriptive purposes only and does not imply en dorsement by the U.S. Government. e Arkansas W ater Resources Center W ater Q uality Laboratory provided analytical support. Special thanks are given to the Bella Vista Property Owners Association for use of the re search site and non-monetary support and to the Boston Mountain Grotto for mapping Blowing Springs Cave. K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS

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ACTA CARSOLOGICA 42/2-3 2013 275 Davis, R., Brahana, J., & J. Johnston, 2000: Ground wa ter in northwest Arkansas: Minimizing nutrient con tamination from non-point sources in karst terrane.Arkansas W ater Resources Center, Report Number: MSC. Doctor, D., Alexander, E., Petric, M., Kogovsek, J., Ur banc, J., Lojen, S. & W Stichler, 2006: Q uantica tion of karst aquifer discharge components during storm events through end-member mixing analysis using natural chemistry and stable isotopes as trac ers.Hydrogeology Journal, 14, 1171. Gehre, M., Geilmann, H., Richter, J., W erner, R., & W Brand, 2004: Continuous ow 2 H/ 1 H and 18 O/ 16 O analysis of water samples with dual inlet precision.Rapid Communications in Mass Spectrometry, 18, 2650. Gibson. J., Briks, S. & T. Edwards, 2008: Global predic tion of t A and t 2 Ht 18 O evaporation slopes for lakes and soil water for seasonality.Global Biogeo chemical Cycles, 22, 1. Harvey, F., 2001: Use of NADP archive samples to deter mine isotope composition of precipitation: char acterizing the meteoric input function for use in ground water studies.Ground W ater, 39, 3, 380 390. Harvey, F. & J. W elker, 2000: Stable isotopic composition of precipitation in the semi-arid north-central por tion of the US Great Plains.Journal of Hydrology, 238, 90. Iqbal M. & N. Krothe, 1995: Inltration mechanisms related to agricultural waste transport through the soil mantle to karst aquifers of southern Indiana, USA.Journal of Hydrology, 164, 171. Jones, J., Sudicky, E., Brookeld, A. & Y. Park, 2006: An assessment of the tracer-based approach to quan tifying groundwater contributions to streamow.W ater Resources Research, 42, 1. Kendall, C., McDonnell, J. & W Gu, 2001: A look inside black box hydrograph separation models: a study at the Hydrohill catchment.Hydrological Process es, 15, 1877. Knierim, K., 2009: Seasonal variation of carbon and nu trient transfer in a northwestern Arkansas cave. MS thesis. e University of Arkansas, pp. 141. Kowalczk, A. & P. Froelich, 2010: Cave air ventilation and CO2 outgassing by radon-222 modeling: how fast do caves breathe?.Earth and Planetary Science Letters, 289, 209. Laincz, J., 2011: Investigation of Nitrate P rocessing in the Interow Z one of Mantled Karst, Northwestern Ar kansas.U. S. Geological Survey, Report Number: Scientic Investigations Report 2011, 75. Lakey, B. & N. Krothe, 1996: Stable isotopic variation of storm discharge from a perennial karst spring, Indi ana.W ater Resources Research, 32, 721. Lee, E. & N. Krothe, 2001: A four-component mixing model for water in a karst terrain in south-central Indiana, USA: using solute concentration and stable isotopes as tracers.Chemical Geology, 179, 129 143. Lee, E. & N. Krothe, 2003: Delineating the karstic ow system in the upper Lost River drainage basin, south central Indiana: using sulphate and t 34 S SO4 as tracers.Applied Geochemistry, 19, 145. Long., A., 2009: Hydrograph separation for karst water sheds using a two-domain rainfall-discharge mod el.Journal of Hydrology, 364, 249. Luhman, A., Covington, M., Alexander, S., Chai, S., Schwartz, B., Groten, J. & E. Alexander, 2012: Com paring conservative and nonconservative tracers in karst and using them to estimate ow path geom etry.Journal of Hydrology, 448, 201. Mahler, B. & B. Garner, 2009: Using nitrate to quantify quick ow in a karst aquifer.Ground W ater, 47, 3, 350. Merlivat, L. & J. Jouzel, 1979: Global climatic interpre tation of the deuterium-oxygen 18 relationship for precipitation.Journal of Geophysical Research, 84, C8, 5029. National Atmospheric Deposition Program, 2007: Na tional Trends Network Data, Annual PrecipitationW eighted Means.[Online] Available from: http:// nadp.sws.uiuc.edu/NTN/ntnData.aspx [Accessed 4 th February 2013]. National Oceanic and Atmospheric Administration, 2009: Arkansas Northwest Division 01, 1895.[Online] Available from: http://www7.ncdc.noaa. gov/CDO/CDODivisionalSelect.jsp [Accessed 3 rd March 2010]. Natural Resources Conservation Service, United States Department of Agriculture, 2012: W eb Soil Survey: Custom Soil Resource Report for Benton County, Arkansas.[Online] Available from: http://web soilsurvey.nrcs.usda.gov/. [Accessed 20 th January 2013]. Nelson, S., 2000: A simple, practical methodology for routine VSMO W/SLAP normalization of water samples analyzed by continuous ow methods.Rapid Communications in Mass Spectrometry, 14, 12, 1044. Ogunkoya, O. & A. Jenkins, 1993: Analysis of storm hy drograph and ow pathways using a three-compo nent hydrograph separation model.Journal of Hy drology, 142, 71. U SING ISOTOPES OF DISSOLVED INORGANIC CARBON SPECIES AND W ATER TO SEPARATE SOURCES OF RECHARGE IN ...

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ACTA CARSOLOGICA 42/2-3 2013 276 Panno, S., Hackley, K., Hwang, H., & W Kelly, 2001: De termination of the sources of nitrate contamination in karst springs using isotopic and chemical indica tors.Chemical Geology, 179, 113. Patricola, C. & K. Cook, 2013: Mid-twenty-rst century warm season climate change in the Central United States. Part I: regional and global model predic tions.Climate Dynamics, 40, 551. Peterson, E., Davis, R., Brahana, J. & H. Orndor, 2002: Movement of nitrate through regolith covered karst terrane, Northwest Arkansas.Journal of Hydrol ogy, 256, 35. Pollock, E., Knierim, K. & P. Hays, 2011: Seasonal carbon dynamics in a northwestern Arkansas cave: linking climate and cave conditions.U. S. Geological Sur vey, Report Number: Scientic Investigations Re port 2011, 37. Raeisi, E., Groves, C. & J. Meiman, 2007: Eects of partial and full pipe ow on hydrochemographs of Logs don river, Mammoth Cave Kentucky USA.Journal of Hydrology, 337, 1. Rantz, S., 1982: Measurement and Computation of Streamow: V olume 1. Measurement of Stage and Discharge.U.S. Geological Survey, Report Number: Geological Survey W ater-Supply Paper 2175. Rice, K. & G. Hornberger, 1998: Comparison of hydro chemical tracers to estimate source contributions to peak ow in a small, forested, headwater catch ment.W ater Resources Research, 34, 7, 1755 1766. Simeral, D., 2013: U.S. Drought Monitor: January 15, 2013.[Online] Available from: http://www. droughtmonitor.unl.edu/archive.html [Accessed 23 rd January 2013]. Simpkins, W ., 1995: Isotopic composition of precipita tion in central Iowa.Journal of Hydrology, 172, 185. Sklash, M., & R. Farvolden, 1979: e role of ground water in storm runo.Journal of Hydrology, 43, 45. Sklash, M., Farvolden, R., and P. Fritz, 1976: A concep tual model of water response to rainfall, developed through the use of oxygen-18 as a natural tracer.Canadian Journal of Earth Sciences, 13, 271. Sptl, C., Fairchild, I. & A. Tooth, 2005: Cave air control on dripwater geochemistry, Obir Caves (Austria): Implications for speleothem deposition in dynami cally ventilated caves.Geochimica et Cosmochim ica Acta, 69, 2451. St-Jean, G., 2003: Automated quantitative and isotopic ( 13 C) analysis of dissolved inorganic carbon and dissolved organic carbon in continuous-ow using a total organic carbon analyser.Rapid Communi cations in Mass Spectrometry, 17, 419. Taylor, D. S., Goodwin, D. P., Bitting, C. J., Handford, R., & M., Slay, 2009: e Ozark Plateaus: Arkansas.In: Palmer, A. N. & M. V. Palmer (eds.) Caves and Karst of the USA National Speleological Society, Inc., pp. 172, Huntsville, Alabama. Tetzla, D., Soulsby, C., Hrachowitz, M., & M. Speed, 2011: Relative inuence of upland and lowland headwaters on the isotope hydrology and transit times of larger catchments.Journal of Hydrology, 400, 438. Trek, B., Veselic, M., and J. Pezdic, 2006: e vulner ability of karst springs: a case study of the Hubelj Spring (SW Slovenia).Environmental Geology, 49, 865. U.S. Department of Agriculture, 2007: State Fact Sheets: Arkansas.[Online] Available from: http://www. ers.usda.gov/data-products/state-fact-sheets/statedata.aspx?StateFIPS=05&StateName=Arkansas#P6 6692c6d09db47faa4f648925911c323_2_657iT24R0 x0 [Accessed 23 rd January 2013. K ATHERINE J. KNIERIM, ERIK POLLOCK & P HILLIP D. H AYS



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I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE CUMBERLAND P LATEAU OF KENTUCKY USA O GLJIKOVI IZOTOPI V KRA KEM VODONOSNIKU CUMBERLANDSKE PLANOTE KENTUCKY ZDA Lee J. FLOREA 1 Izvleek UDK UDK: 546.26.027*14:551.444(736.9) Lee J. Florea: Ogljikovi izotopi v krakem vodonosniku CumOgljikovi izotopi v krakem vodonosniku CumCum berlandske planote Kentucky, ZDA V kraki podtalnici zaledja potoka Otter Creek (Cumberlandska planota, Kentucky, ZDA), smo merili koncentracijo in izotopsko sestavo raztopljenega organskega (DOC) in anorganskega (DIC) ogljika. Podatke smo primerjali z ravnoteno kemijo karbonatov in sadre. Ugotovili smo, da koncentraci ja DOC pada s pretokom vode, da lahko izotopsko obogati tev DIC vzdol ve reakcijskih poti, neposredno poveemo z nasienostjo raztopine in da ima pri raztapljanju pomembno vlogo oksidacija reduciranega vepla. DOC izhaja iz vegetacije tipa C3, s povpreno t 13 C DOC DIC izhaja iz pedogene ga CO 2 in iz raztopljenih karbonatov v obliki HCO 3 Dotoki v vodonosnik so nenasieni na kalcit in imajo izrazito nizk koncentracijo DIC, manj kot 1 mmol/L. Na izvirih so vode v ravnoteju oz. prenasiene na kalcit, koncentracija DIC pa do see 3 mmol/L. Na mestih, kjer iz vode izhaja CO 2 in se izloa kalcit, so izotopske vrednosti DIC med ,3 in ,4 kar zamegljuje enostavna razmerja med t 13 C DIC pretokom in stopnjo nasienja. V veplenih vodah so koncentracije DIC med 2 mmol/L in odstopajo od vrednosti, ki bi jih priakovali pri raztapljanju karbonatov z ogljikovo kislino, v smeri vrednosti v matini karbonatni kamnini. Oitno se del DIC iz karbonatov sproa pri reakciji z vepleno kislino, ki nastaja ob oksidaciji reduciranega vepla iz plitvih slanic nanih polj. Kljune besede: raztopjen organski in anorganski ogljik, re dukcija vepla, geokemija, indeks nasienja. 1 Department of Geological Sciences, Ball State University, e-mail: lorea@bsu.edu Received/Prejeto: 25.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 277, POSTOJNA 2013 Abstract UDC UDK: 546.26.027*14:551.444(736.9) Lee J. Florea: Isotopes of Carbon in a Karst Aquifer of the Cumberland Plateau of Kentucky, USA In this study, the concentration and isotopic composition of dissolved organic carbon (DOC) and dissolved inorganic car bon (DIC) are measured in the karst groundwater of the Ot ter Creek watershed of the Cumberland Plateau of Kentucky, USA. Comparisons among these data and with the geochemis try of carbonate and gypsum equilibrium reactions reveal that DOC concentration is inversely related to discharge, multiple reaction pathways provide DIC with isotopic enrichment that may be directly related to mineral saturation, and oxidation of reduced sulfur is possible for dissolution. DOC is derived from C3 vegetation with an average t 13 C DOC of DIC in groundwater is derived from both pedogenic CO 2 and HCO 3 from dissolved carbonate. At input sites to the karst aquifers DIC concentrations are expectedly low, less than 1 mmol/L, in waters that are undersaturated with respect to calcite. At the output of these karst aquifers DIC concentrations reach 3 mmol/L in waters that are at or above calcite saturation. Values of t 13 C DIC range between .3 and .4 with CO 2 degassing and calcite precipitation at some sites obfuscating a simple re lationship between t 13 C DIC discharge, and mineral saturation. In addition, concentrations of DIC in sulfur seeps within the watershed range between 2 mmol/L with t 13 C DIC values in some samples skewed more toward the anticipated value of carbonate bedrock than would be expected from reactions with carbonic acid alone. is suggests that the oxidation of re duced sulfur from shallow oileld brines liberates bedrock DIC through reactions with sulfuric acid. Keywords: dissolved organic carbon, dissolved inorganic car bon, sulfur redox, ion geochemistry, saturation index.

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ACTA CARSOLOGICA 42/2-3 2013 278 Carbonate aquifers, hosting 60% of the worlds proven petroleum reservoirs, 40% of known gas reserves, and in uencing approximately 25% of the worlds drinking wa ter (Ford and W illiams 2007), are particularly vital to our understanding of the global carbon cycle. A signicant fraction of carbon transport through karst is in dissolved form. Dissolved organic carbon (DOC) comprises or ganic acids and carbohydrates leached out of vegetation and soil. e composition of that DOC depends largely upon the overlying climate and land use. Dissolved in organic carbon (DIC) includes products of mineral re actions with acidity in water. Of primary importance is the carbonate equilibrium reaction with calcite and dolo mite, or H 2 CO 3 (aq) + [(1 x ) Ca, xMg ] CO 3 (sol) (1 x ) Ca (aq) + xMg (aq) + 2 HCO 3 (aq) (1) where the subscripts x and 1-x are proportional to the magnitude of calcite and dolomite within a karst aqui fer, respectively. In Equation 1 CO 2 sequestered into the aqueous system from the atmosphere, organic oxidation, or microbial respiration in the soil reacts with carbon ate bedrock to release carbonate ions into solution. In the pH ranges typical of water in karst aquifers, bicarbonate (HCO 3 ) is the principal dissolved ion. More recently, some emphasis has shied toward carbonate aquifers with secondary porosity that has partly evolved via the oxidation of reduced sulfur, via the following reaction: H 2 SO 4 (aq) + 2[(1 x ) Ca, xMg ]C O 3 (sol) 2(1 x ) Ca (aq) + 2 xMg (aq) + SO 4 (aq) + 2 HCO 3 (aq) (2) In such systems, such as the classic example of the karst of the Guadalupe Mountains of New Mexico (Hill 1990), CO 2 is not sequestered from the atmosphere and DIC is liberated from the carbonate bedrock alone. STABLE ISOTOPES OF CARBON In nature, carbon occurs as two stable isotopes, 12 C and 13 C with the abundance of the heavier isotope of approxi mately 1.1%. Mass-spectrometry can distinguish between these isotopes and compute the enrichment or depletion of the heavier isotope of carbon in a sample as compared the Vienna Pee Dee Belemnite (VPDB), denoted by t 13 C and calculated using (3) Values of t 13 C are reported in parts per thousand or per mille ( VPDB). Various processes in nature may fractionate the heavier or lighter stable isotope of carbon (Kendall & Caldwell 1998) In vegetation, the photosynthetic process pref erentially uptakes the lighter isotope during carbon fixation (Schlesinger 1997; Ehleringer & Cerling 2000). The resulting t 13 C of organic carbon in vegeta tion is depleted in the heavier isotope. The magnitude of this fractionation depends upon the nature of the photosynthetic pathway. In humid landscapes, most native plants utilize the C3 pathway, which yields or ganic matter with t 13 C values between -23 and -27. C4-type vegetation, in contrast, is more adapted to arid conditions, fixes less CO 2 during photosynthesis than the C3 pathway, and consequently has t 13 C val ues between -10 and -14. In shallow ground water, values of t 13 C in dissolved organic and inorganic carbon (t 13 C DOC and t 13 C DIC ) re ect a combination of microbiologic reactions, limestone dissolution, and gas-water exchange processes. W hen carbon dioxide CO 2 produced in soil by oxidation or by biogenic reactions is dissolved into water, the process preferentially selects the heavier isotope. e t 13 C DIC resulting enrichment is from combined microbial, diu sion, and equilibrium fractionation and may be as much as +6.4 (Clark & Fritz 1997). is pedogenic-derived DIC reacts with carbonate bedrock via Equation 1 to produce a DIC in groundwater from each source, with the stoichiometry in a simple reaction mandating a 50% blend of carbon in the products at mineral saturation. Undersaturated solutions would theoretically have val ues of t 13 C DIC that more reect soil CO 2 Oversaturated solutions may be more enriched in bedrock-derived DIC compared to a sample at saturation. e exact nature of t 13 C DIC in karst groundwater can be signicantly more complex than simple mixing and governed by phases of CO 2 enrichment or degassing along the owpath (Marlier & OLeary 1984), by alter native chemical weathering phenomena like those that release DIC from carbonate bedrock into an aqueous so lution without corresponding soil CO 2 (e.g., Equation 2 ), or by weathering of certain silicates that sequesters CO 2 from the atmosphere without the addition of bedrock DIC (e.g., weathering of wollastinite, CaSiO 3 Berner et al. 1983). In summary, changes in end-member con tributions to DOC and DIC (e.g., vegetative cover and bedrock composition) as well as variations in the hy draulic function of the underlying aquifer system from droughts or storm events may manifest as changes to the values of t 13 C DIC I NTRODUCTION LEE J. FLOREA

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ACTA CARSOLOGICA 42/2-3 2013 279 HYDROGEOLOGIC S ETTING Data in this study come from water samples collected within the Otter Creek watershed along the western es carpment of the Cumberland Plateau in W ayne County of southeastern Kentucky, USA (Fig. 1). Otter Creek is a third-order tributary of the Cumberland River. Staged incision of the Cumberland River associated with inter glacial phases of Plio-Pleistocene glaciation has been the primary control on the geomorphic evolution of the re gion (Anthony & Granger 2004) and has resulted in tiered cave systems (Simpson & Florea 2009). ese caves have developed within middle-Mississippian carbonates of the Slade Formation, including, from oldest to youngest, the St. Louis, Ste. Genevieve, Kidder, and Bangor Limestone members (Ettensohn et al. 1984). ese are underlain is paper considers the spatial and temporal variation of dissolved organic and inorganic carbon from one portion of the karst within the western mar gin of the Cumberland Plateau in southeast Kentucky as a component of a larger-scale investigation of carbon ux from karst in the Appalachian lowland plateaus. To that end, the geochemistry of water samples from sites considered by Dugan et al. (2012) and Florea (2013a) are used to compute calcite and gypsum saturation in dices, calcite saturation ratios, as well as the concentra tion of the species of inorganic carbon. Complementa ry to these data are measurements of dissolved organic carbon (DOC) and the stable isotopes of dissolved carbon in both organic and inorganic form (t 13 C DOC and t 13 C DIC ) in these same samples. Variations in these data are investigated as they pertain to sampling site and aquifer characteristics, as well as seasonal patterns manifest, in part, as variations in discharge at selected sites. Sunnybrook Anticline Monticello Otter CreekBeaver Creek AlphaSR 90 Scale: 5 miles NT Karst Springs Tufa Springs Sulfur Seeps Ground Water Flow Surface Water Flow Known Oil Reservoir SC SS HB SSS SCS DH1 DH2 BCA BH BC1 CT LH SB OC3 DC TC Fig. 1: Study area in Wayne County, KY USA. Index map of Kentucky illustrates the location of Wayne County. e light gray is the Otter Creek watershed. e Redmond Creek karst aqui fer outlined in black includes the sampling sites for Sandy Springs and Stream Cave and are indi cated by SS and SC, respectively. Other sampling sites, including karst springs, tufa springs, and sulfur seeps are labled with codes decribed in the text. e axis of the Sunnybrook Anticline is in dicated as a black line. Ground water ow paths are inferred by dye tracing and known cave survey. Oil reservoirs in shallow Mississippian-age strata identi ed by Abbott (1921) are denoted by gray ovals. I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY

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ACTA CARSOLOGICA 42/2-3 2013 280 by the early-Mississippian calcareous shale of the Salem-W arsaw and Fort Payne Formations and over lain by the late-Mississippian calcareous shale of the Paragon Formation. Regionally, early Pennsylvanian sandstones and conglomerates of the Lee Formation comprise the plateau surface. Relief in the study area exceeds 230 m with ridge tops above 530 m and val ley oors below 300 m. e active ow system in the karst aquifers of the Otter Creek watershed comprise dendritic net works of tributary conduits (Palmer 1991) that stair step through the stratigraphy (Crawford 1984) and emerge briey at the land surface where the owpath at the base of the Bangor Limestone traverses the Hartselle Formation before sinking again into the Kidder Limestone (W alden et al. 2007). ese tribu taries coalesce into sinuous, master conduits that may parallel hillside contours (Sasowsky 1994) and discharge at gravity-ow springs at the base of the plateau (Crawford 1984). e karst of the Cumber land Plateau is largely epigene; in other words, the source of acidity is derived from meteoric recharge and driven largely by the carbonate equilibrium reac tion presented in Equation 1 Florea (2013b) consid ers the timing and mode of this recharge using stable isotopes of oxygen and hydrogen (t 18 O and t 2 H) and reveals a strong seasonal bias toward winter months when evapotranspiration is reduced. Fig. 2: Time series plots of selected eld, ion, and carbon data from Stream Cave (SC open squares and solid line) and Sandy Springs (SS open diamonds and dashed line). All y-axes share the dates at the bottom of the gure. LEE J. FLOREA

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ACTA CARSOLOGICA 42/2-3 2013 281 METHODS SAMPLES AND ANALYTICAL METHODS Samples comprising this study include 16 bi-monthly samples collected in 2010 from two sites in the Redmond Creek karst aquifer, Stream Cave (SC) and Sandy Springs (SS) considered by Florea (2013a and b). Measurements of instantaneous discharge (Q) computed using an acoustic ow meter and standard USGS gauging techniques (e.g., Rantz 1982) accompany these samples. At Stream Cave, water emerges from a hillside spring in the Bangor Limestone, traverses the Hartselle Forma tion, and sinks within 100 m into the underlying Kid der Limestone. is site represents one of many inputs to the Redmond Creek karst aquifer. Sandy Springs, in contrast, is the principal outlet for the Redmond Creek karst aquifer. Complementing these data are samples col lected in 2011 from three sulfur seeps: Slickford Bridge (SB), Bertram Hollow (BH), and Beaver Creek #1 (BC1). In January of 2012, additional samples from 13 springs in the Otter Creek watershed complete the dataset. ese samples come from SS, SB, Otter Creek #3 (OC3), Don Carter Seep (DC), Blowing Cave (BCA), Tonyas Cave (TC), Coal Trace (CT), Herlan Buck Springs (HB), Triple S Cave (SSS), Spelunger Cave Springs (SCS), Dry Hol low #1 (DH1), Dry Hollow #2 (DH2), and Light House Spring (LH). ree of these sites (SB, OC3, and DC) represent sulfur seeps. Nine additional sites (SS, BCA, TC, CT, HB, SSS, SCS, DH1, DH2) are classic, gravityow karst springs. e remaining site (LH) is a large tufa spring. W ith the exception of SC, all springs in this study arise from at or near the base of the St. Louis limestone, the regional base of the primary carbonate sequence in the Carboniferous. Data from the samples are summarized in Tab. 1. ese data comprise ionic measurements using HACH eld titrations for bicarbonate (HCO 3 ) as well as Ca 2+ and Mg 2+ concentrations in the 2012 samples, ion chro matography for 2010 samples, and a HACH DR 2800 spectrophotometer for select ions in the 2012 samples. Field measurements of pH and temperature (T) complement these data. Samples for ion analysis were ltered using a 0.45-m membrane and collected us ing 250 mL HDPE bottles and kept at 4C until time of analysis. Samples for cations were preserved using 2 mL of 6N HNO 3 Precision and accuracy were not reported from the lab contracted for the ion analyses. Computed charge balance values are reported in Tab. 1 for samples where appropriate. For each ltered sample, a split was stored in a 30mL glass bottle, treated with CuSO 4 as an anti-micro bial agent, and stored at 4C with a paralm seal until analysis for t 13 C DIC A similar split sample was analyzed for t 18 O and t 2 H with results summarized in Florea (2013b). Concurrent with each of the 16 sets of 2010 2011 data from SC and SS, samples for measurement of DOC and t 13 C DOC were collected in a 1L HDPE bottle spiked with 1 mL of 12 N HCl to prevent microbial ac tivity. In the lab, these samples were dehydrated and the remaining solids treated with H 2 SO 3 to eervesce CO 2 In addition to the springs from epigene karst aqui fers, preliminary geochemical investigations in the Ot ter Creek watershed by Dugan et al. (2012) provide a rst look at geochemical data of water chemistry from travertine (tufa) springs and sulfur seeps, that are in part spatially controlled by the Sunnybrook Anticline (Fig. 1). is anticline has an amplitude of approximately 30 m, is oriented N-NE, and is parallel to the trend of the Cumberland Escarpment. W aters at the tufa springs can be oversaturated with respect to calcite and where they emerge, calcite precipitates. ese springs probably resurge from long, strike-parallel owpaths on the west ank of the anticline. Similarly, some caves in the region contain signicant travertine deposits within active wa ter passages, including rimstone dams and owstone, suggesting periods of calcite oversaturation linked with chemical changes that lead to mineral precipitation. Florea (2013a) investigates the nature of water chemistry in the Otter Creek watershed, in particu lar from the Redmond Creek karst aquifer (Fig. 1), and using a comparison of reaction products and principal component analysis concludes that, although dissolution via Equation 1 dominates the chemistry, dissolution via Equation 2 with the sulfur derived from the entrain ment of shallow brines, is possible at the local scale. is process is particularly important adjacent to sulfur seeps in caves and streams. ese documented sulfur seeps are largely concentrated on the east ank of the Sunnybrook Anticline in the direction of the Appalachian Basin (Fig. 1). e presence of these seeps is a manifestation of shallow petroleum reservoirs in lower Mississippian strata that underlie the carbonates that host the karst aquifers (Fig. 2). I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY

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ACTA CARSOLOGICA 42/2-3 2013 282 from the fraction of inorganic carbon in the solid resi due. Secondary dehydration provided organic resin for isotopic analysis. Measurements of the stable isotopes of carbon were conducted at the Stable Isotope Laboratory at the Uni versity of South Florida. For t 13 C DIC 12 mL borosilicate glass vials with septa were lled with 1 mL of 85% phos phoric acid. Aer a ush of helium gas, 0.2 mL of sample was introduced into the vial with a syringe. e CO 2 produced in the reaction between the acid and the sam ple DIC was analyzed using a ermo-Finnigan IRMS attached to a Gas Bench II. For t 13 C DOC weighed masses of organic resin were loaded into tin capsules and com busted in an Elemental Analyzer. e CO 2 produced in the combustion was analyzed in the IRMS. Isotopic val ues were compared to the VPDB standard using simulta neous measurements from reference samples of Carrara marble and NBS 18 standards for t 13 C DIC and Fergie-CN for t 13 C DOC Measurements of t 13 C DIC and t 13 C DOC are accurate to within .1. LEE J. FLOREA Ca2+Mg2+K+Na+HCO3 -Fl-Cl-NO3 -SO4 2-DO C13CDOC 13CDIC (oC) mmol/L mmol/L mmol/L mmol/L mmol/L mmol/L mmol/L mmol/L mmol/L mg/L V PDB VPDB SS NL070710A 16.33 7.95 1.120 0.225 0.026 0.155 1.61 0.008 0.137 0.053 0.088 -26.5 -7.9 SS NL072010A 12.54 7.08 0.948 0.196 0.024 0.1 12 1.51 0.008 0.090 0.038 0.079 1.6 -27.8 -12.4 SS NL080310A 14.60 7.82 0.987 0.197 0.024 0.178 1.87 0.008 0.139 0.040 0.099 1.7 -26.5 -7.5 SS NL081710A 13.80 7.69 1.259 0.219 0.026 0.092 2.39 0.008 0.087 0.075 0.067 1.0 -26.9 -8.0 SS NL0831 10A 14.55 7.38 1.057 0.195 0.023 0.108 0.007 0.081 0.050 0.063 0.8 -26.5 -8.1 SS NL091410A 14.80 7.46 1.154 0.219 0.024 0.148 2.09 0.007 0.125 0.054 0.094 0.6 -27.0 -8.8 SS NL092810A 15.21 7.95 0.413 0.204 0.002 0.206 2.05 0.003 0.205 0.064 0.086 0.7 -26.0 -9.2 SS NL101210A 14.34 7.91 1.290 0.230 0.024 0.260 2.61 0.004 0.245 0.070 0.098 0.5 -27.8 -7.3 SS NL102610A 14.24 7.93 1.300 0.239 0.025 0.260 2.52 0.004 0.261 0.076 0.124 0.5 -26.5 -8.5 SS NL1 10910A 13.44 7.52 1.053 0.197 0.024 0.163 0.003 0.154 0.059 0.086 1.1 -29.2 -8.6 SS NL1 12310A 12.18 7.75 0.914 0.168 0.025 0.083 1.83 0.004 0.080 0.042 0.079 1.7 -26.5 -9.9 SS NL120710A 0.697 0.136 0.020 0.188 1.61 0.012 0.170 0.032 0.099 0.5 -27.1 -8.6 SS NL1221 10A 10.50 8.34 0.582 0.1 13 0.017 0.208 0.99 0.002 0.199 0.041 0.1 17 0.7 -27.8 -9.0 SS NL010411 A 10.22 7.47 0.523 0.104 0.016 0.179 1.77 0.001 0.172 0.030 0.107 1.1 -26.9 -9.7 SS NL01 181 1A 10.58 7.73 0.651 0.151 0.016 0.1 13 1.39 0.002 0.099 0.022 0.093 1.6 -26.2 -8.4 SS NL020811 A 9.64 0.557 0.135 0.015 0.130 1.09 0.003 0.1 12 0.018 0.099 0.5 -27.4 -9.3 SC NL070710B 14.27 6.40 0.226 0.102 0.015 0.031 0.30 0.007 0.032 0.020 0.039 -7.0 SC NL072010B 18.07 6.93 0.097 0.035 0.014 0.021 0.19 0.007 0.028 0.012 0.033 2.1 -8.9 SC NL080310B 17.06 7.47 0.167 0.049 0.015 0.024 0.007 0.029 0.015 0.036 1.3 -8.4 SC NL081710B 16.12 7.07 0.225 0.060 0.017 0.031 0.47 0.007 0.031 0.019 0.035 0.2 -8.2 SC NL0831 10B 16.85 7.31 0.193 0.060 0.014 0.032 0.006 0.034 0.017 0.042 0.2 -27.1 -7.6 SC NL091410B 15.95 7.65 0.147 0.039 0.014 0.017 0.37 0.006 0.041 0.027 0.038 1.1 -26.4 -7.3 SC NL092810B 15.50 8.03 0.177 0.049 0.013 0.033 0.38 0.002 0.031 0.01 1 0.035 0.9 -7.5 SC NL101210B 14.14 7.69 0.200 0.049 0.013 0.031 0.45 0.003 0.032 0.016 0.038 0.5 -6.3 SC NL102610B 13.53 7.49 0.1 19 0.033 0.027 0.021 0.31 0.002 0.034 0.024 0.042 2.5 -7.7 SC NL1 10910B 1 1.89 7.1 1 0.127 0.032 0.016 0.026 0.002 0.025 0.006 0.039 1.1 -26.5 -6.7 SC NL1 12310B 1 1.14 7.47 0.064 0.035 0.024 0.023 0.27 0.002 0.022 0.004 0.039 3.0 -28.1 -9.4 SC NL120710B 0.071 0.042 0.012 0.023 0.21 0.003 0.020 0.004 0.047 0.7 -7.7 SC NL1221 10B 5.05 7.08 0.073 0.042 0.01 1 0.021 0.13 0.001 0.021 0.006 0.056 0.7 -8.2 SC NL010411 B 4.97 6.97 0.069 0.045 0.01 1 0.025 0.81 0.001 0.017 0.002 0.054 0.9 -8.0 SC NL01 181 1B 3.19 7.19 0.072 0.045 0.010 0.024 0.33 0.001 0.020 0.001 0.051 0.7 -26.6 -8.1 SC NL020811 B 3.84 7.34 0.071 0.047 0.01 1 0.023 0.33 0.002 0.017 0.01 1 0.052 0.7 -8.0 SB LF05131 1B 13.84 6.94 6.849 0.706 0.034 1.036 2.70 0.017 0.988 0.003 3.518 -5.7 BH LF05151 1D 13.09 7.13 3.602 0.486 0.029 0.798 2.10 0.016 0.348 0.015 2.246 -4.5 BC1 LF05151 1A 15.20 6.59 6.084 1.536 0.1 13 4.794 4.11 0.021 1.893 0.000 4.302 -5.6 BCA LF012012B 12.50 8.26 0.970 0.148 2.17 0.052 -1 1.7 TC LF012012C 12.20 7.97 0.798 0.086 1.84 0.135 -13.1 CT LF012012D 12.50 7.64 1.180 0.156 1.66 0.687 -1 1.4 SS LF012012E 10.80 7.93 0.571 0.060 1.22 0.125 -8.7 HB LF012012F 1 1.70 7.66 1.170 0.092 2.52 0.073 -12.9 SSS LF012012G 1 1.30 7.80 0.535 0.1 16 1.74 0.062 -1 1.2 SCS LF012012H 12.20 7.83 0.621 0.060 1.34 0.062 -1 1.3 DH1 LF012012I 10.90 7.93 0.551 0.052 1.19 0.187 -10.6 DH2 LF012012J 10.80 7.90 0.607 0.056 1.04 0.198 -10.6 SB LF012012K 13.80 7.16 6.679 0.240 3.27 6.870 -10.7 OC3 LF012012L 13.40 7.48 8.176 0.336 4.00 7.91 1 -8.9 DC LF012012M 6.90 7.31 7.257 1.219 3.92 8.1 19 -9.3 LH LF012012O 13.80 8.22 1.521 0.140 2.98 0.094 -12.2 Cations Anions Carbon Site Sample Name T emp pH Table 1. Selected field, ion, and carbon data from sample collection. Chg Bal % 18 14 6 6 8 27 3 4 3 4 5 20 1 4 23 1 2 13 1 3 13 21 12 2 56 25 26 21 13 16 1 9 6 8 3 18 4 13 4 10 8 9 2 D ate 07/07/10 07/20/10 08/03/10 08/17/10 08/31 //1 0 09/14/10 09/28/10 10/12/10 10/26/10 1 1/09/10 1 1/23/10 12/07/10 12/21/ 10 01/04/11 01 / 18/11 02/08/11 07/07/10 07/20/10 08/03/10 08/17/10 08/31/ 10 09/14/10 09/28/10 10/12/10 10/26/10 1 1/09/10 1 1/23/10 12/07/10 12/21/ 10 01/04/11 01 / 18/11 02/08/11 05/13/11 05/15/11 05/15/11 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 Tab. 1. Selected eld, ion, and carbon data from sample collection. ANALYSIS OF ANALYTICAL RESULTS e saturation index with respect to calcite and gyp sum and the calcite saturation ratio for each sample were computed using the Debye Hckle relationship to compute molar activities of ionic species as summa rized by Ford and W illiams (2007) and W hite (1988) for the carbonate equilibrium reactions. Additionally for each sample, the partial pressure of CO 2 in solution, PCO 2 was computed using the temperature-dependent equilibrium equations for pK CO2 pK 1 pK 2 and pK c summarized in Ford and W illiams (2007, p. 48) and derived from Garrels and Christ (1965) and Plummer and Blusenberg (1982). Using Henry Law, the equiva lent molar concentration of CO 2 in solution was com puted from PCO 2 e sum of the molar concentrations of CO 2 and HCO 3 1 in solution is a measure of the total DIC in solution. Summaries of computed geochemical values are provided in Tab. 2.

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ACTA CARSOLOGICA 42/2-3 2013 283 I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY SIcal C/CsP CO2CO2DIC atm mmol/L mmol/L SS NL070710A 0.06 0.99 9.86E-04 4.72E-05 1.653 SS NL072010A -0.90 0.45 6.59E-03 3.18E-04 1.828 SS NL080310A -0.06 0.90 1.52E-03 6.89E-05 1.939 SS NL081710A 0.01 0.95 2.57E-03 1.20E-04 2.510 SS NL0831 10A SS NL091410A -0.31 0.73 3.88E-03 1.75E-04 2.265 SS NL092810A SS NL101210A 0.27 1.18 1.70E-03 7.79E-05 2.688 SS NL102610A 0.28 1.19 1.56E-03 7.19E-05 2.592 SS NL1 10910A SS NL1 12310A -0.17 0.83 1.70E-03 8.30E-05 1.913 SS NL120710A SS NL1221 10A -0.02 0.93 2.34E-04 1.20E-05 1.002 SS NL010411 A -0.68 0.55 3.09E-03 1.60E-04 1.930 SS NL01 181 1A -0.44 0.67 1.34E-03 6.84E-05 1.458 SS NL020811 A SC NL070710B -2.85 0.10 6.57E-03 3.02E-04 0.602 SC NL072010B -2.86 0.09 1.30E-03 5.35E-05 0.243 SC NL080310B SC NL081710B -1.99 0.19 2.25E-03 9.80E-05 0.568 SC NL0831 10B SC NL091410B -1.68 0.24 4.67E-04 2.04E-05 0.390 SC NL092810B -1.22 0.36 1.99E-04 8.80E-06 0.389 SC NL101210B -1.43 0.30 5.06E-04 2.33E-05 0.473 SC NL102610B -2.01 0.19 5.51E-04 2.58E-05 0.336 SC NL1 10910B SC NL1 12310B -2.35 0.1 4 4.91E-04 2.47E-05 0.295 SC NL120710B SC NL1221 10B -3.00 0.09 5.43E-04 3.29E-05 0.163 SC NL010411 B -2.36 0.1 4 4.33E-03 2.63E-04 1.073 SC NL01 181 1B -2.50 0.13 1.05E-03 6.73E-05 0.397 SC NL020811 B -2.36 0.1 4 7.47E-04 4.70E-05 0.377 SB LF051311 B -0.11 0.87 1.51E-02 7.03E-04 3.403 BH LF051511 D -0.25 0.78 7.76E-03 3.69E-04 2.469 BC1 LF051511 A -0.31 0.74 5.17E-02 2.31E-03 6.420 BCA LF0120312B 0.44 1.35 6.25E-04 3.02E-05 2.200 TC LF0120312C 0.01 0.95 1.03E-03 5.04E-05 1.890 CT LF0120312D -0.23 0.79 1.97E-03 9.53E-05 1.755 SS LF0120312E -0.34 0.72 7.46E-04 3.79E-05 1.258 HB LF0120312F -0.02 0.94 2.85E-03 1.41E-04 2.661 SSS LF0120312G -0.35 0.71 1.44E-03 7.22E-05 1.812 SCS LF0120312H -0.37 0.71 1.05E-03 5.11E-05 1.391 DH1 LF0120312I -0.37 0.71 7.28E-04 3.69E-05 1.227 DH2 LF0120312J -0.41 0.68 6.80E-04 3.46E-05 1.075 SB LF0120312K 0.16 1.08 1.09E-02 5.08E-04 3.778 OC3 LF0120312L 0.63 1.59 6.28E-03 2.96E-04 4.296 DC LF0120312M 0.40 1.33 8.48E-03 4.85E-04 4.405 LH LF0120312O 0.71 1.68 9.43E-04 4.39E-05 3.024 Site Sample Name Table 2. Summary of analytical data. D ate 07/07/10 07/20/10 08/03/10 08/17/10 08/31 //1 0 09/14/10 09/28/10 10/12/10 10/26/10 1 1/09/10 1 1/23/10 12/07/10 12/21/ 10 01/04/11 01 / 18/11 02/08/11 07/07/10 07/20/10 08/03/10 08/17/10 08/31/ 10 09/14/10 09/28/10 10/12/10 10/26/10 1 1/09/10 1 1/23/10 12/07/10 12/21/ 10 01/04/11 01 / 18/11 02/08/11 05/13/11 05/15/11 05/1511 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 01/20/12 SIgyp -2.65 -2.75 -2.65 -2.73 --2.61 -2.57 -2.46 -2.77 -2.77 -2.86 -2.82 -3.59 -3.99 -3.64 -3.77 -3.73 -3.65 -3.81 -4.11 -3.88 -3.95 -3.93 -3.94 -0.53 -0.91 -0.54 -2.99 -2.59 -1.76 -2.74 -2.72 -3.07 -3.00 -2.58 -2.51 -0.29 -0.18 -0.22 -2.52 Tab. 2. Summary of analytical data.

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ACTA CARSOLOGICA 42/2-3 2013 284 Fig. 3: Analytical measurements of A) 13 C DIC B) 13 C DOC and C) DOC compared against mea sured value of discharge with Stream Cave (SC) shown as open squares and Sandy Springs (SS) shown as open diamonds. A linear regression is shown for values of DOC at SC in panel C. Matched logarithmic regressions for values of 13 C DIC are shown for SC and SS in panel A. Most generally, the results of this study reveal the blend of dissolved carbon that is transmitted though the large ly epigene karst of the Cumberland Plateau. ese data further reveal information on the source and timing of carbon transport. As a baseline for further research, the data characterize the magnitude of carbon ux in present conditions. Since rates of chemical weathering are sensi tive to environmental and climate conditions, reference studies as this provide one important regional context to help understand the potential impacts of changes in land use (Zhang 2011) and atmospheric CO 2 (Cao et al. 2012) upon temperature, precipitation, plant community struc ture, and resultant DOC and DIC in karst aquifers. DOC AND t 13 C DOC Values of DOC are within the expected range of karst waters as observed by Simon et al. (2007). Values at the aquifer input (SC) are consistently higher (aver age and standard deviation = 1.110.82 mg/L) than at the aquifer output (SS) (average and standard devia tion = 0.97.47 mg/L). is suggests a conversion of DOC into DIC between aquifer input and output that is at least in part due to microbial activity. DOC val ues at SC are particularly sensitive to discharge (Fig. 3). During August 2010 when dry conditions reduced the base ow at SC to less than 0.8 L/s, DOC levels were close to zero. In contrast, higher discharge conditions LEE J. FLOREA VARIATION OF CARBON IN KARST A QUIFERS 0 y = 0.40x 0.56 R2 = 0.65 0 1 2 3 10 20 30 40 50 DOC (mg/L)Discharge (L/s) -30.0 -29.0 -28.0 -27.0 -26.0 -25.0 y = -0.76ln(x) 6.5 R2 =0.24 y = -0.76ln(x) 6.5 R2 = 0.36 -13.0 -12.0 -11.0 -10.0 -9.0 -8.0 -7.0 -6.0 -5.0 13CDIC ( VPDB) 60 70 80 13CDOC ( VPDB) A B C

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ACTA CARSOLOGICA 42/2-3 2013 285 in October and November 2010, combined with organic loading from falling deciduous leaves, produced DOC values up to 3 mg/L. No similar relationship between DOC and Q is visible at SS; however, similar timing of high and low values of DOC is clear from the time se ries data (Fig. 3). Despite scatter, the contribution of C3-type veg etation is clearly demonstrated by the values of t 13 C DOC observed at both SC and SS (average and standard devi ation = -27.0.8). No relationship between t 13 C DOC and Q is visible in Fig. 3, suggesting that the source of DOC does not shi either seasonally from agriculture or according to availability of water. is conforms to existing land use in the watersheds for both SC and SS, which are minimally impacted by human activities, being primarily woodland with a history of selected timber harvesting dating back to the early 1900s (Kay Koger, landowner personal communication). e estimated population density for Redmond Creek is ~ 0.01 ha -1 and is a similar density for the larger Otter Creek watershed. t 13 C DIC AND DIC CONCENTRATION VERSUS DISCHARGE t 13 C DIC values from the Otter Creek watershed gener ally coincide with carbonic acid dissolution of limestone. e concentration of this DIC generally decreases with increasing Q Data from SC and SS in 2010 weakly suggest that t 13 C DIC at the aquifer output is inversely pro portional with discharge (Fig. 3) indicating that undersaturated waters during higher discharge conditions have excess soil-derived PCO 2 and contribute more depleted t 13 C DIC values than during base ow conditions. Con versely, base ow conditions are more enriched in 13 C, with weak logarithmic regressions for SC and SS that ap proach -6.5 at a Q = 1 L/s (Fig. 3). t 13 C DIC values are somewhat dependent upon the DIC concentration at SC and SS, but not among the other sites in this dataset, including tufa springs and sul fur seeps (Fig. 5). DIC in the form of HCO 3 dominates the water chemistry (Tab. 1 and 2). e concentration of DIC among all samples increases with calcite saturation and, on average, exceeds 2 mmol/L in saturated waters (Fig. 5). Along the owpath between SC and SS, DIC in CO 2 form is converted to HCO 3 via the reaction in Equa tion 1. Careful inspection of data in Tab. 2 demonstrates greater CO 2 concentration at SS when compared to SC (average and standard deviations = 1.1x10 4 .85x10 4 and 0.8x10 4 .98x10 4 respectively); however, at SC the CO 2 comprises a greater fraction of the DIC in the sample (Fig. 5). t 13 C DIC VERSUS CALCITE SATURATION W alden (1999), during an investigation of groundwa ter from the Redmond Creek karst aquifer, used ratios of 87 Sr/ 86 Sr to dierentiate between conduit and diuse ow. Her results expectedly found a trend toward val ues similar to the host limestone along the aquifer ow path. Following this logic, values of SI and C/Cs should in theory approach, or even exceed, saturation along that same owpath. Data from Tab. 2 demonstrate that this is likely the case with values of SI at SC (average and stand ard deviation = -2.62.59), an input to the Redmond Creek aquifer, consistently lower than at SS (average and standard deviation = -0.56.39), the primary output for the Redmond Creek aquifer. I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY -5 -4 -3 -2 -1 0 1 2 -5 -4 -3 -2 -1 0 1 2 SI calcite SI gypsum Oversaturated gypsum Undersaturated calcite Oversaturated gypsum Oversaturated calcite Undersaturated gypsum Undersaturated calcite Undersaturated gypsum Oversaturated calcite Fig. 4: Calculated values of satu ration index for calcite (x-axis) and gypsum (y-axis) at Stream Cave (SC open squares), Sandy Springs (SS open diamonds), and sulfur seeps (open circles) from 2010. Also shown are DIC measurements from 2012 samples including those from karst springs (lled diamonds), the tufa spring (lled triangle), and the sulfur seeps (lled cir cles). Undersaturated and over saturated elds are identied.

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ACTA CARSOLOGICA 42/2-3 2013 286 LEE J. FLOREA -14.6 y = 2.9x 8.1 R2= 0.62 y = 0.68x 6.3 R2= 0.24 -14.60 -13.60 -12.60 -1 1.60 -10.60 -9.60 -8.60 -7.60 -6.60 -5.60 -4.60 -3.60 -4.00 -3.50 -3.00 -2.50 -2.00 -1.50 -1.00 -0.50 0.00 0.50 1.00 13CDIC ( VPDB) SI calcite y = 0.89ln(x) 6.3 R2 = 0.24 y = 3.73ln(x) 8.1R2 = 0.62 -20.6 -19.6 -18.6 -17.6 -16.6 -15.6 -14.6 -13.6 -12.6 -1 1.6 -10.6 -9.6 -8.6 -7.6 -6.6 -5.6 -4.6 -3.6 DIC (mmol/L) DIC (mmol/L) 13CDIC ( VPDB) 13CDIC ( VPDB) HCO3 1(mmol/L )-13.6 -12.6 -1 1.6 -10.6 -9.6 -8.6 -7.6 -6.6 -5.6 -4.6 -3.6 -14.6 -13.6 -12.6 -1 1.6 -10.6 -9.6 -8.6 -7.6 -6.6 -5.6 -4.6 -3.6 13CDIC ( VPDB ) 0 0.5 1 1.5 2 2.5 3 0 1 2 3 4 5 6 7 0 0.2 0.4 0.6 0.8 1 1.2 1.4 1.6 1.8 0 0.5 1 1.5 2 2.5 3 3.5 4 4.5 A B C D E -5 -4.5 -4 -3.5 -3 -2.5 -2 -1.5 -1 -0.5 0 0.5 1 SI gypsum -14.60 -13.60 -12.60 -1 1.60 -10.60 -9.60 -8.60 -7.60 -6.60 -5.60 -4.60 -3.60 13CDIC ( VPDB) y = 4.11x 7.9 R2= 0.64 y = 5.30x 5.4 R2= 0.19 FC/Cs y = 3.6x 10.7 R2= 0.80 y = 16.5x 5.8 R2= 0.99 y = 4.77ln(x) 11.0R2 = 0.93 Correspondingly, values of t 13 C DIC should increasingly reect DIC derived from the bedrock. W ith an approximated shi of +6.4 from CO 2 in the soil to CO 2 in so lution (Clark and Fritz 1997), water with C/Cs = 0 that is in equilibrium with soil CO 2 should have a t 13 C DIC of -20.6. eoreti cally, karst water with a SI = 0 or C/Cs = 1 would have a 50% blend of DIC from the soil and from the bedrock following Equation 1. W ithout prior knowledge of the bedrock end member, the value of t 13 C DIC expected at sat uration is unknown. However, the data from both SC and SS appear to converge at satura tion to a t 13 C DIC value of -6.3 and -8.1 for SC and SS, respectively (Fig. 5), which is similar to the values observed at base ow conditions in Section 5.3 (Fig. 3). Assuming that Equation 1 is the only reaction pathway for producing DIC, these results would im ply a bedrock source of +8.4 for SC and +5.8 for SS, considerably more enriched in 13 C than typical values for marine car bonates that are typically close to that of the VPBD standard (e.g., Hoefs 1997). Since t 13 C of carbonates may vary by several per mille, a comparative investigation of the t 13 C signal of bedrock is currently under way to quantify the additional enrichment, if present. Fig. 5: Analytical measurements of dissolved in organic carbon (DIC) at Stream Cave (SC open squares), Sandy Springs (SS open diamonds), and sulfur seeps (open circles) from 2010. Also shown are DIC measurements from 2012 samples including those from karst springs (lled diamonds), the tufa spring (lled triangle), and the sulfur seeps (lled circles). From top to bottom: A) V alues of 13 C DIC at SC and SS are compared against DIC concentration in the form of CO 2 (open symbols) and H CO 3(gray-lled symbols). B) V alues of 13 C DIC are compared against total DIC concentration. C) V alues of 13 C DIC are com pared against C/Cs with logarithmic regressions shown for data from SC and SS. D) DIC concentra tion in the form of H CO 3are compared against the calcite saturation ratio (C/Cs). E) V alues of 13 C DIC are compared against the saturation index with respect to calcite; linear regressions are shown for the data from SC and SS. F) V alues of 13 C DIC are compared against the saturation index with respect to gypsum; linear regressions are shown for the data from SC and SS.

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ACTA CARSOLOGICA 42/2-3 2013 287 C ONCLUSIONS In the Otter Creek watershed of the Cumberland Plateau of Kentucky, USA, the concentration of dissolved organ ic carbon is less than 3 mg/L, and in the case of some sites is inversely proportional to discharge. C3 vegetation is the source of this DOC, with an average t 13 C DOC of DIC concentrations that may exceed 4 mmol/L are inversely proportional to discharge and directly pro portional to the saturation of the aqueous solution with respect to calcite and gypsum at some sites. Values of t 13 C DIC reect soil and bedrock sources and conform to measurements expected from the carbonate equilibrium reactions. Increasing enrichment of 13 C at two sites in this study proceeds with increased saturation of groundwa ter with respect to calcite. Sulfur seeps follow this same trend at higher base ow conditions. ese same sulfur seeps are enriched in 13 C at low ow, but undersaturated, thus suggesting that the oxidation of reduced sulfur may enrich these sties in bedrock DIC. Other locations appear to follow independent DIC pathways. Some sites, such as tufa springs may have depleted t 13 C DIC from CO 2 degas sing and calcite precipitation. One possible source of additional t 13 C DIC enrich ment is evident in the data from the sulfur seeps. ese sites are closer to saturation with respect to gypsum (Fig. 4). In 2011 these sites, despite high values of Ca 2+ and HCO 3 in solution (Tab. 1), are undersaturated with respect to calcite (Tab. 2). ey are correspondingly enriched with respect to 13 C (Fig. 5). is implies ad ditional enrichment in bedrock carbon without a corre sponding additional contribution from PCO 2 from the soil. Such a process may be driven by the oxidation of dissolved suldes, which can decrease the pH and drive additional carbonate dissolution via Equation 2 Interest ingly, the 2012 data from the sulfur seeps reveal oversat uration with respect to calcite and decreased enrichment of 13 C. ese samples were collected in higher discharge conditions during the winter months. ey follow trends similar to SC and SS and converge on a value t 13 C DIC of -11 at SI = 0 and C/Cs = 1 (Fig. 5). Collectively these data suggest the potential for lessened oxidation of re duced sulfur at the time of the 2012 samples. e particular reason why the 2012 sulfur seep samples are more depleted in 13 C compared to the samples from SS, which are in turn more depleted in 13 C than the samples from SC is unclear. Logic suggests that the trend should be the opposite; the potential in uence of Equation 2 upon t 13 C DIC should increase at sites with known inuence of sulfur-rich brines, which is certainly the case the sulfur seeps and observed to a lesser degree at sites within the cave that contributes to SS. One possible mechanism that could generate the trends in Fig. 5 is that the average bedrock t 13 C DIC may vary by several per mille within the contributing water shed for each site. A second possibility is the potential for the oxidation of pyrite within the bedrock within the watershed of SC, known to be present in signicant quantities in the Bangor Limestone and the organicrich shale and coal that overlie the Bangor. A third pos sibility is the potential for methanogenesis within the deeper ow systems that contribute some percentage of ow to SS and the sulfur seeps where reduction of organics may lead to depleted values of t 13 C DIC Finally, the samples from SC, which represent waters that have limited interaction with bedrock, may simply not be in equilibrium with soil CO 2 Each of the postulates re mains untested at this time. Of the remaining samples during 2012 from all other sites, the values of t 13 C DIC are equal to, or more de pleted than those observed at SS in 2010 (Fig. 5). Five of those sites, CT, SSS, SCS, DH1, DH2, are clus tered and similar in range to the data from SS. ese ve springs have similar physical characteristics (strati graphic position, discharge, etc.) to SS. Four other sites, BCA, TC, HB, and LH, are depleted in 13 C compared to the other samples at the same calcite saturation. ese springs may collect waters from aquifers that have some what dierent bedrock isotope composition than SS or each other. e sample from LH is signicantly oversatu rated with respect to calcite and is an active tufa spring. Fractionation of the heavier isotope during CaCO 3 pre cipitation combined with degassing of CO 2 enriched in the lighter isotope potentially drives this sample toward depletion. e remaining three depleted sites may also experience similar fractionation from degassing and cal cite precipitation to varying degrees and underscores the complexity of interpreting a limited number of t 13 C DIC values from a site. I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY

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ACTA CARSOLOGICA 42/2-3 2013 288 REFERENCES Abbott, W .R., 1921: Oil Field MapW ayne and Clinton Counties, Kentucky. Anthony, D.M. & D.E. Granger, 2004: A Late Tertiary or igin for multilevel caves along the western escarp ment of the Cumberland Plateau, Tennessee and Kentucky, established by cosmogenic 26 Al and 10 Be.Journal of Cave and Karst Studies, 66, 2, 46. Berner, R.A., Lasaga, A.C. & R.M. Garrels, 1983: e car bonatesilicate cycle and its eect on atmospheric carbon dioxide over the past 100 million years.American Journal of Science, 284, 641. Clark, I.D. & P. Fritz, 1997: Environmental Isotopes in H y drogeology.Lewis Publishers, Boca Raton, 328 pp. Crawford, N.C., 1984: Karst landform development along the Cumberland Plateau Escarpment of Tennessee. In: RG LeFleur (ed.), Groundwater as a geomorphic agent. Boston, Allen and Unwin, Inc., pp. 294. Dugan, C.R., Florea, L.J. & W .D. W alden, 2012: A Geo chemical Investigation of Springs within the Ot ter Creek watershed: W ayne County, Southeastern Kentucky.Geological Society of America Abstracts with Programs, 44, 4, 26. Ehleringer, J.R., Buchmann, N. & L.B. Flanagan, 2000: Carbon isotope ratios in belowground carbon cycle processes.Ecological Applications, 10, 2, 412. Ettensohn, F.R., Rice C.R., Dever, G.R. Jr. & D.R. Ches nut, 1984: Slade and Paragon Formations; New stratigraphic nomenclature for Mississippian rocks along the Cumberland Escarpment in Kentucky.U.S. Geological Survey Bulletin 1605B, 37 pp. Florea, L.J.l., 2013: Selective recharge and isotopic com position of shallow groundwater within temperate, epigenic carbonate aquifers. J ournal of H ydrology 489: 201. doi 10.1016/j.jhydrol.2013.03.008 Florea, L.J., 2013a: Investigations into the Potential for Hypogene Speleogenesis in the Cumberland Pla teau of Southeast Kentucky, U.S.A. In, P roceedings of the 16 th International Congress of Speleology, Brno, Czeck Republic, J uly 2013, pp. 356. Ford, D.C. & P. W illiams, 2007: Karst H ydrogeology and Geomorphology.John W iley & Sons, 562 pp. Garrels, R.M. & C.M. Christ, 1965: Solutions, Minerals, and Equillibria. Harpers Geoscience Series. Harper and Row, New York. 450 pp. Hill, C.A., 1990: Sulfuric acid speleogenesis of Carlsbad Cavern and its relationship to hydrocarbons, Dela ware Basin, New Mexico and Texas.American As sociation of Petroleum Geologists Bulletin, 74, 11, 1685. Hoefs, J., 1997: Stable isotope geochemistry.; 4 th Ed., Springer, New York, 200 pp. Cao, J.H., Yuan, D.X., Groves, C., Huang, F., Huang H. & Q Lu, 2012: Carbon uxes and sinks: e consump tion of atmospheric and soil CO 2 by carbonate rock dissolution.Acta Geologica Sinica, 86, 963. Kendall, C. & E.A. Caldwell, 1998: Fundamentals of Iso tope Geochemistry. In: Kendall, C. & J. J. McDon nell, (eds.), Isotope Tracers in Catchment H ydrology Elsevier Science B.V ., Amsterdam, pp. 51. Marlier, J.F. & M.H. OLeary, 1984: Carbon kinetic iso tope eects on the hydration of carbon dioxide and the dehydration of bicarbonate ion.Journal of the American Chemical Society, 106, 5054. Palmer, A.N., 1991: Origin and morphology of limestone caves.Geological Society of America Bulletin, 103, 1. e author acknowledges the major contributions by Bill W alden, Chasity Stinson, and Nick Lawhon during sample collection as well as the conversations, hospital ity, and access provided by landowners Tim Pyles and Kay Koger and cavers Rick Gordon, Deb Moore, Harry Gopel, and Eric W eaver. Finally, the author appreciates the collaborations with Jonathan W ynn and Bogdan Onac at the USF Stable Isotope Laboratory. Funding for this work was provided by a WKU Provost Incentive grant, a WKU start-up index, and a Ball State University start-up package. W elcome comments from two anony mous reviewers greatly improved the ow and content of this manuscript. A CKNO WLEDGEMENTS LEE J. FLOREA

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ACTA CARSOLOGICA 42/2-3 2013 289 I SOTOPES OF C ARBON IN A K ARST AQUIFER OF THE C UMBERLAND P LATEAU OF K ENTUCKY Plummer, L. & N.E. Busenberg, 1982: e Solubilities of calcite, aragonite and vaterite in CO 2 H 2 O solu tions between 0 and 90C, and an evaluation of the aqueous model for the system CaCO 3 CO 2 H 2 O.Geochimica et Geochimica Acta, 46, 1011 1040. Rantz, S.E., 1982: Measurement and computation of streamow, Volume 1, Measurement of stage and discharge.United States Geological Survey W ater Supply Paper 2175, 284 p. Sasowsky, I.D. & W .B. W hite, 1994: e role of stress re lease fracturing in the development of cavernous porosity in carbonate aquifers.W ater Resources Research, 30, 12, 3523. Schlesinger, W .H., 1997: Biogeochemistry: An Analysis of Global Change.Academic Press, San Diego, 443 pp. Simon, K.S., Pipan, T. & D.C. Culver, 2007: A conceptual model of the ow and distribution of organic car bon in caves.Journal of Cave and Karst Studies, 69, 2, 279. Simpson, L.C. & L.J. Florea, 2009: e Cumberland Pla teau of Eastern Kentucky.In: Palmer, A.N. & MV Palmer (eds.), Caves and Karst of America. Nation al Speleological Society, pp. 70. W alden, K., 1999: Strontium isotopes in Redmond Creek Cave, Monticello, Kentucky.Honors esis, Ohio State University, 66 p. W alden, W .D., W alden, K.M. & L.J. Florea, 2007: e Caves and Karst of Redmond Creek, W ayne Coun ty, Kentucky.National Speleological Society Con vention Program Guide, p. 107. W hite, W .B., 1988: Geomorphology and H ydrology of Karst Terrains New York, Oxford University Press, 464 pp. Zhang, C., 2011: Carbonate rock dissolution rates in dif ferent landuses and their carbon sink eect.Chi nese Science Bulletin, doi: 10.1007/s11434



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C ONTRIBUTION OF NON TROGLOBIOTIC TERRESTRIAL INVERTEBRATES TO CARBON INPUT IN HYPOGEAN HABITATS P RISPEVEK PREZIMUJO IH NETROGLOBIONTSKIH KOPENSKIH NEVRETEN ARJEV K VNOSU OGLJIKA V PODZEMELJSKE HABITATE Tone NOVAK 1 Franc JANEKOVI 1 & Saka LIPOV EK 1,2 Izvleek UDK 592:546.26:551.44 Tone Novak, Franc Janekovi & Saka Lipovek: Prispevek prezimujoih netroglobiontskih kopenskih nevretenarjev k vnosu ogljika v podzemeljske habitate V enajstih najpomembnejih vrstah kopenskih nevretenarjev v jamah Slovenije smo prouili razlike v biomasi, suhi masi, energijski vsebnosti in koliini ogljika v zimskem obdobju. Te podatke smo obravnavali v kombinaciji z abundanco v 54 jamah in rudnikih, da bi ocenili koliino organskega ogljika v njih ter vnos ogljika v te habitate. V srednji Evropi so najpo membneje vrste, ki vnaajo ogljik v podzemlje, Troglophilus cavicola, T. neglectus, Faustina illyrica, Amilenus aurantiacus in Scoliopteryx libatrix. V nasprotju s splono domnevo, prispe vajo poginuli osebki le 0,15 % C glede na skupno migrirajoo biomaso, ki ga troglobionti ne morejo neposredno uporabiti zaradi okub z entomopatogenimi glivami. Skupaj z uplenjeni mi osebki vnos ne presega 0,3 % biomase teh vrst. Te razmere je treba skrbno prouiti v vsakem krakem predelu, kjer se mi grirajoa favna razlikuje od ostalih obmoij. Kljune besede : biomasa, dihanje, jame, tok ogljika, netroglo biontske vrste, organski ogljik. 1 Department of Biology, Faculty of Natural Sciences and Mathematics, University of Maribor, Koroka cesta 160, SI-2000 Maribor, Slovenia, e-mail: tone.novak@uni-mb.si; franc.janzekovic@uni-mb.si 2 Medical Faculty, University of Maribor, Slomkov trg 15, SI-2000 Maribor, Slovenia, e-mail: sasa.lipovsek@uni-mb.si Received/Prejeto: 1.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 301, POSTOJNA 2013 Abstract UD 592:546.26:551.44 C Tone Novak, Franc Janekovi & Saka Lipovek: Contribu tion of non-troglobiotic terrestrial invertebrates to carbon in put in hypogean habitats Eleven of the most important terrestrial invertebrate species in Slovenian caves were analyzed for dierences in their fresh and dry biomass, energy content and carbon bulk during win ter. ese data were combined with the species abundance in 54 caves and adits in order to estimate their organic carbon bulk and carbon input into these habitats. In Central Euro pean caves, Troglophilus cavicola, T. neglectus, Faustina illyrica, Amilenus aurantiacus and Scoliopteryx libatrix are the most im portant vectors of carbon between the epigean and hypogean habitats. In contrast to the general assumption, carbon total contribution to caves via dead bodies is only 0.15% of total mi gratory biomass, and it is not directly available to troglobionts because of infection with entomopathogenic fungi. In winter, together with predated migratory specimens, carbon input does not exceed 0.3% of the total migratory biomass. is situ ation should be carefully examined in every karstic region in habited by distinctive fauna. Keywords: biomass, carbon ux, caves, non-troglobiotic spe cies, organic carbon, respiration. I NTRODUCTION Maintenance of life, manifested through metabolic proc esses, is directly involved in carbon cycling, a substantial domain for biological and ecological studies. In the last two decades, biosequestration, the capture and storage of atmospheric CO 2 by biotic processes and systems, has at tracted much attention. is is because biosequestration is closely related to global warming, which currently rep resents one of the main threats to the Earths biota, hu mans included, and which is caused in a large part by an thropogenic CO 2 emissions (U.S. DOE 2008; Houghton

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ACTA CARSOLOGICA 42/2-3 2013 302 T ONE NOVAK, FRANC JANEKOVI & SA KA LIPOV EK 2009; Riebeek 2011). e natural carbon cycle includes carbon ow through inorganic and organic compounds, i. e., through abiotic and biotic systems, although our understanding of changes in terrestrial biomass in this respect remains rather rudimentary (Perruchoud et al. 2000; Houghton et al. 2009). Soil organic carbon is the main global carbon stock in recent carbon cycling; more than half of cycling carbon is currently stored in soils (Perruchoud et al. 2000; Schmidt et al. 2011). In hypogean habitats, in the absence of photosyn thesis, carbon compounds are fewer and less diverse in comparison to epigean habitats (Culver & Pipan 2009). So far, chemoautotrophy has been conrmed in only a very few caves (Srbu et al. 1996, Engel 2005, Por 2007). e diversity and abundance of the biota in most oth er, heterotrophic hypogean habitats depend on alloch thonous energy resources. ese habitats, cave streams for example, are more likely to be carbon rather than nutrient-limited (Simon & Beneld 2002). Among het erotrophic organisms in these habitats, the ow of car bon includes the uptake of substrates rich in organic carbon and other nutrients as well as their prey, their metabolism and incorporation of nutrients into the body, and a permanent, gradual loss of carbon from the body through elimination and respiration. Organic matter is conveyed into the subterranean environment by water, gravity, wind or active move ments of organisms. W ater and soil are the main abi otic vectors supplying the subterranean area with food, whilst growing roots and migrating animals are the most important biotic resources (Tercafs 2001; Culver & Pipan 2009; Fong 2011). In addition, guano supports a specic subterranean community, the guanobionts (De harveng & Bedos 2000; Gnaspini 2005). Food takes the form of dissolved organic carbon (DOC), ne (FPOM) and coarse particulate organic matter (CPOM), eggs, ca davers, excrement and predated specimens. A conspicu ous amount of organic carbon is incorporated in nontroglobiotic animals (i. e., trogloxenes and troglophiles; see Novak et al. 2012 for comments) that move daily or seasonally between surface and subterranean habitats and supply food directly to the subterranean communi ties (Fong 2011). Careful accounting of inputs and outputs to a karst basin and measures of standing stocks and uxes (trans port and respiration) within basins are required to im prove our knowledge of the carbon cycle within karstic ecosystems (Simon et al. 2007). From this perspective, the contribution of terrestrial trogloxenic species, spe cies that periodically leave the cave, to subterranean communities has been underestimated (Fong 2011) and understudied. Till recently, most such studies have involved estimating the amount of excrement/guano or cadavers or eggs laid in caves (e. g., Jequier 1964; Taylor et al. 2005). Novak & Kutor (1982) published the rst measurements of the energy content in selected nontroglobiotic species. In dormancy, metabolism slows down consider ably, and development and reproduction are suppressed (Denlinger 2002). is period is then appropriate for estimating energy requirements in animals maintaining basal metabolism, i. e., the minimal loss of organic car bon via respiration. Some non-troglomorphic terrestrial animals have been studied for their energy metabolism during overwintering (Lipovek et al. 2004, 2008, 2009, 2011; Novak et al. 2004). In this contribution, our goal is to estimate how these taxa are involved in carbon ow with respect to their contribution of organic carbon to hypogean habi tats, and what is the carbon lost via respiration during their winter sojourn in these habitats. W e hypothesized that more abundant species and those with larger masses are more important in this respect. MATERIALS AND METHODS For this study, we combined two data sets. e rst is a eld data set on the more abundant invertebrates (list ed in Novak et al. 2012) from 51 caves and three abanfrom 51 caves and three abanfrom 51 caves and three aban doned mines in central and northern Slovenia (listed and mapped in Novak 2005; in the following: caves). e second data set includes biometric, biochemical and cy tological data provided by a separate study of 11of the more abundant non-troglomorphic invertebrate spe cies which inhabit these caves during winter or which are important for their large mass (Tab. 1). Fresh mass was used in calculating their biomass bulk, and their dry mass to assess their carbon bulk and loss during winter. e caves were investigated between 1977 and 2001. W e sampled in January, April, July and October in a total of 617 sampling sections, every 3.5 m, on average. Faunal records were provided on two visits within 45 hrs by observing the cave walls, ceiling and oor and ap plying standardized, baited pit-fall trapping and Berlese funnels. Altogether, 2468 records were provided, refer ring to 173,008 individuals of 600 estimated species in

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ACTA CARSOLOGICA 42/2-3 2013 303 RESULTS Table 2 shows individual fresh mass in the 10 species from the 54 caves plus Chilostoma from a cave in Ita ly and t-test between the series at the beginning and at the end of overwintering; their dry mass is presented in Tab. 3. Estimated average individual energy budgets are shown in Tab. 4. e selected species demonstrate vari ous types of life cycles with respect to active or inactive life in hypogean habitats (active state, quiescence, dia pause), strategies and the span of dormancy (Tab. 5), and strategies and level of eciency in using energy-supply ing compounds. Numbers of dead individuals and their dead biomass are also presented in Tab. 5. Fresh mass increased in most species overwintering deeper inside caves. Given this nding, the measurements for T. neglec tus should be repeated on account of their discrepancy. Both snails and both butteries most probably decreased their fresh mass, owing to the diminished water content, due to a drop in relative humidity in the cave entrance sections. Both Troglophilus moult during overwintering (Novak & Kutor 1983), but T. cavicola larvae had obvi ously fed before sampling at the end of overwintering, probably because the sampling date was possibly past the end of winter. e same was true in F. illyrica Aer overwintering, the total fresh mass of these 11 species increased by 6.71%, while their dry mass de creased by 10.09%, and energy by 10.57%. eir total organic carbon was 474.49 g at the beginning, and its respiration loss was 47.86 g by the end of winter. eir total dry necromass contributed 5.59 g of organic carbon to the caves, and yielded 1.07/0.93% of the fresh mass at C ONTRIBUTION OF NON TROGLOBIOTIC TERRESTRIAL INVERTEBRATES TO CARBON INPUT IN HYPOGEAN HABITATS total (details in Novak et al. 2004; Novak 2005). During winter, numbers of individuals in all the species varied by ca. 5% to 40%, mostly because of migration between the caves and adjacent habitats inaccessible to humans; this variation is not part of the present study. In order to avoid these disturbances, we estimated the biochemical changes as if the abundances did not change during over wintering. Data on dead organisms and excrement were estimated because the study was not designed to monitor these. Predated individuals are discussed in Novak et al. (2010). By 2013, 64 caves in total had been investigated in the same way (own unpublished data). In four of these monthly monitoring provided important supporting in formation. Specimens for biometrical, biochemical and cy tological study were collected between 2003 and 2011 from seven caves in Slovenia (locality centroid 46 N, 15 E, altitude 600 m), four among them from the list of 54 caves, and from a cave in northeast ern Italy (45 N, 13 E). Four among these caves were from the list of 54 caves. Specimens of each stage and sexdepending on the specieswere measured for their size, fresh and dry mass, and lipid, glycogen and water content at the beginning (November), the middle (January) and the end of overwintering (March). ree to ve additional specimens were analyzed histologi cally and cytologically in the control. For these analy ses, specimens were killed by exposure for two hours to C. e dry mass was determined aer 15 days of vacuum desiccation under P 2 O 5 (details in Lipovek et al. 2004, 2008, 2009, 2011; Novak et al. 2004). Invertebrate dry mass is ca. 50% carbon, on average (Salonen et al. 1976), we therefore estimated carbon content as half the dry mass, and the carbon loss as 50% the dry mass lost during overwintering. e energetic input was calcu lated upon the biomass and energetic data in selected speciesseven among those from the 11-species-list (No vak & Kutor 1982) (Tab. 4). T-tests were used to identify signicant dier ences in total fresh and dry mass between the beginning and the end of overwinter ing. e data analysis was carried out with the statisti cal soware SPSS 19.0. Tab. 1: Species investigated for their participation in carbon and energy uxes in subterranean environments in winter. Higher taxon Family Species Gastropoda Helicidae Chilostoma ( Josephinella ) lefeburiana (Frussac, 1821) Faustina illyrica (Stabile, 1864) Opiliones Phalangiidae Amilenus aurantiacus (Simon, 1881) Gyas annulatus (Olivier, 1791) Gyas titanus Simon, 1879 Araneae Tetragnathidae Meta menardi (Latreille, 1804) Coleoptera Carabidae Laemostenus schreibersii (Kster, 1846) Lepidoptera Geometridae Triphosa dubitata (Linnaeus, 1758) Noctuidae Scoliopteryx libatrix (Linnaeus, 1758) Orthoptera Rhaphidophoridae Troglophilus cavicola (Kollar, 1833) Troglophilus neglectus Krauss, 1879

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ACTA CARSOLOGICA 42/2-3 2013 304 the beginning/the end of overwintering, and 3.43/3.69% of the dry mass, respectively. Energy costs were ca. 0.76/ 0.71 kJ/g, if calculated from the biomass, and these values were 2.51/2.80 kJ/g for the dry mass. e estimat ed carbon input via dead invertebrate individuals con tributed 23.15 g fresh mass (which represented 0.74% of the total fresh mass), 11.18 g (1.18%) dry mass and 5.59 g C (1.18%) in the caves. All the cadavers were in fected by entomopathogenic fungi. Carbon lost via respi ration was 73.71 g, i. e., 2.34% of the fresh, and 7.75% of the dry mass at the beginning of overwintering. e total assemblage of the 11 species at the beginning and end of overwintering did not dier signicantly in fresh mass (t =1.17; p=0.245; df=253), but did dier signicantly in dry mass (t=2.28; p=0.023; df=253) and energy content (t=2.34; p=0.020; df=253). e relative importance of species with respect to their abundance, fresh and dry mass and energy bulk is presented in Tab. 6. Tab. 2: Individual fresh mass in the 11 species under investigation. t-test between the series at the beginning and at the end of overwin tering. Signicant p-values are in bold. Species N of individuals; beginningend of overwintering Beginning MeanSD Min-Max [mg] End MeanSD Min-Max [mg] Average dierence beginningend of over-wintering [%] t; p df F. illyrica 5 1944.3.6 1910.9.8 1672.5.2 1546.1.5 .98 6.50 <0.001 8 C. lefeburiana 5 2202.4.7 1901.1.5 1582.6.4 1209.8 .14 3.11 0.013 9 A. aurantiacus 5 14.7.6 12.2.3 16.1.1 15.2.4 9.52 1.70 0.128 8 A. aurantiacus 5 20.3.5 17.5.1 22.2.2 19.1 24.9 9.36 1.280 0.236 8 G. annulatus 10 29.89.29 22.5.2 30.72.24 26.2.7 2.78 0.38 0.712 17 G. titanus 21 24 42.9.13 32.7 44.8.0 29.8.3 4.43 0.43 0.667 43 M. menardi 10 248.2.2 164.8.8 290.6.5 171.6 432.2 17.08 1.13 0.271 20 L. schreibersii 5 61.8.2 49.7.4 60.6.4 36.3.1 .94 0.32 0.755 13 T. dubitata 10 58.8.6 46.6.4 50.6.4 39.2.6 .95 2.39 0.028 19 S. libatrix 10 261.1.6 223.4.4 228.8.7 182.5.4 .37 2.34 0.031 18 T. cavicola 6 477.8.2 402.1.4 596.1.1 521.2 24.76 3.76 0.005 9 T. cavicola 5 541.2.3 504.1.4 611.7.8 534.9.4 13.03 2.19 0.060 8 T. cavicola larvae 10 87.1.3 77.5.2 149.3.6 101.4.9 71.41 7.96 < 0.001 18 T. neglectus 4 504.7.9 407.7.3 442.5.7 325.3.7 .32 0.81 0.047 4 T. neglectus 5 649.3.0 428.4.7 350.8.8 71.4.8 .97 2.20 0.056 9 T. neglectus larvae 13 133.3.7 48.6.6 149.0.3 73.6.9 11.78 0.56 0.565 25 Excrement of both snail species, the two Troglophilus and some other non-troglobionts, like Diplopoda and Collembolawere relatively abundant close behind the entrance, but were very rare deeper within the caves. In the warm half of the year, e. g., T. neglectus defecated during their daily sojourn up to a few meters behind the entrance, contributing nothing to the troglobiotic com munity. On the other hand, before overwintering, both Troglophilus defecated; their scarce fecal pellets (about two per 100 cave crickets) were occasionally found deeper inside caves at the overwintering sites, and were consumed by troglobiotic Collembola, Diplopoda and leiodid Coleoptera. ese also attracted Pseudoscorpi ones, Diplura and carabid Coleoptera. ese pellets con tributed to the troglobiotic community once a year, but their total mass in all the caves was estimated at less than 200 mg/year. T ONE NOVAK, FRANC JANEKOVI & SA KA LIPOV EK

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ACTA CARSOLOGICA 42/2-3 2013 305 Tab. 3: Individual dry mass in the 11 species under investigation. t-test between the series at the beginning and at the end of overwinter ing. Signcant p-values are in bold. Species N of individuals; beginningend of overwintering Beginning MeanSD Min-Max [mg] End MeanSD Min-Max [mg] Average dierence beginningend of over-wintering [%] t; p df F. illyrica 5 241.7.7 218.6.2 298.1.2 276.3 27.0 23.33 3.58 0.007 8 C. lefeburiana 5 328.7.1 278.4.3 218.2.0 172.9.9 .62 4.53 0.001 9 A. aurantiacus 5 6.5.7 5.3.2 4.5.4 4.1.1 .77 5.18 0.001 8 A. aurantiacus 5 8.9.3 7.3.9 6.0.9 5.1.3 .58 4.11 0.003 8 G. annulatus 10 7.91.32 5.6.2 7.42.02 6.2.1 .19 0.58 0.669 17 G. titanus 21 24 12.6.8 9.7.8 13.1.3 7.8.6 3.97 0.07 0.943 43 M. menardi 10 109.7.8 54.8.6 109.3.6 49.5.7 .36 0.04 0.965 20 L. schreibersii 5 24.8.6 16.5.0 21.3.6 13.1.3 .11 0.75 0.466 13 T. dubitata 10 26.4.7 21.0.7 22.3.4 16.3.2 .53 0.37 0.028 19 S. libatrix 10 144.1.2 107.3.3 121.1.3 74.5.6 .96 2.21 0.041 18 T. cavicola 6 165.7.6 122.2.7 135.4.9 125.8.6 .29 2.16 0.059 9 T. cavicola 5 172.2.5 154.9.7 148.6.6 129.8.4 .7 2.58 0.032 8 T. cavicola l. 10 31.4.9 26.9.5 42.0.7 28.6.4 33.76 4.65 < 0.001 18 T. neglectus 4 143.6.4 101.7.2 112.8.0 73.2.4 .45 0.93 0.407 4 T. neglectus 5 201.2.1 174.7 90.6.2 21.0.5 .97 2.99 0.015 9 T. neglectus l. 13 45.6.0 15.1 36.8.5 17.8.0 .3 0.91 0.371 25 Tab. 4: Estimated average individual energy budgets for the 11 species under investigation at the beginning and at the end of overwin tering. *According to Novak & Kutor (1982), 1 estimated mean budget for arthropods, 2 provisional budget. Species E/g dry mass [J/g]* Beginning [J/ indiv.] End [J/indiv.] Average dierence beginningend of over-wintering [%] F. illyrica 20000 2 4752.0 5962.0 25.46 C. lefeburiana 20000 2 6573.2 4364.0 .61 A. aurantiacus 25000 1 193.0 132.3 .45 G. annulatus 25000 1 257.5 173.6 .58 G. titanus 25000 1 315.4 328.2 4.06 M. menardi 26449 2901.5 2890.4 .38 L. schreibersii 23483 530.1 550.5 3.85 T. dubitata 25554 675.1 564.1 .44 S. libatrix 27127 3909.8 3284.5 .99 T. cavicola 24478 4054.8 3314.3 .26 T. cavicola 24579 4233.0 3651.5 .74 T. cavicola larvae 23495 707.7 1006.1 42.16 T. neglectus 21348 3065.6 2408.1 .45 T. neglectus 22967 4621.0 2081.2 .96 T. neglectus larvae 21118 962.2 776.4 .31 C ONTRIBUTION OF NON TROGLOBIOTIC TERRESTRIAL INVERTEBRATES TO CARBON INPUT IN HYPOGEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 306 Tab. 5: Mean abundance of individuals (N), total fresh and dry mass at the beginning of overwintering, and method of overwintering in the 11 invertebrate species: a active hypogean ecophase, d diapause, q quiescence. Species N of indiv. in 54 caves Method Total fresh mass Total dry mass N of dead indiv. Dry necromass [g] Beginning [g] End [g] Dierence [%] Beginning [g] End [g] Dierence [%] F. illyrica 333 q 647.45 556.94 .98 79.12 99.27 25.47 0 0.00 C. lefeburiana 12 q 26.43 18.99 .15 3.94 2.62 .50 0 0.00 A. aurantiacus 8406 d 147.19 161.14 9.48 64.89 44.47 .47 430 3.31 G. annulatus 86 q 2.57 2.64 2.72 0.89 0.60 .58 0 0.00 G. titanus 109 q 4.68 4.88 4.27 1.38 1.43 3.62 10 0.08 M. menardi 772 a 191.60 224.34 17.09 84.69 79.66 .94 10 1.10 L. schreibersii 147 d 8.42 10.31 22.45 3.32 3.45 3.92 1 0.02 T. dubitata 210 d 12.35 10.58 .33 5.55 4.64 .40 20 0.53 S. libatrix 309 q 80.69 70.71 .37 44.54 37.41 .01 40 5.76 T. cavicola 874 d 417.57 520.96 24.76 144.78 118.34 .26 0 0.00 T. cavicola 1102 d 596.42 674.12 13.03 189.79 163.71 .74 0 0.00 T. cavicola larvae 3772 q 328.43 563.20 71.48 113.61 161.52 42.17 3 0.09 T. cavicola 5748 1342.42 1758.28 30.98 448.18 443.57 .03 0.09 T. neglectus 377 q 190.28 166.82 .33 54.14 42.53 .44 1 0.14 T. neglectus 476 q 309.05 166.99 .97 95.77 43.13 .97 0 0.00 T. neglectus larvae 1373 q 182.98 204.64 11.84 62.56 50.48 .31 2 0.09 T. neglectus 2226 682.31 538.45 .08 212.47 136.14 .93 0.23 Total 3146.11 3357.26 6.71 948.97 853.26 .09 11.18 Tab. 6: e % contribution of species among the 11 species to total number of species, fresh and dry mass, carbon and energy in caves in winter. N Fresh mass Dry massCarbon Energy A. aurantiacus 45.79 T. cavicola 42.65 T. cavicola 47.16 T. cavicola 48.15 T. cavicola 31.31 T. neglectus 21.68 T. neglectus 22.36 T. neglectus 20.70 T. neglectus 12.13 F. illyrica 20.61 M. menardi 8.91 M. menardi 9.92 M. menardi 4.21 M. menardi 6.09 F. illyrica 8.47 A. aurantiacus 7.18 F. illyrica 1.81 A. aurantiacus 4.68 A. aurantiacus 6.83 F. illyrica 7.13 S. libatrix 1.68 S. libatrix 2.56 S. libatrix 4.69 S. libatrix 5.35 T. dubitata 1.14 C. lefeburiana 0.84 T. dubitata 0.58 T. dubitata 0.63 L. schreibersii 0.80 T. dubitata 0.39 C. lefeburiana 0.41 C. lefeburiana 0.35 G. titanus 0.59 L. schreibersii 0.27 L. schreibersii 0.35 L. schreibersii 0.34 G. annulatus 0.47 G. titanus 0.15 G. titanus 0.15 G. titanus 0.15 C. lefeburiana 0.07 G. annulatus 0.09 G. annulatus 0.09 G. annulatus 0.10 DISCUSSION Terrestrial fauna migrating between surface and hypo gean habitatsi. e., the non-troglobiotic taxahave been considered among the most important vectors of carbon supplying subterranean communities (Culver & Pipan 2009; Fong 2011). In contrast to this general logical ap pearance, this statement should be carefully examined in every karstic region settled by distinctive fauna, as we demonstrate in this contribution for Central European caves. Despite ambiguous data in some species and stag es under study, which deserve renewed measurement, we can identify the general trends for non-troglobiotic animals from caves in Slovenia during winter. e total biomass of non-troglobiotic taxa increases slightly dur ing winter, while the dry mass decreases, indicating that utilization of reserve lipids is the most important meta bolic process in these species overwintering in central European caves (Lipovek et al. 2004, 2008, 2009, 2011; Novak et al. 2004). e more frequent species and those with grater biomass share the rst ve places in a rank ing of their importance to subterranean habitats in either abundance, fresh and dry mass or energy bulk, compris T ONE NOVAK, FRANC JANEKOVI & SA KA LIPOV EK

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ACTA CARSOLOGICA 42/2-3 2013 307 CONCLUSIONS Terrestrial invertebrates overwintering in caves have hy pothetically been considered an important vector convey ing carbon from the epigean to the hypogean habitat, but in reality they represent a rather limited food resource for hypogean species. In Central European caves, ve among the 11 species contribute 93.08.71% ratios of the total abundance, biomass and dry mass, energy and incorpo rated carbon of the 11 species. ese ve species migrate between the epigean and hypogean habitats and are the most important ones in this respect. Despite this, they contribute a negligible amount of carbon to subterranean fauna. ing 93.08.71% ratios among the 11 species. ese are T. cavicola, T. neglectus, M. menardi, F. illyrica and A. au rantiacus W hile M. menardi (mostly, see the life cycle in Novak et al. 2010) and L. schreibersii are residents of subterranean habitats, the migratory species alternate between epigean and hypogean habitats, with minimal total loss (0.74%) of their biomass in caves during winter. Besides, all the individuals that die in caves are infected by entomopathogenic fungi and are thus not directly available to subterranean animals, at least not to arthro pods, which strictly avoid contact with infested cadavers. Predated individuals (Novak et al. 2010) and the scanty excrement produced during pre-overwintering defeca tion represent the only resource of non-troglobiotic taxa in Slovenian caves. e overall amount of these contri butions does not exceed 1% of their biomass, i. e., about 0.3% dry mass and 0.15% organic carbon. In contrast to the general assumption and our hypothesis, in Central European caves, non-troglobiotic taxa contribute very little to the deep troglobiotic species. Additionally, pre dated individuals contributed maximally 0.15% C. is low carbon input might be the main reason why a major group of troglobionts inhabit shallow subterranean habiinhabit shallow subterranean habi tats (Novak et al. 2012). Many points must be considered in providing a credible estimate of each species contribution to bio mass, dry mass, necromass, energy and/or carbon input in hypogean habitats: its life cycle and type of dormancy, preferred hypogean habitats, the biomass bulk of immi grating and emigrating individuals, the duration of its hypogean ecophase, the strategy for creation and utiliza tion of energy-supporting reserve stocks, the time and space frames of specic investigations and the form of species contribution (prey, necromass or excrement) in the subterranean habitat. In our study, species abun dance was underestimated by at least 30% of the visible population (Novak, unpublished). W hile more accurate data of this kind would improve our quantitative esti mates of mass, energy and carbon ow, they would not change the outcomes on the importance of migratory non-troglobionts to troglobiotic communities. Our results represent actual conditions in a specic period and region. ese conditions change from year to year and vary among dierent regions and caves. In studies on the biomass, energy and/or carbon ow sup plied by non-troglobiotic animals to caves it, is of spe cial importance to check whether these individuals or their products, like eggs, are predated, or if they defecate there, or whether their cadavers are consumed by species which are more or less permanently resident in caves. For analytical purposes, in the context of carbon ux, large invertebrates can be compared with coarse particulate organic carbon, which moves relatively short distances before it is broken down or consumed (Simon & Beneld 2001). Such a distribution is also true for all the species under study, which are most abundant in the upper 10 m beneath the surface (Novak et al. 2012). Gibert (1986) found that dissolved organic carbon is the most important among abiotic carbon sources in subter ranean water. Analogously, it seems justied to consider the possibility that microfauna are the most signicant biotic vector of carbon supply to terrestrial hypogean habitats. ACKNO WLEDGEMENTS W e are indebted to David Culver and two anonymous ref erees for their insightful comments and corrections of the initial version of the manuscript, and to Michelle Gadpaille for valuable improvements to the language. is study was partly supported by the Slovene Research Agency within the research program Biodiversity (grant P1-0078). C ONTRIBUTION OF NON TROGLOBIOTIC TERRESTRIAL INVERTEBRATES TO CARBON INPUT IN HYPOGEAN HABITATS

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ACTA CARSOLOGICA 42/2-3 2013 308 Culver, D.C. & T. Pipan 2009: Biology of Caves and Other Subterranean H abitats.Oxford University Press, Oxford, U.K., pp. 254. Deharveng, L. & A. Bedos, 2000: e cave fauna of southeast Asia. Origin, evolution and ecology. In: W ilkens, H., D.C. Culver & W .F. Humpreys (Eds.): Subterranean ecosystems. Elsevier, pp. 603. Denlinger, D.L., 2002: Regulation of diapause.Annual Review of Entomology 47, 93. Engel. A.S., 2005: Observations on the biodiversity of suldic karst habitats.Journal of Cave and Karst Studies, 69, 187. Fong, D.W ., 2011: Management of subterranean fauna in karst.In: van Beynen, P.E. (Ed.), Karst Manage ment. Springer, pp. 201. Gibert, J., 1986: cologie dun systme karstique juras sien. Hydrogologie, drive animale, transits de matires, dynamique de la population de Niphargus (Crustace Amphipode).Mmoires de Biospolo gie, 13, 1. Gnaspini, P., 2005: Guano community. In: Encyclopedia of Caves W hite, W .B. &. D.C. Culver (Eds.). Spring er, pp. 201. Houghton, R.A., F. Hall & S.J. Goetz, 2009: Impor tance of biomass in the global carbon cycle.Journal of Geophysics Research 114, G00E03, doi:10.1029/2009JG000935. Jequier, J.-P., 1964: tude cologique et statistique de la faune terrestre dune caverne du Jura Suisse au cours dune anne dobservations.Revue Suisse de Zoologie, 71, 13. Lavoie, K.H., K.L. Helf & T.L. Poulson, 2007: e biology and ecology of North American cave crickets.Jo urnal of Cave and Karst Studies, 69, 114. Lipovek Delakorda, S., I. Letofsky-Pabst, T. Novak, M. Giovanelli, F. Hofer & M.A. Pabst, 2008: Applica tion of elemental microanalysis to elucidate the role of spherites in the digestive gland of the helicid snail Chilostoma lefeburiana.Journal of Microscopy, 231, 38. DOI: 10.1111/j.1365-2818.2008.02015.x Lipovek Delakorda, S., T. Novak, F. Janekovi, L. Seni & M.A. Pabst, 2004: A contribution to the function al morphology of the midgut gland in phalangiid harvestmen Gyas annulatus and Gyas titanus dur ing their life cycle.Tissue & Cell, 36, 275. doi:10.1016/j.tice.2004.04.003 Lipovek, S., I. Letofsky-Pabst, T. Novak, F. Hofer & M.A. Pabst, 2009: Structure of the Malpighian tu bule cells and annual changes in the structure and chemical composition of their spherites in the cave cricket Troglophilus neglectus Krauss, 1878 (Rhaphidophoridae, Saltatoria).Arthropod Struc ture and Development 38, 315. doi:10.1016/j. asd.2009.02.001 Lipovek, S., T. Novak, F. Janekovi & M.A. Pabst, 2011: Role of the fat body in the cave crickets Troglophi lus cavicola and Troglophilus neglectus (Rhaphido phoridae, Saltatoria) during overwintering.Ar thropod Structure and Development, 40, 54. doi:10.1016/j.asd.2010.09.002 Novak, T., 2005: Terrestrial fauna from cavities in North ern and Central Slovenia, and a review of systemati cally ecologically investigated cavities.Acta Carso logica, 34, 169. Novak, T. & V. Kutor, 1982: Contribution a la connais sance de la biomasse et du bilan nergetique de la faune des entres de grotte en Slovenie (Yougosla vie).Mmoires de Biospologie, 8, 82. Novak, T. & V. Kutor, 1983: On Troglophilus (Rhaphi dophoridae, Saltatoria) from North Slovenia (YU). Mmoires de Biospologie, 10, 127. Novak, T., S. Lipovek Delakorda, L. Seni, M.A. Pabst & F. Janekovi, 2004: Adaptations in phalangiid harvestmen Gyas annulatus and G. titanus to their preferred water current adjacent habitats. Acta Oecologica, 26, 45. Novak, T., M. Perc, S. Lipovek & F. Janekovi, 2012: Duality of terrestrial subterranean fauna. Interna tional Journal of Speleology, 41, 181. Novak, T., T. Tkavc, M. Kuntner, A.E. Arnett, S. Lipovek Delakorda, M. Perc & F. Janekovi, 2010: Niche partitioning in orbweaving spiders Meta menardi and Metellina merianae (Tetragnathidae).Acta Oecologica, 36, 522. Perruchoud, D., L. W althert, S. Zimmermann & P. Lscher, 2000: Contemporary carbon stocks of mineral forest soils in the Swiss Alps.Biogeochem istry, 50, 111. Por, F.D., 2007: Ophel: a groundwater biome based on chemoautotrophic resources. e global signi cance of the Ayyalon cave nds, Israel.Hydrobio logia, 592, 1. Riebeek, H, 2011: e carbon cycle. http://earthobser vatory.nasa.gov/Features/CarbonCycle/ REFERENCES T ONE NOVAK, FRANC JANEKOVI & SA KA LIPOV EK

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ACTA CARSOLOGICA 42/2-3 2013 309 Salonen, K., J. Sarvala, I. Hakala & M.-L. Viljanen, 1976: e relation of energy and organic carbon in aqua tic invertebrates.Limnology and Oceanography, 21, 724. Srbu, S.M., T.C. Kane & B.K. Kinkle, 1996: A chemoau totrophically based cave ecosystem.Science, 272, 1953. Schmidt, M.W .I., M.S.Torn, S. Abiven, T. Dittmar, G. Guggenberger, I.A. Janssens, M. Kleber, I. KgelKnabner, J. Lehmann, D.A.C. Manning, P. Nannip ieri, D.P. Rasse, S. W einer & S.E. Trumbore, 2011: Persistence of soil organic matter as an ecosystem property.Nature, 478, 49 Simon, K.S. & E.F. Beneld, 2001: Leaf and wood break down in cave streams.Journal of the North Ameri can Benthological Society, 482, 31. Simon, K.S. & E.F. Beneld, 2002: Ammonium retention and whole stream metabolism in cave streams.Hy drobiologia, 482, 31. Simon, K.S., T. Pipan & D.C. Culver, 2007: A conceptual model of the ow and distribution of organic car bon in caves.Journal of Cave and Karst Studies, 69, 279. Taylor, S.J., J. Krejca & M.L. Denight, 2005: Foraging and range habitat use of Ceuthophilus secretus (Or thoptera Rhaphidophoridae), a key trogloxene in central Texas cave communities.American Mid land Naturalist, 154, 97. Tercafs, R.R., 2001: e protection of the subterranean environment. Conservation principles and manage ment tools.P.S. Publishers, Luxembourg, 400 p. U.S.DOE, 2008: Carbon cycling and biosequestration: Integrating Biology and climate through system science.Report from the March 2008 W orkshop, Doe/SC, U.S. Deptartment of Energy Oce of Science, 133 p. C ONTRIBUTION OF NON TROGLOBIOTIC TERRESTRIAL INVERTEBRATES TO CARBON INPUT IN HYPOGEAN HABITATS



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USING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN SEMI ARID SYSTEMS E D W ARDS PLATEAU T E X AS USA UPORABA HIDROGEOKEMI NIH IN EKOHIDROLO KIH ODZIVOV ZA RAZUMEVANJE EPIKRA KIH PROCESOV V POLSU NIH SISTEMIH PLANOTA E D W ARDS T EKSAS ZDA Benjamin F. SCHW ARTZ 1* Susanne SCHWINNING 1 Brett GERARD 1 Kelly R. KUKO W SKI 1 Chasity L. STINSON 1 & Heather C. DAMMEYER 1 Izvleek UDK 556.32:581.11(736.4) Benjamin F. Schwartz, Susanne Schwinning, Brett Gerard, Kelly R. Kukowski, Chasity L. Stinson & Heather C. Dam meyer: Uporaba hidrogeokeminih in ekohidrolokih odzivov za razumevanje epikrakih procesov v polsunih sistemih, planota Edwards, Teksas, ZDA Epikras je prepustna meja med povrinskim in podzemeljskim okoljem in se lahko pojmuje kot vadozno kritino obmoje epigenih krakih sistemov, ki se pod netopnim pokrovom ne razvijajo. Iz hidrolokega vidika je to obmoje pogosto pojmo vano kot prepustno le v eno smer (navzdol), vendar povezave med vodnimi potmi skozi epikras in koreninskimi sistemi le snatih rastlin pomenijo, da se voda premika navzgor in navzdol po epikrasu. Vendar pa je dinamika teh tokov zapletena in zelo odvisna od spremenljivosti v prostorski strukturi epikrasa, znailnosti vegetacije, kot tudi v asovni spremenljivosti pada vin in izhlapevanja. Tukaj povzemamo spoznanja iz raziska v na ve mestih na planoti Edwards v osrednjem Teksasu, ki zdruuje izotopske, hidrogeokemine in ekozioloke metodologije. 1) Gosta gozdna vegetacija na golih mestih ali mestih s tanko prstjo (0-30 cm) deloma rpa vodo iz epikrasa. 2) Vendar pa je drevesna transpiracija obiajno prekinjena v asu suhih poletij, kar kae, da se rastlinam dostopen del uskladiene vode v epikrasu izrpava hitro, tudi ko trajno kapljanje v jamah e nakazuje, da je voda e vedno prisotna v epikrasu. 3) Vodne poti, ki napajajo kapljanje v jamah so hitro prekinjene in iz ven dosega rastlinskih korenin. 4) Globoka inltracija in pol njenje teh sistemov ni efektivno brez monih in stalnih pada vin, ki presegajo doloeno mejno vrednost. Pragovi so mono povezani s predhodno potencialno evapotranspiracijo in pa davinami, kar kae na odvisnost od stopnje vlage v obmoju epikrake cone evapotranspiracije. Epikras in nezasiena cona sta v tej regiji lahko pojmovani kot spremenljivo nasien sistem s skladienjem v razpokah, poroznem mediju in v plitvih viseih vodonosnikih. Veina teh je drevesnemu koreninskemu 1 Texas State University, Department of Biology, 601 University Drive, San Marcos, T X 78666 *corresponding author: e-mail: bs37@txstate.edu Received/Prejeto: 23.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 315, POSTOJNA 2013 Abstract UDC 556.32:581.11(736.4) Benjamin F. Schwartz, Susanne Schwinning, Brett Gerard, Kelly R. Kukowski, Chasity L. Stinson & Heather C. Dam meyer: Using hydrogeochemical and ecohydrologic responses to understand epikarst process in semi-arid systems, Edwards plateau, Texas, USA e epikarst is a permeable boundary between surface and subsurface environments and can be conceptualized as the va dose critical zone of epigenic karst systems which have not de veloped under insoluble cover. From a hydrologic perspective, this boundary is oen thought of as being permeable in one direction only (down), but connectivity between the ow paths of water through the epikarst and the root systems of woody plants means that water moves both up and down across the epikarst. However, the dynamics of these ows are complex and highly dependent on variability in the spatial structure of the epikarst, vegetation characteristics, as well as temporal variability in precipitation and evaporative demand. Here we summarize insights gained from working at several sites on the Edwards Plateau of Central Texas, combining isotopic, hydro geochemical, and ecophysiological methodologies. 1) Dense woodland vegetation at sites with thin to absent soils (0-30 cm) is in part supported by water uptake from the epikarst. 2) How ever, tree transpiration typically becomes water-limited in dry summers, suggesting that the plant-available fraction of stored water in the epikarst depletes quickly, even when sustained cave drip rates indicate that water is still present in the epikarst. 3) Flow paths for water that feeds cave drips become rapidly disconnected from the evaporation zone of the epikarst and out of reach for plant roots. 4) Deep inltration and recharge does not occur in these systems without heavy or continuous precipitation that exceeds some threshold value. resholds are strongly correlated with antecedent potential evapotranspira tion and rainfall, suggesting control by the moisture status of the epikarst evapotranspiration zone. e epikarst and unsatu

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ACTA CARSOLOGICA 42/2-3 2013 316 rated zone in this region can be conceptualized as a variably saturated system with storage in fractures, matrix porosity, and in shallow perched aquifers, most of which is inaccessible to the root systems of trees, although woody vegetation may con trol recharge thresholds. Keywords : hydraulic disconnection, precipitation thresholds, root zone, plant water use, recharge, epikarst storage, baromet ric pressure. sistemu nedostopna, eprav lahko lesna vegetacija nadzira me jne vrednosti polnjenja. Kljune besede : hidravlina prekinitev, pragovi padavin, kore ninska cona, rastlinska uporaba vode, polnjenje, skladienje v epikrasu, zrani tlak. I NTRODUCTION e Critical Zone is the thin veneer of Earth that extends from the top of the vegetation to the base of weathered bedrock where fresh water ows, soils are formed from rocks, and terrestrial life ourishes (NSF 2012). As an in tegrated system, the Critical Zone is poorly understood in general, but especially so in systems where the transi tion from soil to soluble bedrock materials is blurred by the dominance of a wide hybrid zone, part soil and part fractured, weathered, and porous bedrock. ese transi tion zones are commonly found in regions where karst is exposed at the surface and are known as the epikarst (Jones et al. 2003). Unlike horizons of weathered non-soluble bedrock beneath soils, the epikarst is a complex network of rock matrix, empty cavities, soil pockets and ow paths of variable hydraulic conductivity (Klimchouk 2004; Es trada-Medina et al. 2013). Importantly, the epikarst also contains plant roots. us, the epikarst not only facili tates the transport of surface water and chemical constit uents to underlying aquifers, it also stores water, trans forms organic and inorganic materials, and makes water and nutrients available for uptake by plants. In short, the epikarst is a highly dynamic boundary between surface and groundwater environments that modies, and is modied by, the transfer of water. e hydrologic function of the epikarst has been studied using a variety of methods ranging from basinwide studies (e.g., Mangin 1975; W illiams 1983) to incave studies of drips and inlets. Studies at both scales have concluded that the epikarst is an important stor age component in karst systems and that the hydraulic properties of the epikarst are substantially dierent from those of the underlying conduit-dominated ow and transport system (e.g., Smart & Friederich 1986; Lee & Krothe 2001; Arbel et al. 2010). Comparatively, the epikarst has higher storage, variably saturated condi tions, higher fracture density, and relative homogeneity of hydraulic ow pathways. One aspect that has received relatively little atten tion is the interaction between epikarst and vegetation (Schwinning 2010). Trees and shrubs take up water and nutrients from the epikarst and by doing so, inu ence the dynamics and chemical composition of water draining out of the epikarst. e few existing studies suggest that root access to water stored in the epikarst is extremely heterogeneous; probably constrained by the sizes and distribution of conduits and the durabil ity of weathered rock. Studies from karst areas in sea sonally dry climates of Mexico, China, and the USA suggest limited use of epikarst water by trees, chiey to supplement water during the dry season, but not enough to completely alleviate water stress (Q uerejeta et al. 2006; Q uerejeta et al. 2007; Schwinning 2008; Nie et al. 2011; Rong et al. 2011; Deng et al. 2012; Kukowski et al. 2013). In addition, these studies showed that root ing depth and depth of water uptake vary by species and with the age of individuals. W ith few exceptions, the stable isotope ratio analysis of water extracted from woody plants suggested relatively shallow, evaporative ly enriched water sources in the epikarst. At an epikarst site on the Edwards Plateau of Central Texas, USA, variation in precipitation-adjusted annual evaporation was more strongly tied to soil depth than to tree den sity (Heilman et al. 2009). So far, only one study in the Balcones Fault Zone of the Edwards Plateau of Central Texas, USA, found roots tapping directly into deep con duits in the phreatic zone at 25 30 m depth (Jackson et al. 1999), and this situation seems to be the exception rather than the rule. In this paper, we summarize ndings from three sites in central Texas, USA, where we combined meth ods from hydrogeology, geochemistry, ecohydrology and statistical modeling to characterize the epikarst as a complex, dynamic storage system, interacting with sur face vegetation. Our overall goal was to determine the degree of hydrological coupling between dierent com ponents of the Critical Zone, including woody vegeta tion at the surface, a thin layer of soil, the epikarst, and rates of epikarst drainage as estimated from speleothem drip rates inside caves. B. F. S CH W ARTZ S. S CH W INNING B. G ERARD K. R. K UKO W SKI C. L. STINSON & H. C. D AMMEYER

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ACTA CARSOLOGICA 42/2-3 2013 317 S ITE D ESCRIPTIONS Our study sites are located in the Hill Country of Central Texas, USA. e rst, McCarty Cave, is just south of the city limits of San Marcos, T X. is shallow cave lies 3 10 m beneath a forested area dominated by cedar elm (Ulmus crassifolia), escarpment live oak (Quer cus fusiformis), and Ashe juniper ( J uniperus ashei), with fewer numbers of Texas persimmon (Diospyros texana), netleaf hackberry (Celtis laevigata var. reticulate), and agarita (Berberis trifoliolata) Soils are shallow and rocky, ranging from 0 to 30 cm, and overlie epikarst formed in the Edwards limestone within the recharge zone of the Edwards aquifer, a conned aquifer that supplies fresh water to 2 million people between Austin and San An tonio along the Balcones fault zone. A domestic well was available for sampling 50 m from the cave. e second site is Cave W ithout A Name (CW AN), approximately 17 km northeast of Boerne, T X. is commercial cave lies 15 to 30 m below the surface, and is formed in the lower Glen Rose Limestone, which com prises the upper part of the Middle Trinity group in the extensive Trinity Aquifer system underlying much of the Central Texas Hill Country (Veni 1994; Mace et al. 2000). e Trinity Aquifer is important not only for groundand surface water resources in the Hill Country, but it also contributes substantially to the Edwards Aquifer via cross-formational ow along the Balcones Fault Zone (Kuniansky and Holligan 1994). Vegetation and soil type and thickness were similar to McCarty Cave, but the site had greater tree diversity, including two more species of oak, both deciduous: Texas oak (Quercus texana) and shin oak (Quercus sinuata). CW AN also contains a hy drologically active branchwork stream system contain ing 4.5 km of nearly water-lled main stem and tributary conduits. Exploration continues in the upstream reaches of the system. e third site is Headquarters Cave on Camp Bullis, north of San Antonio. is cave lies beneath a hillside, and is formed across the contact zone between the Ed wards and underlying Upper Glen Rose Limestones. e overstory vegetation is composed chiey of escarpment live oak and Ashe juniper, and understory vegetation, soil characteristics and thickness are similar to the other two sites. However, just prior to this study, nearly all ju niper trees were removed from a ca. 2,800 m 2 area over lying and surrounding the cave. e region is classied as having semi-arid to subtropical climate. Mean annual precipitation (1961-1990) in San Marcos, T X is ~864 mm + 100mm ( http://www. wrcc.dri.edu/precip.htm l ), which is similar to the Boerne and San Antonio areas. Mean August maximum tem perature is 35 C and mean January minimum is 4 C (Dixon 2000). Our primary research questions were: 1) where are the water sources for forest trees located in the soil to bedrock continuum and, consequently, how does tree water use likely aect deep inltration and recharge in the Edwards Plateau, 2) how do environmental param eters control inltration and recharge, and 3) can hydro geochemical data recorded inside caves be used to rene conceptual models of epikarst form and function? To an swer these questions, we used a variety of measurements, including liquid water stable isotopes and sap-ow to determine the source of tree water, and cumulative po tential evapotranspiration and precipitation data to de termine precipitation thresholds required for groundwa ter recharge. Finally, through comparing three sites with slightly dierent geologic and hydrological characteris tics, two in Edwards limestone, one in Lower Glen Rose Limestone, we also aimed to examine the inuence of geologic factors on these hydrologic processes. METHODS At each site, a suite of instruments was installed and data collected continuously at 10 or 15 minute intervals. Surface instrumentation included a full weather station (precipitation, wind speed and direction, solar radiation, temperature, and relative humidity; Onset Computer Corp., Bourne, MA, USA), two pairs of EC-5 soil mois ture sensors (Decagon Devices, Inc., Pullman, W A, USA) at 15 and 25 cm depths, and sap ow sensors on all major tree species (one sensor per tree; six trees per species). Inside the caves, conduits and speleothem drip sites were instrumented to measure drip rate or discharge, specic conductivity and temperature of drip water, and U SING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN ...

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ACTA CARSOLOGICA 42/2-3 2013 318 air temperature, which were reported to Hobo datalog gers (Micro Station H-21-002, Onset Computer Corp., Bourne, MA, USA) or recorded on Schlumberger CTD Diver instruments. At C W AN, 6 sites were monitored, including dripping speleothems formed on porous ceil ings (2), dripping speleothems connected to small frac tures (2), a small conduit that discharged water only af ter heavy precipitation and overland ow, and the cave stream. Drip water was collected on tarps and funneled into tipping-bucket rain gauges. Conduit and stream discharges were measured at weirs using paired pres sure transducers and barometric pressure loggers, and stage-discharge relationships that allow measured wa ter depth (stage) to be converted to discharge in units of L/s. At McCarty Cave, the well and two adjacent drip sites were monitored. W aters dripped directly from two sites on the porous bedrock ceiling with surface areas of 3 and 1 m 2 respectively. In Headquarters Cave, one dripping speleothem was monitored at the bottom of the cave. At all sites, water samples were periodically collect ed (hourly to monthly, depending on site and hydrologic conditions) for analysis of major ions, dissolved nutri ents, and liquid water stable isotopes. Precipitation was also collected at all caves. Granier type sap ow sensors (Granier 1985) were inserted 1 cm deep into the sapwood approximately at breast height and the trunk wrapped in reective bubble insulation to reduce temperature artifacts. is method measures the velocity of sap ow rise in trees during the day, which is considered a relative measure of wholetree transpiration rate. On the same trees, pre-dawn leaf water potentials were periodically measured using a Scholander pressure chamber (Model 1000, PMS In struments, Albany, OR, USA) and stem samples taken aer dawn for the stable isotope ratio analyses of stem water. Stem samples were stored in screw cap vials and stored frozen before being cryogenically extracted un der vacuum pressure (Ehleringer et al. 2000). Extracted stem, precipitation, and cave water samples were ana lyzed on an LGR DLT-100 Liquid W ater Stable Isotope Analyzer (Los Gatos Research, Inc., Mountain View, CA, USA) with internal standards created using the LGR cer tied standards. Values are expressed in standard delta notation relative to the V-SMO W standard (Gonantini 1978). Precision was estimated at 0.5 per mil for tD and 0.3 per mil for t 18 O (one standard deviation). Statistical analyses were performed using JMP and R soware packages (JMP 1989-2007; R Development Core Team 2012). RESULTS AND D ISCUSSION T REE W ATER SOURCES AND TRANSPIRATION RATE W e applied two methods to deduce the likely water sourc es of trees: W e measured stable isotope ratios of extracted stem water and compared them with precipitation and drip water samples to identify (or rule out) water sources. is technique relies on the fact that plant water uptake by roots and pressure-driven ow through the sapwood does not involve isotopic fractionation (Ehleringer et al. 2000). erefore, stem water, if sampled far upstream from evaporation sites (e.g., leaves), is considered a vol ume-weighted linear mixture of all plant water sources. Additionally, we measured the seasonal dynamics of sap water rise through trees and compared it to the dynamics of soil moisture and epikarst drainage rates to determine whether tree transpiration was more closely linked with the pulse dynamics of the soil and shallow epikarst, or the dampened dynamics of deeper water stores. Comparison of isotope ratios ruled out cave drip, stream, or well water as a possible source of water for trees. During the summer drought of 2009, neither at the McCarty Cave site nor at Camp Bullis were drip water sources a likely source; instead, extracted sap water in dicated evaporatively enriched water sources, implying that soil and the top of the epikarst were the predomi nant water sources. At McCarty Cave, sap water became progressively more enriched during the summer drought and farther removed from the local meteoric water line (Fig. 1). is suggests that the origin of the transpiration ux was precipitation stored relatively close to the sur face, subjected to direct evaporation, and diminishing over the course of summer. At Camp Bullis, evaporative enrichment was not as extreme, indicating that trees had access to more stable, deeper water sources (Fig. 1). Ac cordingly, minimal predawn water potential values mea sured in August for Ashe juniper and live oak were lower at the McCarty Cave site than at Camp Bullis (live oak: -4.2 MPa 0.4 se (McCarty) vs. -3.1 MPa 0.3 se (C. Bullis); Ashe juniper: -8.03 MPa 0.08 se (McCarty) vs. -7.1 MPa 0.2 se (C. Bullis)). Aer the rst substantial rainfall events in Sep tember 2009 (70 mm), stem water isotope ratios at the McCarty Cave site moved back towards the Local Mete oric W ater Line (LMWL) and precipitation values, and B. F. S CH W ARTZ S. S CH W INNING B. G ERARD K. R. K UKO W SKI C. L. STINSON & H. C. D AMMEYER

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ACTA CARSOLOGICA 42/2-3 2013 319 plant water potentials recovered immediately (data not shown), indicating once again that trees responded to soil and shallow epikarst water. Overall, deeper epikarst water could be ruled out as a likely source for transpi ration, since mean drip isotope values, presumed to be a sample of water stored in the lower epikarst, did not correspond to tree water isotope ratios at any time, not even immediately aer a rain event. is, combined with the fact that the isotope ratios of stem waters were highly variable according to meteorological conditions, while those of the drips were not, underscores our conclusion that the water available to the trees is largely separate from the (presumably deeper) sources supplying water to the drips. Sap ow velocity (indicative of tree transpiration) declined rapidly during the month of May 2009 at both sites. However, rates of decline were dierent for dier ent species, and there were also dierences between sites (Fig. 2). At the McCarty Cave site, live oak transpiration declined most rapidly during the month of June, while Ashe juniper transpiration declined gradually over the entire summer. By contrast, at Camp Bullis, transpiration of Ashe juniper declined more rapidly than of live oak. is may have been because live oak was deeper-rooted than juniper at this site, with greater access to relatively more stable water sources at depth. is interpretation is also consistent with the stable isotope and water po tential data. Escarpment live oak can grow deeper roots than Ashe juniper, if given a chance (Jackson et al. 1999). However, at McCarty Cave, there may not have been a chance, due to the less fractured rock substrate (Kukows ki et al. 2013). W here some parts of the cave are just 4 m below ground, we saw no tree roots protruding from the ceiling, suggesting that the maximum rooting depth was 4 m at most, and probably much less. A trenching study in the same region showed that, irrespective of soil depth, most tree roots are no deeper than 40 cm, al though a few individual roots were observed to penetrate Fig. 1: Monthly liquid water stable isotope data from stem water extractions for three tree species at McCarty Cave (le panel) and two species at H eadquarters Cave (right panel) showing evaporative enrichment over the summer drought of 2009. Stem water isotopes return to near starting values aer ~70 mm of precipitation in early September, 2009. U SING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN ...

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ACTA CARSOLOGICA 42/2-3 2013 320 Fig. 2: Relationships between sap ow rates, soil moisture at 15 cm (solid line) and 25 cm (dashed line), precipitation, and drip rates at in-cave sites, for McCarty Cave and H eadquarters Cave during the 2009 drought. Soil data were lost for a period of 1 month at Mc Carty Cave. deeper through rock ssures (Heilman et al. 2009). us, it appears that live oak may have escaped more severe water stress at Camp Bullis by having a few roots pen etrate deeper into the epikarst compared to Ashe juniper, although even these deeper water sources appeared to decline over the course of summer of 2009 (Fig. 2b). At both cave sites, sap ow velocities followed simi lar dynamics as soil moisture; both were declining from spring into summer, but remained responsive to midsummer rain events. W hile in-cave drip rates were also declining over the same interval, there were also some dierences relative to sap ow rates. At McCarty Cave, the steepest decline in drip rates was observed in May, at least a month ahead of the most precipitous decline in sap ow. is may indicate that epikarst storage was already low going into the spring of 2009, so that aer rainfall events in spring, there was more water close to the surface than stored deeper in the epikarst. Drip rates increased slightly aer a midsummer rain event, but were generally extremely low aer May. In fact, the inter mittent increase in drip rate could have been caused by an inux of moist air, which temporarily increased rates of condensation on the cave ceiling. At Camp Bullis, two large discharge pulses dominated drip rates in May, and from then on drip rates gradually declined, largely in parallel with the decline of transpiration. However, drip rates did not respond to the two mid-summer rain events, indicating a greater degree of separation between surface and lower epikarst dynamics. It is of note that sap ow in escarpment live oak, at both sites, seemed most closely linked to drip dynamics, following the early de cline in drips rate at McCarty Cave and a more gradual decline at Camp Bullis Cave. is may indicate a greater reliance of this species on water stored closer to the bot tom of the epikarst. W e conclude that tree transpiration is generally in separable from moisture dynamics close to the surface, where water is stored in soil and soil pockets that are eas ily accessed by plant roots. Accessibility by roots to wa ter from lower reaches of the epikarst may vary greatly with epikarst characteristics, but appears to be rare over all, except when specic site and epikarst conditions are met. For example, roots can access deeper epikarst wa ter sources in situations where deeper or larger fractures and/or faults allow vertical penetration of several me ters or more. Relatively dense woodlands seem possible across epikarst types in the Texas Hill Country, as this re gion has experienced wide-spread woody encroachment in the past 100 years (Van Auken 2000). However, we ob served during the exceptionally severe drought of 2011 that local tree mortality rates dier greatly with epikarst characteristics. For example, at the McCarty Cave site, one third of all live oak trees and 6% of juniper trees died in 2011 (Kukowski et al. 2013), but at the Camp Bullis Cave site almost all trees survived. us, under extreme drought conditions, when all near-surface stores of wa ter have dried up, trees may survive by having at least a few roots tap into remaining water pools deeper in the epikarst, provided that these pools exist and local geo logic conditions permit root penetration. W e believe that a signicant portion of storage supplying drips is in the rock matrix and ne fractures, which are largely inaccessible to plant roots. Cretaceous aged carbonates in the Hill Country have relatively high primary and secondary porosity when compared to Pa leozoic carbonates of the USA and elsewhere, and have B. F. S CH W ARTZ S. S CH W INNING B. G ERARD K. R. K UKO W SKI C. L. STINSON & H. C. D AMMEYER

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ACTA CARSOLOGICA 42/2-3 2013 321 many weathered and/or clay-rich beds which contain visible or palpable moisture when drilled or broken even during severe drought. is source of water in the epikarst may be either completely severed from the reach of plant roots, or may serve as a signicant source of transpiration only under the most extreme drought con ditions, aer all other sources have been depleted. Together with other lines of evidence (Heilman et al. 2009), our data contribute to the growing body of literature showing that trees in the Texas Hill Country, and in similar karst areas worldwide, have little capacity for utilizing deeper groundwater, and rely instead upon limited and rapidly depletable shallow water sources in the soil and upper epikarst, and their evolved abilities to survive frequent drought through drought tolerance or avoidance. e hydraulic disconnection between surface and subsurface: how quickly does it form, and what precipitation thresholds must be overcome before re charge occurs? Drip and discharge data from our in-cave sites clearly illustrate that a hydraulic disconnection occurs within 1-2 weeks of the last rain. Hydraulic disconnec tion is characterized by an accumulating moisture decit in the soil and upper epikarst where the bulk of the root zone is located. is dry zone grows over time if there is little or no precipitation, requiring increasingly more water to re-establish a hydraulic connection between the surface and groundwater. If a subsequent precipitation event is insucient to overcome the accumulated mois ture decit, inltrating water cannot break-through into deeper regions of the epikarst, even though regions below the dry zone may still be draining and producing drips. W e determined the precipitation threshold for break-through in September 2011, at the peak of an ex ceptionally severe drought, at McCarty Cave. W e applied the equivalent of a four-hour long, 100 mm rainfall event, ve times during a nine-day period, over a 60 m 2 area ly ing ~ 4.0 m directly above the drip sites in the cave. e rate of application was relatively slow compared to typical rain events in central Texas, which sometimes produce as much as 80 mm in as little as 15 minutes (pers. obs.). In our experiment, 520 mm of articial precipitation were required before we saw a response; a sharp increase in drip rate that was maintained at elevated rates rst by continued application of water, and later by natural precipitation events over the following several months. Consistent with our earlier observations of natural pre cipitation events, vegetation recovered rapidly (visibly greening) within days of the rst water application, long before moisture reached the cave. e break-through threshold of 520 mm, equivalent to almost 60% of aver age annual precipitation, may have been an extreme up per limit, since it was obtained during a time of record drought conditions in Texas. Surprisingly, we observed that the accumulated moisture decit controlled not only diuse and direct inltration, but also macropore ow during overland ow conditions. Aer the 2011 drought, in which only 210 mm of rain fell over a period of 390 days, a single high-intensity precipitation event of 70 mm produced no measurable response at any of the in-cave drip sites. Under no antecedent conditions has a precipitation event smaller than 5 mm resulted in an in-cave response. W e analyzed data from our third eld site at CW AN to predict how much precipitation is required for a re charge response to occur aer varying antecedent con ditions. W e used a logistic regression model with data from three in-cave discharge sites ranging from the main stream to a matrixand fracture-dominated drip. is technique models the probability of a binary response ([1, 0], or [Yes, No]) using continuous predictor vari ables. W e tested a variety of environmental variables in various combinations and used the sample size corrected Akaike Information Criterion (AICc) to select the best model. Predictor variables included cumulative PenmanMonteith Potential Evapotranspiration (PET) for 2 28 weeks, incremented in two-week intervals, the amount of precipitation in a single event, soil moisture prior to a precipitation event, the length of a precipitation event (as a measure of intensity), and discharge at a site immedi ately prior to a rain event. A mixed eect model for data from the three sites incorporating the 14 week sum of PET, soil moisture prior to a rain event, and sum of precipitation during a single event, was able to predict with 88.7% accuracy (n = 115) whether or not a recharge response occurred as a result of a precipitation event (Gerard 2012). W e validated this model by parameterizing a mixed eect model with a randomly selected half of the full dataset from the three sites (n = 57, accuracy = 87.7%), and used it to predict the remaining half of the data (n = 58, ac curacy = 91.4%). ese results indicate that the model has similar predictive power outside of the data used to build it, and the high level of accuracy and consistency across data sets suggests that inltration and recharge at this site, and possibly quite generally, are controlled by relatively simple, quantiable meteorological variables. e epikarst as a variable saturated and perched aquifer At CW AN, between September 2009 and late 2012, we observed a variety of hydrogeologic responses to natural precipitation. ese included uctuations in drip rates on daily, seasonal, or longer time scales, likely as sociated with drainage from large open ssures at the U SING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN ...

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ACTA CARSOLOGICA 42/2-3 2013 322 Fig. 4: A portion of CWAN White Grapes drip data illustrating relationships between changes in drip rate, barometric pressure, and specic conductance that support our hypothesis of gasses trapped in a variably saturated perched aquifer. Changes in pressure cause gas volumes to expand and contract, forcing water into (low SC drip periods) and out of fractures and vuggy porosity (high SC periods). surface receiving water through overland ow, from perched, variably saturated epikarstic aquifers, and from the rock matrix, respectively. ese observations sup port the conceptual model put forth by some research ers (W illiams 1983; Klimchouk 2004) that describes the epikarst as a variably saturated fractured system with perched aquifers draining to underlying conduits via widely spaced large fractures and shas. For example, we found evidence for a perched aquifer in the two-phase discharge dynamics of W hite Grapes drip at CW AN. is drip site is characterized by rapid increases in discharge aer rain, followed by initially high and exponentially decreasing ow (periods of >1000 mL/hr marked A in Fig. 3), a rapid decrease in discharge rates to 300 ml/hr, and nally by exponentially decreasing ow with a atter recession slope than before. e rapid decrease suggests sudden depletion of storage in a perched and highly transmissive aquifer. However, we also found physical and hydrologic ev idence of long-term storage in matrix and fracture/bed ding plane porosity, which is characteristic of the regions Cretaceous limestones. Bedrock exposed in natural cave passages and broken or blasted outcrops at all three sites reveals high matrix and vuggy porosity (when compared to typical Paleozoic carbonates). Evidence supporting matrix/fracture/bedding storage was found in several sets of drip data (Fig. 3), where drips continued to ow, albeit with exponentially decreasing rates during the 390 days of drought. In that time, only one precipitation event caused a measurable recharge response at any of Fig. 3: H ydrographs for the White Grapes and Last Switch drip sites, and the Main Stream dis charge site, in CWAN over a three year period, and cumulative precipitation above the cave. Regions marked with A illustrate a plateau in drip rates that is maintained at >1000 mL/hr at the White Grapes Site prior to behavior attrib uted to the drainage of a perched aquifer (sud den declines in drip rate aer plateaus marked A). Large scale noise in regions marked B is attributed to barometric pressure eects on drip rates from variably saturated fracture/matrix. e Last Switch data represent drip ow domi nated by matrix ow, Main Stream represents both base ow and direct recharge via conduits, and White Grapes represents a mixture of rapid direct recharge, perched aquifer drainage, and drainage from matrix and fracture storage. Gaps represent missing data due to instrument or log ger failure. B. F. S CH W ARTZ S. S CH W INNING B. G ERARD K. R. K UKO W SKI C. L. STINSON & H. C. D AMMEYER

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ACTA CARSOLOGICA 42/2-3 2013 323 the sites (a 70 mm rainfall caused a small overland ow/ direct recharge response at Main Stream on Septem ber 28, 2012). us, soil and upper epikarst storage was so depleted that almost no precipitation event passed through to contribute to deep inltration and drainage. However, drips sites remained active, suggesting a highly persistent source of water in deeper matrix/fracture stor age. A second line of evidence for storage in matrix and fractures is found in portions of the W hite Grapes hy drograph data marked B in Fig. 3 that appear noisy. Attempting to explain these high-frequency signals, we discovered inverse relationships between baromet ric pressure, drip rate, and specic conductivity (SC) (Fig. 4). To our knowledge, this phenomenon has only been reported in a few previous drip studies (Genty and Deandre 1998), but we found these signals at all drip sites and the main stream site in CW AN. us, the phe nomenon seems to be pervasive and possibly driven by ubiquitous processes that may have been overlooked or ignored in other karst studies. W e recently proposed a simple model to explain the inverse relationship between drip rate and baromet ric pressure based on a mechanism involving variable saturation of the epikarst (Gerard 2012). Because the epikarst is part of the unsaturated zone, matrix, fracture, and small perched aquifers within it are subject to spatial and temporal variation in saturation. As epikarst ma trix and fractures drain during drought, gasses replace water in the voids. en, during recharge events, a ris ing perched water table traps gas pockets in dead-end fractures and vuggy porosity. W ith atmospheric pres sure (P atm )changes at the surface synchronously aect ing pressure changes in the cave (though with a slightly damped signal), trapped gas pockets compress or expand in response to increases or decreases in P atm W hen P atm increases, gas pockets compress and water moves from main ow paths into dead-end fractures and vugs. As P atm decreases, gas pockets expand and water is pushed back into the main ow paths from where some fraction drains out and reaches the drip monitoring site. In eect, the periodic uctuation between gas compression and expansion pumps water into and out of dead-end com partments, explaining the inverse relationship between P atm and drip rate. If this model is correct, drips should have higher SC during periods of low P atm when water is pushed out of the dead-end compartments. is is because this water will have had longer residence time in dead-end compartments (relative to the main ow paths) and in creased time to dissolve rock. W e tested this prediction and found it to be supported by the data (Fig. 4). us, drip rate dynamics are consistent with the lower regions of the epikarst functioning like a variably saturated, perched aquifer. C ONCLUSIONS T HE EPIKARST AS A CRITICAL ZONE Our work in the Central Texas Hill Country has revealed an epikarst system that is controlled by complex interac tions between physical, hydrogeochemical and biological processes. e biological dimension is oen overlooked in karst research, but we nd that a detailed understand ing of karst ecohydrology is critical to predicting the dy namics of epikarst drainage and storage, and subsequent recharge to the aquifer. For future research, we therefore promote a concep tual model of the epikarst system as an integrated criti cal zone, extending from the treetop to the base of the epikarst, with multiple internal interactions and feed backs. W ithin this critical zone, trees in the Edwards Plateau primarily access soil and shallow epikarst water within the upper one to two meters of the surface. W hen cumulative precipitation overcomes any accumulated moisture decit within this zone, inltrating water moves below this depth and is essentially inaccessible to evapo transpiration. It then contributes to a deeper and wellmixed storage compartment in which matrix and frac ture porosity and permeability dominate, and residence times are probably on the order of several years or more. is variably (temporally and spatially) saturated zone extends from the base of the rooting zone downward to the regional water table, and within this range, some ar eas may also contain one or more perched aquifers. Our study sites are representative of much of the Central Texas Hill Country where Edwards or Glen Rose Fm. carbonates outcrop, providing a foundation for future studies in the region. In ongoing and future re search, we seek to apply and further rene the method ologies and models we developed for siteto small drain age basin-scales and apply them to larger karstic drainage basins in the Hill Country. Further, we plan to explore the role of epikarst and variation in its properties across the landscape in vegetation response to extreme climatic events, exemplied by the Texas Drought of 2011. U SING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN ...

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ACTA CARSOLOGICA 42/2-3 2013 324 A CKNO WLEDGMENTS W e thank landowners Jim and Shannon Brotherton for allowing unlimited access to McCarty Cave, Tom Sum mers (owner) and Mike Burrell (Manager) for allowing and facilitating access to CW AN, Matthew L. Cooksey and Chris ibodaux for allowing access to Headquar ters Cave and Camp Bullis, and ZARA Environmental for sharing drip data at Headquarters Cave. anks also to Philip Ramirez, Gabrielle Timmins, Benjamin Tobin, Benjamin Hutchins, Jacob Martin, and many others for their assistance in installing and maintaining equipment, and collecting samples above and below ground. Fund ing for this research was provided in part by a State of Texas THECB ARP grant (Project # 003615-0021-2007), and a USGS 104g grant (Award # G09AP00147). e potential eects of barometric pressure on rates of epikarst evolution is another area of continuing re search interest. How important is the periodic pumping of uid into and out of dead-end compartments for sol ute transport out of the epikarst and into the lower con duit system? REFERENCES Arbel, Y., Greenbaum, N., Lange, J. & M. Inbar, 2010: In ltration processes and ow rates in developed karst vadose zone using tracers in cave drips. Earth Sur face Processes and Landforms, 35, 1682. Deng, Y., Jiang, Z.C. & X.M. Qin, 2012: W ater source partitioning among trees growing on carbonate rock in a subtropical region of Guangxi, China. Environmental Earth Sciences, 66(2), 635. Dixon, R., 2000: Climatology of the Freeman Ranch, Hays County, Texas. Freeman Ranch Publication Series No 3, pp 2, Southwest State Univer sity, San Marcos, Texas: 2. Ehleringer, J., Roden, J. & T. Dawson, 2000: Assessing ecosystem level water relations through stableiso tope ratio analysis. Methods in Ecosystem Science. O. Sala, R. Jackson, H. Mooney and R. Howarth. New York, Springer: 181. Estrada-Medina, H., Graham, R.C., Allen, M.F., Jime nez-Osornio, J.J. & S. Robles-Casolco, 2013: e importance of limestone bedrock and dissolution karst features on tree root distribution in northern Yucatan, Mexico. Plant and Soil, 362(1), 37. Genty, D. & G. Deandre, 1998: Drip ow variations un der a stalactite of the Pre Nol cave (Belgium). Evi dence of seasonal variations and air pressure con straints. Journal of Hydrology, 211, 208. Gerard, B., 2012: Eects of Environmental Parameters and Precipitation Dynamics on Inltration and Recharge into the Trinity Aquifer of Central Texas. Department of Biology. San Marcos, Texas State University. M.S.: 61. Gonantini, R., 1978: Standards for Stable Isotope Measurements in Natural Compounds. Nature, 271(5645), 534. Granier, A., 1985: A New Method of Sap Flow Measure ment in Tree Stems. Annales Des Sciences Forest ieres, 42(2), 193. Heilman, J.L., McInnes, K.J., Kjelgaard, J.F., Owens, M.K. & S. Schwinning, 2009: Energy balance and water use in a subtropical karst woodland on the Ed wards Plateau, Texas. Journal of Hydrology, 373, 426. Jackson, R.B., Moore, L.A., Homan, W .A., Pockman, W .T. & C.R. Linder, 1999: Ecosystem rooting depth determined with caves and DNA. Proceedings of the National Academy of Sciences of the United States of America, 96, 11387. JMP, 1989: Version 9, SAS Institute Inc. Cary, NC. Jones, W .K., Culver, D.C. & J.S. Herman, 2003: In: Intro duction. Epikarst, Shepherdstown, WV, Karst W a ters Institute. Klimchouk, A.B., 2004: Towards dening, delimiting and classifying epikarst: Its origin, processes and variants of geomorphic evolution. Speleogenesis and Evolution of Karst Aquifers www.speleogen esis.info 2(1). Kukowski, K.R., Schwinning, S. & B.F. Schwartz, 2013: Hydraulic responses to extreme drought conditions in three co-dominant tree species in shallow soil over bedrock. Oecologia 171:819. B. F. S CH W ARTZ S. S CH W INNING B. G ERARD K. R. K UKO W SKI C. L. STINSON & H. C. D AMMEYER

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ACTA CARSOLOGICA 42/2-3 2013 325 Kuniansky, E.L. & K.Q Holligan, 1994: Simulations of ow in the Edwards-Trinity aquifer system and con tiguous hydraulically connected units, west-central Texas. 93. W .-R. I. Report, U.S. Geological Survey. 93: 40. Lee, E.S. & N.C. Krothe, 2001: A four-component mix ing model for water in a karst terrain in south-cen tral Indiana, USA. Using solute concentration and stable isotopes as tracers.Chemical Geology, 179, 129. Mace, R.E., Chowdhury, A.H., Anaya, R. & S.C. W ay, 2000: Groundwater Availability of the Trinity Aqui fer, Hill Country Area, Texas: Numerical Simula tions rough 2050., Texas W ater Development Board. Report: 119. Mangin, A., 1975: Contribution a l'etude hydrody namique des aquifers karstiques: DES thesis, Univ. Dijon, France, (Ann, Speleo. 1974, v. 29, nos. 3 and 4, 1975, v. 30, n. 1). DES. Nie, Y.P., Chen, H.S., W ang, K.L., Tan, W ., Deng, P.Y. & J. Yang, 2011: Seasonal water use patterns of woody species growing on the continuous dolostone out crops and nearby thin soils in subtropical China.Plant and Soil, 341(1), 399. NSF, 2012. Critical Zone Observatories (CZO). Re trieved 12-28-2012, 2012, from http://www.nsf.gov/ pubs/2012/nsf12575/nsf12575.htm Q uerejeta, J.I., Estrada-Medina, H., Allen, M.F. & J.J. Jimenez-Osornio, 2007: W ater source partition ing among trees growing on shallow karst soils in a seasonally dry tropical climate.Oecologia, 152(1), 26. Q uerejeta, J.I., Estrada-Medina, H., Allen, M.F., Jime nez-Osornio, J.J. & R. Ruenes, 2006: Utilization of bedrock water by Brosimum alicastrum trees grow ing on shallow soil atop limestone in a dry tropical climate.Plant and Soil, 287(1), 187. R Development Core Team, 2012: A Language and envi ronment for statistical computing. Vienna, Austria, from http://www.R-project.org Rong, L., Chen, X., Chen, H., W ang, S.J. & X.L. Du, 2011: Isotopic analysis of water sources of mountainous plant uptake in a karst plateau of southwest China.Hydrological Processes, 25(23), 3666. Schwinning, S., 2008: e water relations of two ever green tree species in a karst savanna.Oecologia, 158, 373. Schwinning, S., 2010: Ecohydrology Bearings Invited Commentary: e ecohydrology of roots in rocks.Ecohydrology, 3, 238. Smart, P. L. & H. Friederich, 1986: In: W ater movement and storage in the unsaturated zone of a maturely karstied carbonate aquifer, Mendip Hills, Eng land. Proceedings of the Environmental Problems in Karst Terranes and their Solutions Conference, Dublin, Ohio, National W ater W ell Association. Van Auken, O.W ., 2000: Shrub invasion of North Ameri can semiarid grasslands.Annual Review of Ecol ogy and Systematics, 31, 197. Veni, G., 1994: Geomorphology, H ydrogeology, Geochem istry, and Evolution of the Karstic Lower Glen Rose Aquifer, South-central Texas. Department of Geo sciences, Pennsylvania State University. Doctoral Dissertation: 721. W illiams, P.W ., 1983: e role of the subcutaneous zone in karst hydrology.Journal of Hydrology, 61, 45. U SING HYDROGEOCHEMICAL AND ECOHYDROLOGIC RESPONSES TO UNDERSTAND EPIKARST PROCESS IN ...



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S PRING DISCHARGE RECORDS A CASE STUDY H IDROGRAMI KRA KIH IZVIROV : TUDIJA NA PRIMERU IZVIRA D EVILS I CEBO X rM ISSOURI ZDAf Carol M. WICKS 1 Izvleek UDK 556.36:551.44(737.8) Carol M. Wicks: Hidrogrami krakih izvirov: tudija na prime ru izvira Devils Icebox (Missouri, ZDA) Hidrogrami so vsota procesov in reakcij v krakem zaledju izvira. lanek nudi kratek pregled uporabe hidrogramov pri doloitvi notranje strukture krakega vodonosnika, izboljanju modelov napajanja in praznenja ter doloanja hidrodinaminih parametrov krakega zaledja. Pri tem uporabimo podatke iz krakega izvira Devils Icebox. V primeru tega izvira hidro grami ne povedo veliko o notranji strukturi vodonosnika, zato pa ve o napajanju zaledja. Model polnjenja in praznenja vodono snika ni vrnil vhodnih podatkov iz katerih smo doloili model ske parametre in je neuporaben za napovedovanje. Upo raba principa ohranitve mase za doloanje hidrodinaminih parame trov, je dober pristop, a so v naem primeru manjkali nekateri kljuni podatki. Kraki hidrologi potrebujejo ve kvantitativnih podatkov sledenj in dolge asovne nize podatkov o dotoku in iztoku z visoko asovno loljivostjo. Kljune besede: Hidrogram, hidrologija, kras. 1 E235 Howe-Russell-Knien Geosciences Complex, Department of Geology and Geophysics, Louisiana State University, Baton Rouge, LA 70803, e-mail: cwicks@lsu.edu Received/Prejeto: 15.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 339, POSTOJNA 2013 Abstract UDC 556.36:551.44(737.8) Carol M. Wicks: Spring discharge records a case study Spring discharge records integrate of all the processes and the reactions occurring within a karst basin. A brief summary of the use of discharge records as a means to constrain the inter nal structure of karst basins, as means to constrain rainfallruno models for karst basin, and as a means to determine the value of hydrodynamic parameters of karst basins is presented. Data collected from Devils Icebox, a karst basin spring in Mis souri, USA, were used to assess these approaches to character izing karst basins. For Devils Icebox, most of the discharge re sponses do not record information about the internal structure of the basin rather the responses record information about the recharge to the basin. A rainfall-runo model failed to repro duce the data from which model parameters were derived and has little utility in a predictive mode. Use of conservation of mass equations as a means to derive hydrodynamic parameters is a useful approach, although critical data are lacking. More generally, karst hydrologists need quantitative tracer data and long-term, high-resolution temporal data of the input(s) to and the output(s) from karst basins. Keywords: Hydrograph, hydrology, karst. I NTRODUCTION From the early 1900s to the present, the structure and functioning of karst basins (springsheds) has been in ferred from the physical responses of those basins to recharge events (Ashton 1966; Ford & W illiams 2007; Hess & W hite 1988; Vesper & W hite 2003). e histori cal and continued use of hydrographs has been driven by the idea that the output from a karst basin records information about the reactions and the processes that

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ACTA CARSOLOGICA 42/2-3 2013 340 occur within the basin ( W hite 2002; 2007). us, the spring discharge was and is seen as an integration of all the processes and the reactions occurring within that basin. e present article is an overview of the use of hydrographs as a means to constrain the internal struc ture of basins (W hite & Deike 1989), as means to con strain rainfall-runo models (Labat et al. 1999), and as a means to determine the value of hydrodynamic pa rameters (Dreiss 1989b; Ferrick 2005). INTERNAL STRUCTURE W hether the physical and chemical responses of karst basins to recharge events carry information about the internal structure of a karst basin or carry information about the intensity and duration of the recharge event has been the subject of considerable research. Early interpretations were based on the assumption that the responses carried information about the internal struc ture of the basin. Thus, an exponential decay relation was used to derive values for characteristic response times for draining the conduits, the fractures, and the rock matrix that are present in karst basin (Ford & W illiams 2007). Subsequently, researchers interpreted the responses in terms of unexplorable passages (Ash ton 1966), the proportion of air-filled passages (Brown 1970; Brown 1973), the presence of constrictions (Ha lihan & W icks 1998; Halihan et al. 1998; Vineyard, 1958) or fine structures (Hess & W hite 1988), the ge ometry of submerged conduits (Grasso et al. 2003a; Grasso et al. 2003b) or of the basin (Kovcs & Per rochet 2008; Kovcs et al. 2005), and the ratio of the surface area to volume of conduits (Birk & Hergarten 2010; Birk et al. 2004). As the number of feasible in terpretations increased, the purposefully ambiguous terms of quick and slow flow were used to describe the observed physical responses of karst basins to recharge events ( W hite 2007). Yet, all of these interpretations are based on the assumption that the spring responses actually record information about the internal struc ture of the basin. Recently, a dimensionless number that can be used to determine whether the discharge hydrograph does record information about the inter nal structure of a basin or about the input hydrograph to that basin has been defined (Covington et al. 2009); however use the dimensionless number requires infor mation about both the input function and the spring response. Given the current state of knowledge, how should the physical and chemical responses of karst basins be interpreted? RAINFALLRUNOFF MODELS e goal of rainfall-runo modeling is to predict dis charge (output) from a basin for a given recharge (input) event (Dooge 1959). e rainfall-runo models were rst applied to karst basins by analyzing the records of daily rainfall into and springow out of three karst ba sins (Knisel 1972). e rainfall-runo model is based the denition of a transfer function (kernel function, unit hydrograph). Commonly, a single, linear transfer function has been derived (Dreiss 1983; Hoke & W icks 1997; W icks & Bohm 2000). Some researchers are de veloping nonlinear transfer functions (Denic-Jukic & Jukic 2003) or combining rainfall-runo modeling with hydrograph separation to derive the most useful rela tion between rainfall and runo (Pinault et al. 2001a; Pinault et al. 2001b). However, numerous studies have pointed out that the response of karst basins is inher ently non-linear and non-stationary (Labat et al. 2000a; 2000b; 2002). us, the question is Can rainfall-runo models be applied to karst basins?. HYDRODYNAMIC CHARACTERISTICS e conservation of mass equations are widely used throughout the hydrologic sciences to explain the diu sion of ood surges (Ferrick 2005; Ferrick & Goodman 1998) and the advection and dispersion of solutes (Din gman 1984; Freeze & Cherry 1979). W hen solved us ing appropriate initial and boundary conditions, these governing equations permit the determination of the hydrodynamic characteristics of streams and ground water. W ithin karst hydrology, application of these equations has been associated with movement of trac ers (Dreiss 1989b; Field & Leij 2012). e movement of ood surges, which have been noted to travel quickly (Ford & W illiams 2007; W hite 1988) and the transport of solutes or contaminants (not tracers) have rarely been investigated using conservation of mass equations. W hat can karst hydrologists learn through thoughtful application of time moment analysis or the advectiondispersion type equations? OBJECTIVES e objective of the current research is to begin to ad dress the questions: How should the physical and chemical responses of karst basins to recharge events be interpreted?, Can rainfall-runo models be applied to karst basins?, and W hat can karst hydrologists learn through thoughtful application of time moment analy sis or the advection-dispersion type equations?. e approach taken is that of a case study in which all of these questions are addressed for the same basin, Devils Icebox basin. CAROL M. WICKS

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ACTA CARSOLOGICA 42/2-3 2013 341 Devils Icebox basin is in Boone County in central Mis souri USA and it is where the endemic pink planaria are found. Even though the basin is small (32 km 2 ) and the length of the mapped passage is modest (~10km), the number of scientic studies that have been conducted in the basin is relatively high. e studies include role of chert in controlling passage development and location (Hargrove 1968), water quality (Lerch et al. 2001; W icks & Engeln 1997), sediment transport (Dogwiler & W icks 2004), and the studies of the pink planaria (Sutton 2004; W icks et al. 2010). A detailed description of the basin is provided in the cited works and will not be repeated here. Briey, the areal extent of the Devils Icebox basin is well dened (Vandike 1983; Vandike & Schulte 1984) and the location of cave passages and stream channels are well known (Deike et al. 1960). ere is a primary stream channel linking the losing stream to the spring and that stream channel is wide (relative to depth) and sinuous with impermeable sides; and the in-cave stream has a free surface. Nearly all of the wa ter owing along the in-cave stream is from the losing sur face stream and only a minor portion is delivered via sink holes and seepage (W icks 1997a). e discharge and specic conductance data that were collected from Devils Icebox are presented in Lerch et al. (2001) and in Fig. 1. Nine storm events were recorded over the time period and those storms are indicated in Fig. 1. MATERIALS AND METHODS INTERPRETATIONS INTERNAL STRUCTURE An exponential decay relation has been used to derive values for the characteristic response time for draining conduits, fractures, and rock matrix that are present in a karst basin (Ford & W illiams 2007). Plotting the loga rithm of discharge against time allowed the identication of three linear segments (conduits, fractures, and rock matrix). For Devils Icebox, only two linear segments are obvious (Fig. 2). Certainly, the conceptual model of con duits, fractures, and rock matrix is valid; however, the in terpretation of the exponential approach is problematic as the draining of one of the components is not apparent in the response of the basin to recharge events (Fig. 2). e discharge from the Devils Icebox was de scribed as ow past a constriction and draining of pooled water (Vineyard 1958). In a test of that concep tual model, high-ow events were successfully modeled using a reservoir-constriction model (Halihan & W icks Fig. 1: Record of discharge (plot ted as logarithm of discharge; upper panel) and specic con ductance (lower panel) from Devils Icebox cave for April 1999 to March 2000 including nine re charge events. S PRING DISCHARGE RECORDS A CASE STUDY

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ACTA CARSOLOGICA 42/2-3 2013 342 1998; Halihan et al. 1998). However, Covington et al. (2009) have shown that many of the physical responses to recharge events recorded at the Devils Icebox record information about the input function and do not reveal information about the internal structure of the basin, such as the presence of constrictions. Can these dier ent interpretations, that of Halihan and that of Coving ton, be reconciled? Under typical ow conditions, such as those used in the study by Covington et al. (2009), the discharge from the Devils Icebox reects input hydro graphs. Under high ow conditions, such as those used in the study by Halihan and W icks (1998), the spring hydrographs do record the presence of the owpath constriction and pipe-full conditions within the nor mally air-lled passageways. For Devils Icebox, spring hydrographs record information about the input to the basin and not information about the internal structure of the basin. For rarer extremely high discharge events, the spring discharge does record information about the internal constrictions and pipe-full conditions. Moving forward, karst hydrologists should instru ment both springs and losing reaches for extended peri ods, including droughts and oods, in order to place any individual results into appropriate hydrologic context. In the simplest case, the data recorded would include depth (stage) of water as a function of time. Researchers could then easily compare the output hydrograph (spring dis charge) to the input hydrograph (losing reach hydro graph) and determine whether the spring hydrographs were recording information about the input hydrograph or about the internal structure. Having the input hydro CAROL M. WICKS Fig. 2. P lot of the natural logarithm of the discharge for each of the nine recharge events as a function of time since maximum discharge. Note that there are two linear segments; one for higher discharges (early times) and the other for lower discharges (later times). V alues of the mean plus and minus one standard devia tion of the slopes are given. Fig. 3: a) Unit hydrographs de rived from each of the nine re charge events noted in Fig. 1 are plotted against time. b) e mea sured discharge (thin line; lehand vertical axis) is compared to the discharged calculated by using a representative unit hy drographs, the rainfall record ad justed for potential evapotranspi ration, and the basin area (thick line; right-hand vertical axis). Note the magnitude dierence between the two vertical axes, the presence of measured discharge events that were not calculated (near May and J une 1999), and the presence of calculated events that were not real events (near Dec 1999). a) b)

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ACTA CARSOLOGICA 42/2-3 2013 343 graph data is critical to our eorts to knowing how to interpret the response of these basins. RAINFALLRUNOFF MODEL Rainfall-runo models relate aerially distributed rain fall to runo at a particular location along a stream, usually a location with a gaging station, through a trans fer function (Dreiss 1983). Deriving a unit hydrograph is the simplest method of determining a transfer func tion. For Devils Icebox, unit hydrographs (Fig. 3a) were calculated for each of the nine discharge events (Fig. 1). Variation in the unit hydrographs is apparent with some having smaller widths at a dimensionless height of 0.5 compared to other unit hydrographs. is variation points out the diculty of selecting a single unit hydrograph to serve as the representative trans fer function. Further, once that representative trans fer function is selected, the calculations of discharge made using that a single transfer function (along with rainfall record corrected for evapotranspiration and the basin are) do not match the data from which the transfer function was derived (Fig. 3). ere are events that were calculated to have occurred when measured events are lacking (near Dec 1999); and there are meas ured events for which no calculated event occurred (May and June 1999). Fundamentally, a rainfall-runo model is a wa ter balance for a basin. For the Devils Icebox, the sur face area of the basin that drains to the spring is well known (Lerch et al. 2005). at surface area includes land that drains into the losing stream and surface area that drains downward through the sinkhole plain. Even though the basin is well known, the water balance has not been closed, mainly due to an ungaged overow channel that funnels out of the basin during high-ow events (W icks 1997b). Research focused on Big Spring and on Maramec Spring basins (also in Missouri, USA) showed that the match between the observed discharged and the discharge calculated based on rain fall-runo model was poor (Dreiss 1989a; 1989b). Un less the water balance for a particular karst basin can be closed, there is limited use for rainfall-runo model for that basin. For the Devils Icebox, ~90% of the wa ter that exits at the spring during recharge events had a source in the losing stream. us, a surface runospring discharge (a runo-runo) model might be more appropriate for predicting the timing and mag nitude of the peak discharge. is concept also aligns with our understanding that the spring discharge re cords information about the input function (surface runo) and not about the internal structure of the ba sin. Such a runo-runo model requires that the los ing stream be monitored over the same time and us S PRING DISCHARGE RECORDS A CASE STUDY ing the same sampling interval as is used for the spring discharge record. Data from the losing stream were not collected. HYDRODYNAMIC CHARACTERISTICS Solution to the linear diusion equation (conserva tion of mass) that governs the diusion of ood surges (Ferrick 2005; Ferrick & Goodman 1998) permits the determination of the celerity and noninertial diusion coecient. For the Devils Icebox, the calculated celer ity and noninertial diusion coecient ranged from 0.05 0.78 m s and from 0 10 m 2 s respectively (W icks & Loper 2008). us, the celerity at which storm surges move along in-cave streams is within the range and toward the lower values reported for rivers (0.0 to 3.8 m s ; Ferrick 2005). For the Devils Icebox, the values of the diusion coecient are higher than values reported for rivers (0.005 0.80 m 2 s ; Ferrick 2005). e bound ing walls of the in-cave stream provide more resistance to ow than the open-back channels of a river, resulting in slower the movement of ood surges (lower celerity) and enhanced diusion of the ood surge (higher diusion coecients). Solutions to the equations that govern the advection and dispersion of solutes (Freeze & Cherry 1979) permit the determination of the average eective ow velocity, eective dispersion coecient, and eective dispersivity (Dreiss 1989b). e values for these parameters should be calculated from tracer test data (Field 2002a; 2002b). For Devils Icebox, these critical data are not available; however, there is a plethora of specic conductance data (Fig. 1). Useful indicators of hydrodynamic parameters can be obtained by using specic conductance data and the moments about the means method (Dreiss 1989b). e coecient of variation, C v is related to the distri bution and interconnectedness of the travel owpaths in the basin and the skewness coecient, is related to symmetry of travel distances. For the Devils Icebox, there is one main owpath from the losing stream to the spring and the interconnection of that owpath is high, C v is low (0.83 to 1.31) as anticipated. e coecient of skewness for Devils Icebox is also low (0.32 to 0.97), given that there is very little variation in symmetry along one owpath. SUMMARY Even for the well-studied Devils Icebox basin, critical data (quantitative tracer data, long-term records of discharge at the spring and at the losing streams) are lacking; how ever, insight into the physical and chemical processes oc curring within the basin was possible by applying a com bination of approaches. For most recharge events in the Devils Icebox basin, the physical and chemical responses

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ACTA CARSOLOGICA 42/2-3 2013 344 CONCLUSIONS How should the physical and chemical responses of karst basins to recharge events be interpreted? For karst ba sins whether that basin is dominated by in-cave streams or by ow through phreastic conduits, progress can be made by assessing the relation between the discharge hydrograph and the input hydrograph (Covington et al. 2009). If the response records information about the in ternal structure of the basin, then existing techniques can be used to determine the properties of that internal structure. is requires that we monitor the input to and the output from karst basins over the same time peri ods and with the same temporal resolution (LeGrand & Stringeld 1973). Can rainfall-runo models be applied to karst ba sins? Rainfall-runo models (or runo-runo models) require that the water budget for a basin can developed. For many karst basins, water budget cannot be devel oped as the basin area is unknown (or variable depend ing on ow condition) and lack of appropriate methods to correct precipitation for changes in soil moisture and evapotranspiration. Until water budgets are developed, rainfall-runo models are of limited utility. Karst hy drologists need to try to close the water balances for the basins that they study. W hat can karst hydrologists learn through thought ful application of time moment analysis or the advec tion-dispersion type equations? Most karst hydrologists commonly record the data needed to determine the val ue of moments about the mean. W ith quantitative trac er data, eective dispersivity can be calculated (Dreiss 1989b). Karst hydrologists need to reported the values of a few key parameters (moments about the mean), so that we can develop an understanding of how these param eters vary and so that we can compare those calculated values to values from other karst basins, from surface streams (rivers), and from groundwater basins. ese comparison would allow karst hydrologists to place karst hydrology within the broader framework of hydrologic sciences (Herman et al. 2009). ere are fundamental issues to address. W e need to routinely monitor the input(s) to and the output(s) from karst basins for long periods of time. Using those data, we need to develop water budgets for basins. W e need quantitative tracer data for basins, even in wellcharacterized basins, that permits calculation of key pa rameters (dispersivity, eective velocities). ACKNO WLEDGEMENTS e Karst W aters Institute organized a conference Car bon and Boundaries in Karst that provided the impetus for the preparation of this article. e Karst Hydrology Research Group at LSU reviewed an early version of this article and their comments improved the article. e comments of two anonymous reviewers are greatly ap preciated and resulted in an improved article. e sta of Rock Bridge Memorial State Park in Missouri provided access to the cave and permitted this study. NSF Grant EAR #9870423 and AER #1141745 to C.M. W icks pro vided partial support for this study. CAROL M. WICKS of the basin to recharge events should be interpreted as a recorder of information about the input function to that basin and not as a recorder of information about the internal structure of the basin. For infrequent and very large recharge events, the spring discharge does record the development of pipe-full conditions. Developing a runo-runo model (contrasted with a rainfall-runo model) would allow prediction of the timing and dura tion of peak discharge events. e runo-runo model would also sidestep the issue of failure to close the water balance for the basin, as a runo-runo model would only track water that ows into the subsurface, eectively closing the water balance. Such a model aligns with our understanding that the spring hydrographs record the input function. For a basin dominated by ow along a single owpath, the solutions to various conservation of mass equations can provided detailed information about the movement of ood surges and solutes through Devils Icebox.

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ACTA CARSOLOGICA 42/2-3 2013 345 Ashton, K., 1966: e analysis of ow data from karst drainage systems.Transactions of the Cave Re search Group, 7, 36. Birk, S. & S. Hergarten, 2010: Early recession behaviour of spring hydrographs.Journal of Hydrology, 387, 1, 24. DOI: 10.1016/j.jhydrol.2010.03.026. Birk, S., Liedl, R. & M. Sauter, 2004: Identication of localised recharge and conduit ow by combined analysis of hydraulic and physico-chemical spring responses (Urenbrunnen, SW -Germany).Journal of Hydrology, 286, 1, 179. Brown, M.C., 1970: New methods in karst hydrology.Transactions-American Geophysical Union, 51, 4, 285. Brown, M.C., 1973: Mass balance and spectral analysis applied to karst hydrologic networks.W ater Re sources Research, 9, 3, 749. Covington, M., W icks, C. & M. Saar, 2009: A dimension less number describing the eects of recharge and geometry on discharge from simple karstic aqui fers.W ater Resour. Res, 45, 11, W11410. Deike, G., Hopson, H., Sturmfels, G., Dieke, R., Barn holtz, R. & K. Lang, 1960: Devils Icebox, Boone County, Missouri. Chouteau Grotto, National Spe leological Society, 1 sheet, 1:2740 scale. Denic-Jukic, V. & D. Jukic, 2003: Composite transfer functions for karst aquifers.Journal of Hydrology, 274, 1, 80. Dingman, S.L., 1984: Fluvial Hydrology. New York, W .H. Freeman and Company, 383 p. Dogwiler, T. & C.M. W icks, 2004: Sediment entrainment and transport in uviokarst systems.Journal of Hydrology, 295, 1, 163. Dooge, J.C., 1959: A general theory of the unit hy drograph.Journal of Geophysical Research, 64, 241. Dreiss, S.J., 1983: Linear unit-response functions as in dicators of recharge areas for large karst springs.Journal of Hydrology, 61, 31. Dreiss, S.J., 1989a: Regional scale transport in a karst aquifer: 1. Component separation of spring ow hydrographs.W ater Resources Research, 25, 117. Dreiss, S.J., 1989b: Regional scale transport in a karst aquifer: 2. Linear systems and time moment analy sis.W ater Resources Research, 25, 126. Ferrick, M.G., 2005: Simple wave and monoclinal wave models: River ow surge applications and implications.W ater Resources Research, 41, doi:10.1029/2004WR003923, 2005. Ferrick, M.G. & N.J. Goodman, 1998: Analysis of linear and monoclincal river wave solutions.Journal of Hydraulic Engineering, 728. Field, M.S., 2002a: Ecient hydrologic tracer-test design for tracer-mass estimation and sample-collection frequency, 1. Method development.Environmen tal Geology, 42, 7, 827. Field, M.S., 2002b: Ecient hydrologic tracer-test design for tracer-mass estimation and sample-collection frequency, 2. Experimental results.Environmental Geology, 42, 7, 839. Field, M.S. & F.J. Leij, 2012: Solute transport in solution conduits exhibiting multi-peaked breakthrough curves.Journal of Hydrology, 440, 26. Ford, D.C. & P.W W illiams, 2007: Karst Hydrogeology & Geomorphology. John W iley & Sons, Inc., New York, 576. Freeze, R.A. & J.A. Cherry, 1979: Groundwater. PrenticeHall, Englewood Clis, 604. Grasso, D.A., Jeannin, P.-Y. & F. Zwahlen, 2003a: Erra tum to "A deterministic approach to the coupled analysis of karst springs' hydrographs and chemo graphs" [Journal of Hydrology 271 (2003) 65].Journal of Hydrology, 279, 1, 291. Grasso, D.A., Jeannin, P.Y. & F. Zwahlen, 2003b: A deter ministic approach to the coupled analysis of karst springs' hydrographs and chemographs.Journal of Hydrology, 271, 1, 65. Halihan, T. & C.M. W icks, 1998: Modeling of storm re sponses in conduit ow aquifers with reservoirs.Journal of Hydrology, 208, 1, 82. Halihan, T., W icks, C.M. & J.F. Engeln, 1998: Physical re sponse of a karst drainage basin to ood pulses: ex ample of the Devil's Icebox cave system (Missouri, USA).Journal of Hydrology, 204, 1, 24. Hargrove, G., 1968: Relation of a chert zone to develop ment of Devils Icebox Boone County Missouri.Missouri Speleology, 10, 15. Herman, E.K., Toran, L. & W .B. W hite, 2009: Q uantify ing the place of karst aquifers in the groundwater to surface water continuum: A time series analysis study of storm behavior in Pennsylvania water re sources.Journal of Hydrology, 376, 1, 307. Hess, J.W & W .B. W hite, 1988: Storm response of the karstic carbonate aquifer of southcentral Kentucky.Journal of Hydrology, 99, 235. Hoke, J.A. & C.M. W icks, 1997: Contaminant transport in karst aquifers.e Engineering Geology and Hydrogeology of Karst Terranes. Rotterdam: AA Balkema, 189. S PRING DISCHARGE RECORDS A CASE STUDY REFERENCES

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ACTA CARSOLOGICA 42/2-3 2013 346 Knisel, W .G., 1972: Response of karst aquifers to re charge, Hydrology Papers. Colorado State Univer sity, Fort Collins, CO, 1. Kovcs, A. & P. Perrochet, 2008: A quantitative approach to spring hydrograph decomposition.Journal of Hydrology, 352, 1, 16. Kovcs, A., Perrochet, P., Kirly, L. & P.-Y. Jeannin, 2005: A quantitative method for the characterisation of karst aquifers based on spring hydrograph analy sis.Journal of Hydrology, 303, 1, 152. Labat, D., Ababou, R. & A. Mangin, 1999: Linear and nonlinear input/output models for karstic spring ow and ood prediction at dierent time scales.Stochastic Environmental Research and Risk As sessment, 13, 5, 337. Labat, D., Ababou, R. & A. Mangin, 2000a: Rainfall-run o relations for karstic springs. Part I: convolution and spectral analyses.Journal of Hydrology, 238, 3, 123. Labat, D., Ababou, R. & A. Mangin, 2000b: Rainfall-run o relations for karstic springs. Part II: continuous wavelet and discrete orthogonal multiresolution analyses.Journal of Hydrology, 238, 3, 149. Labat, D., Ababou, R. & A. Mangin, 2002: Multireso lution cross-analysis of rainfall rates and karstic spring runos.Comptes Rendus Geoscience, 334, 8, 551. LeGrand, H.E. & V.T. Stringeld, 1973: Karst hydrology A review.Journal of Hydrology, 20, 2, 97. Lerch, R.N., Erickson, J.M. & C.M. W icks, 2001: Inten sive water quality monitoring in two karst water sheds of Boone County, Missouri.Proceedings of the 15th National Cave and Karst Management Symposium (October 16, 2001), National Cave and Karst Management Symposium Steering Com mittee, Tuscon, AZ. Lerch, R.N., W icks, C.M. & P.L. Moss, 2005: Hydro logic characterization of two karst recharge areas in Boone County, Missouri.Journal of Cave and Karst Studies, 67, 158. Pinault, J.L., Pauwels, H. & C. Cann, 2001a: Inverse modeling of the hydrological and the hydrochemi cal behavior of hydrosystems: Application to nitrate transport nd denitrication.W ater Resources Re search, 37, 8, 2179. Pinault, J.L., Plagnes, V., Aquilina, L. & M. Bakalowicz, 2001b: Inverse modeling of the hydrological and the hydrochemical behavior of hydrosystems: Char acterization of karst system functioning.W ater Re sources Research, 37, 8, 2191. Sutton, M., 2004: e pink planarians of Devils Icebox Cave Census Protocols. Cave Research Founda tion, 32. Vandike, J., 1983: Stream gaging results, upper Bonne Femme Creek rating curve, Devils Ice Box Spring, Memorandum. Division of Geology and Land Sur vey, Missouri Department of Natural resources, 1. Vandike, J. & S. Schulte, 1984: Devil's Icebox W ater Q ual ity Study June 1982 to July 1984, Memorandum, 43. Vesper, D.J. & W .B. W hite, 2003: Metal transport to karst springs during storm ow; an example from Fort Campbell, Kentucky/Tennessee, USA.Journal of Hydrology, 276, 1, 20. Vineyard, J., 1958: e reservoir theory of springow.National Speleological Bulletin, 48, 43. W hite, W .B., 1988: Geomorphology and Hydrology of Karst Terrains. Oxford University Press, Inc., 464. W hite, W .B., 2002: Karst hydrology: recent developments and open questions.Engineering Geology, 65, 2, 85. Doi: 10.1016/s0013(01)00116. W hite, W .B., 2007: Brief history of karst hydrogeology: Contributions of the NSS.Journal of Cave and Karst Studies, 69, 1, 13. W hite, W .B. & G.H. Deike, III, 1989: Hydraulic geom etry of cave passages.In: W hite, W .B. &W hite, E.L. (eds.), Karst Hydrology: Concepts from the Mammoth Cave Area, Van Nostrand Reinhold, pp. 223. New York. W icks, C. & D. Loper, 2008: A linear model of pressure pulses in karstic aquifers.AGU Spring Meeting Abstracts, 1, 05. W icks, C., Noltie, D.B., Peterson, E.W & T. Dogwiler, 2010: Disturbances in the habitat of Macrocotyla glandulosa (Kenk).Ecohydrology, 3, 1, 116. W icks, C.M., 1997a: A Hydrologic and Geochemical Model of the Devils Icebox.Missouri Speleology, 39, 3, 1. W icks, C.M., 1997b: Origins of groundwater in a u viokarst basin Bonne Femme Basin in central Mis souri, USA.Hydrogeology Journal, 5, 89. W icks, C.M. & B. Bohm, 2000: Application of unit hy drograph technique to the discharge record at Big Spring, Carter County, Missouri.In: Sasowsky, I.D. & C.M. W icks, (eds.), Groundwater ow and contaminant transport in carbonate aquifers, A.A. Balkema Publisher, pp. 31. W icks, C.M. & J.F. Engeln, 1997: Geochemical evolution of a karst stream in Devils Icebox Cave, Missouri, USA.Journal of Hydrology, 198, 1, 30. CAROL M. WICKS



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P HYSICAL STRUCTURE OF THE E PIKARST F IZI NA STRUKTURA EPIKRASA W illiam K. J ONES 1 Izvleek UDK 551.44:556.3 William K. Jones: Fizina struktura epikrasa Epikras je preperela vrhnja cona krasa s poveano poroznostjo, znailna za tevilne krake pokrajine. Je zgornja meja krasa in hkrati prostor v katerem se kopiijo organske snovi, ki reagirajo s pronicajoo vodo. Epikras skladii in usmerja preniklo vodo v kraki sistem. Hidravlina prepustnost epikra sa pada z globino. Epikras lahko deluje kot visei vodonosnik, v katerem voda lateralno potuje do vertikalnih struktur (razpoke,brezna), skozi katere hitro ali poasi odteka v globi no. Epikras nastaja zaradi razbremenilnih napetosti v tleh ter zinega in keminega preperevanje. Epikras se razvije tudi na svee izdanjenih karbo natih, vendar je razvoj zaradi dodatnega CO 2 hitreji na obmojih, pokritih s prstjo. Nakopiena prst v razpokah je lahko skladie organskih snovi in povzroa zastajanje pronicajoe vode. Debelina epikrasa je navadno do 15 m, eprav lahko, npr. v kamnitih gozdovih, vertikalno preperevanje razpok see precej globlje. Vrtae so hidroloko lahko del epikrasa, lahko pa predstavljajo vrzeli, skozi katere voda neposredno odteka v kras. Voda iz epikrasa hitro (skozi brezna in razirjene razpo ke) ali poasi (skozi sisteme ozkih razpok) odteka v kras, podvrena pa je tudi evapotranpiraciji. Veji del vode, ki ga padavinski dogodki iztisnejo iz epikrasa, je stara skladiena voda, ki jo nadomesti svea voda padavi nskega dogodka. Kljune besede: epikras, polnjenje vodonosnika, vadozna cona, kras. 1 Karst W aters Institute, PO Box 356, W arm Springs, Virginia 24484, USA, e-mail: W kj30@hotmail.com Received/Prejeto: 23.2.2013 COBISS: 1.01 ACTA CARSOLOGICA 42/2-3, 311, POSTOJNA 2013 Abstract UDC 551.44:556.3 William K. Jones: Physical Structure of the Epikarst Epikarst is a weathered zone of enhanced porosity on or near the surface or at the soil/bedrock contact of many karst land scapes. e epikarst is essentially the upper boundary of a karst system but is also a reaction chamber where many organics accumulate and react with the percolating water. e epikarst stores and directs percolating recharge waters to the underly ing karst aquifers. Epikarst permeability decreases with depth below the surface. e epikarst may function as a perched aqui fer with a saturated zone that transmits water laterally for some distance until it drains slowly through fractures or rapidly at sha drains or dolines. Stress-release and physical weathering as well as chemical dissolution play a role in epikarst develop ment. Epikarst may be found on freshly exposed carbonates al though epikarst that develops below a soil cover should form at a faster rate due to increased carbon dioxide produced by vege tation. e accumulation of soil within the fractures may create plugs that retard the downward movement of percolating water and creates a reservoir rich in organic material. e thickness of the epikarst zone typically ranges from a few meters to 15 meters, but vertical weathering of joints may be much deeper and lead to a stone forest type of landscape. Some dolines are hydrologically connected directly to the epikarst while other dolines may drain more directly to the deeper conduit aqui fer and represent a hole in the epikarst. W ater stored in the epikarst may be lost to evapotranspiration, move rapidly down vertical shas or larger joints, or drain out slowly through the soil inllings and small fractures. Much of the water pushed from the epikarst during storms is older water from storage that is displaced by the new event water. Keywords : epikarst, recharge, vadose zone, karst.

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ACTA CARSOLOGICA 42/2-3 2013 312 Epikarst is studied by biologists, hydrologists and geo morphologists. e presence of a near-surface habitat that retained at least some water year round has long been recognized by biospeleologists (Rouch 1968). A brief pa per by Mangin (1973) proposed the term epikarst for the perched aquifer in the upper part of the vadose zone. W illiams (1983) called this the subcutaneous zone and described the structure based on geomorpological ober svations. W illiams viewed the epikarst as a zone with many open fractures that facilitates rapid inltration of precipitation but creates a bottleneck for water move ment at some depth as the frequency and size of the frac tures decrease. An introduction to epikarst is presented in Ford and W illiams (2007). A denition of the epikarst from a 2003 Karst W aters Institute conference and work shop on the subject (Jones et al. 2004) is: Epikarst is located within the vadose zone and is dened as the heterogeneous interface between unconsoli dated material, including soil, regolith, sediment and vege tative debris, and solutionally altered carbonate carbonate rock that is partially saturated with water and capable of delaying or storing and locally rerouting vertical inltra tion to the deeper regional, phreatic zone of the underlying karst aquifer. A number of interesting studies have been pub lished since the 2003 conference and this paper presents a brief review. e literature now contains more quanti tative studies based on multiple year data collections and some old questions have been addressed but some new questions have arisen. A few books directly deal with the epikarst and a number of edited collections have sections on the subject. Biological work on epikarst fauna (especially Co pepods) includes a book by Pipan (2005) and numerous papers. A summary of research on epikarst communities was presented by Culver et al. (2012). A paper on Co pepod distribution as an indicator of epikarst system connectivity should be of interest to hydrogeologists as well as biologists (Pipan & Culver 2007). An introduc tion to the hydrology of the epikarst zone was presented by Bakalowicz (2012). A book by Kogovsek (2010) oers a very detailed study of several sites in Slovenia with data collected over several years. Another study from Slove nia by Trcek (2003) presents an interesting analysis of spring hydrographs. I NTRODUCTION T HE N ATURE OF THE E PIKARST Z ONE e epikarst typically extends from exposed carbonate bedrock on the land surface or from the soil/bedrock contact downward for several to tens of meters. e epikarst may contain a perched aquifer but is situated in the vadose zone (Fig. 1). e permeability of the epikarst decreases with depth and creates a sort of bottleneck for vertically percolating water. e epikarst has very fast and very slow ow routes so water inltrating from a giv en storm event may be partitioned along paths of widely ranging permeabilities. Flow routes through the epikarst may be very fast by ow directly down sha drains to underlying stream conduits and essentially bypass the smaller fractures with little mixing or storage within the epikarst. Some of the inltrating water may be stored for a period of several years and may also migrate some dis tance horizontally within the epikarst zone. W ater emerging from a karst spring is generally a mixture of autogenic and allogenic recharge to the karst aquifer system. Several studies have attempted to iden tify the proportion of epikarst water to total spring dis charge. is is obviously very site specic for a karst basin that receives a high percentage of recharge from sinking streams may have a much lower percentage of epikarst water in the water discharging through a cave stream or spring. Trcek (2003) and Trcek and Krothe (2004) stud ied the Hubelj Spring (Slovenia) and the Orangeville Rise (USA) using oxygen isotopes as a tracer to aid in separating the components of storm hydrographs. e studies found that there was a signicant piston ow eect with new event water moving into storage and dis placing older stored water from the epikarst. e aver age event water was about 22% of the storm discharge at the springs and the epiow component was 41% at the Hubelj Sping and 59% and the Orangeville Rise. A more direct sampling of epikarst water involves collecting water at ceiling drips or rimstone pools within a cave. A multiyear study of the epikarst above Postojna Cave in Slovenia involved sampling several ceiling drips (trickles) in the roof of the cave to collect ow rates, chemical characteristics and sampling for injected uo rescent tracers (Kogovsek 2010). e vadose zone is 100 meters (depth of cave passage) beneath the surface at the WILLIAM K J ONES

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ACTA CARSOLOGICA 42/2-3 2013 313 sampling sites. Kogovsek found that the trickles were very responsive to individual storm events with 72% of the annual precipitation inltrating through the epikarst zone. Storage of event water was very high at this study site and most of the storm response water pushed out at the trickles was older water being displaced by the event water. W hen the antecedent moisture was low prior to a storm event the response at the trickles was delayed until sucient water had moved into the epikarst fractures to ush old water from storage. Monitoring of the trickles following injection of dye on the land surface above the sampling sites resulted in a general pattern of one to three days for the rst appearance of the tracer (depending on ow response at the trickle) but suggested that 98% of the injected mass of the tracer was stored in the poorly Fig. 1: Sketch illustrating water movement through the vadose zone. Inltrating water enters the epikarst at the soil/bedrock contact and may be stored in fractures and transmitted horizontal for some distance to vertical drains. Water may remain in storage in this zone for several years and be pushed out (displaced) by new storm event water. Some of the vertical drains or shas may essentially bypass the epikarst and rapidly deliver event water to the phreatic zone. saturated catchment area within the epikarst. Tracers in jected by spreading on the soil took much longer (over three years) to reach a maximum concentration at the trickles compared to about ve months for the tracer in jected onto bedrock. Dye was still being recovered seven years aer injection for one of the tracer tests. Epikarst probably develops to some extent on most karst landscapes except perhaps in very arid regions. However, Kresic (2013) argues that epikarst, at least the usual description as presented here, is actually the exception and true epikarst is relatively rare. Epikarst developed on very young (eogenetic) carbonates are in uenced by the high primary porosity (Taborosi et al. 2004), but the karren forms become more like classical continental epikarst as the limestone matures (Mylroie P HYSICAL STRUCTURE OF THE E PIKARST

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ACTA CARSOLOGICA 42/2-3 2013 314 REFERENCES Bakalowicz, M., 2012: Epikarst.In: W hite, W & D. Culver (eds.), Encyclopedia of caves .Elsevier, pp. 284, Oxford. Culver, D., Brancelj, A. & T. Pipan, 2012: Epikarst com munities.In: W hite, W & D. Culver (eds.), Ency clopedia of caves. Elsevier, pp. 288, Oxford. Ford, D. & P. W illiams, 2007: Karst hydrogeology and geomorphology.John W iley and Sons, pp. 132, W est Sussex. Gins, A., Knez, M., Slabe, T. & W Dreybrodt (Eds.), 2009: Karst rock features Karren sculpturing.Carso logica 9, pp. 561, Ljubljana. Jocson, J., Jenson, J. & D. Contractor, 2002: Recharge and aquifer responses: Northern Guam Lens Aquifer, Guam, Mariana Islands.Journal of Hydrology, 260, 231. Jones, W ., Culver, D. & J. Herman (Eds.), 2004: Epikarst .Karst W aters Institute Special Publication 9, pp. 160, Charles Town. Kogovsek, J., 2010: Characteristics of percolation through the karst vadose zone.Carsologica 10, pp. 168, Lu bljana. Kresic, N., 2013: Water in karst.McGraw-Hill, pp. 183, New York. Mangin, A., 1973: Sur la dynamique des transferts en aquifere karstique.In: Proceedings of the 6 th Inter national Congress of Speleology, III, 157, Olo mouc. Mylroie, J., Jenson, J., Miklavic, B. & D. Taborosi, 2012: Surface and vadose implications of karstication in eogenetic carbonates.Geological Society of Amer ica Abstracts with Programs, 44, 7, p 434. Pipan, T., 2005: Epikarst A promising habitat. Carso logica 5, pp.101, Ljubljana. Pipan, T. & D. Culver, 2007: Epikarst communities: bio diversity hotspots and potential water tracers.En vironmental Geology, 53, 265. Rouch, R., 1968: Contribution a la connaissance des Harpacticides hypoges (Crustaces Copepodes).Annales de Speleology, 23, 1, 5. Stinson, C., Schwartz, B., Gerard, B., Schwinning, S., Ramirez, P., & G. Timmins, 2012: Preliminary re sults from a multi-tracer epikarst recharge experi ment: McCarty Cave, Texas, USA.Geological Soci ety of America Abstracts with Programs, 44, 7, 435. Trcek, B., 2003: Epikarst zone and karst aquifer behaviour A case study of the Hubelj catchment, Slovenia.Geoloski zavod Slovenije, pp.100, Ljubljana. Trcek, B. & N. Krothe, 2004: Oxygen isotope studies of major karst springs on the Mitchell plain (USA) and the Trnovski Gozd karst plateau (Slovenia).In: Jones, W ., Culver, D. & J. Herman (Eds.), 2004: Epikarst .Karst W aters Institute Special Publication 9, pp. 92, Charles Town. Taborosi, D., Jenson, J. & J. Mylroie, 2004: Karren fea tures in island karst: Guam, Mariana Islands.Z eit schri fur Geomorphologie, 48, pp. 369. W illiams, P., 1983: e role of the subcutaneous zone in karst hydrology.Journal of Hydrology, 61, 45. et al. 2012). A study of the partitioning of percolation water on Guam (Jocson et al. 2002) showed a similar range of travel times (hours to months as found in older karst settings. A study of the Edwards aquifer in Texas (USA) suggests a potential of signicant storage within the bedrock matrix (Stinson et al. 2012) as well as the fractures. Epikarst studies are certainly still in fashion and questions remain concerning the nature of the epikarst in dierent settings. Is there epikarst in regions where caves are formed primarily by hypogenic processes? How does glaciation aect the epikarst? Epikarst will be an important part of future research on the transport of carbon through the vadose zone. A CKNO WLEDGEMENTS Dr. Lee F. Elliott prepared the gure. anks are due to reviewers Drs. Branka Trcek and John Mylroie for sug gestions that improved the manuscript and added several important references. WILLIAM K J ONES


Description
Selected papers from an international and
multidisciplinary symposium on
Carbon and Boundaries in Karst, Carlsbad,
New Mexico, January 7 to January 11, 2013,
organised by Karst Waters Institute (KWI) and the
National Cave and Karst Research Institute (NCRKI).
Guest editor
David C. Culver,
Introduction to the symposium--
"From January 7 to January 11, 2013, the Karst waters
Institute (KwI) and the National Cave and Karst Research
Institute (NCRKI) held an international and multidisciplinary
symposium on Carbon and Boundaries in Karst at NCKRI
headquarters in Carlsbad, New Mexico.
There is growing interest in the dynamics of both
inorganic and organic carbon in karst systems, and especially
in the flux of carbon and nutrients between the surface and
subsurface, and between different components (e.g. epikarst
and vadose zone) in the karst subsurface. This symposium was
about these and other questions connected to carbon in karst
and boundaries in karst. It was especially timely both
because of rapid advances in the field and the importance of
carbon sequestration in global climate change The symposium
highlighted recent advances in biology, geology, and
hydrology that are helping us understand the dynamics of
karst ecosystems, especially with respect to carbon. The
talks were organized around seven main themes:
The Upper Boundary Epikarst
The Lower Boundary Phreatic Zone Lateral Inputs -
Insurgences
Lateral Outputs Resurgences CO
2- Processing and Storage
Organic Carbon Sources and Quality
Synthesis and Large Scale Models
Sixty participants from seven countries attended the
week-long meeting which included an excursion to Carlsbad
Caverns National Park. For the first time at a KwI meeting,
several participants, who were unable to attend in person,
gave their presentations via Skype. The meeting was
highlighted by two keynote presentations:
Groundwater Ecology of Alluvial River Flood plains,
Jack Stanford, Flathead Lake Biological Station, Polson,
Montana
Karst Conduit Matrix Exchange and the Karst hyporheic
zone, John wilson, New Mexico Institute of Mining and
Technoloogy, Socorro, New Mexico.
Two most distinguished karst scientists, William B. White
of Pennsylvania State University and Derek Ford of McMaster
University jointly summed up the meeting. ..."
EDITORIAL
Franci Gabrovsek
ORIGINAL PAPERS
Carbon fluxes in Karst aquifers: Sources, sinks, and the
effect of storm flow/
William B. White
Do carbonate karst terrains affect the global carbon
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Quaternary glacial cycles: Karst processes and the global CO2
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A framework for assessing the role of karst conduit
morphology, hydrology, and evolution in the transport and
storage of carbon and associated sediments/
George Veni
Biological Control on Acid Generation at the Conduit-Bedrock
Boundary in Submerged Caves: Quantification through
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Janet S. Herman, Alexandria G. Hounshell, Rima B.
Franklin, Aaron L. Mills
An approach for collection of nearfield groundwater samples
in submerged limestone caverns/
Aaron L. Mills, Terrence N. Tysall, Janet S.
Herman
Organic matter flux in the epikarst of the Dorvan karst,
France/
Kevin S. Simon
Environmental controls on organic matter production and
transport across surface-subsurface and geochemical
boundaries in the Edwards aquifer, Texas, USA/
Benjamin T. Hutchins, Benjamin F. Schwartz, Annette S.
Engel
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water to separate sources of recharge in a cave spring,
northwestern Arkansas, USA Blowing Spring Cave/
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Isotopes of Carbon in a Karst Aquifer of the Cumberland
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Lee J. Florea
Organic Carbon in Shallow Subterranean Habitats/
Tanja Pipan, David C. Culver
Contribution of non-troglobiotic terrestrial invertebrates to
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Physical Structure of the Epikarst/
William K. Jones
Using hydrogeochemical and ecohydrologic responses to
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plateau, Texas, USA/
Benjamin F. Schwartz, Susanne Schwinning, Brett Gerrard,
Kelly R. Kukowski, Chasity L. Stinson, Heather C.
Dammeyer
Variability of groundwater flow and transport processes in
karst under different hydrologic conditions/
Nataa Ravbar
Spring discharge records a case study/
Carol M. Wick


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