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January 2013 Hypogene Speleogenesis in the Cerna River Basin, SW Romania: A Sedimentological, Mineralogical, and Stable Isotopic Approach Cristina Montana Puscas University of South Florida, [email protected]
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Scholar Commons Citation Puscas, Cristina Montana, "Hypogene Speleogenesis in the Cerna River Basin, SW Romania: A Sedimentological, Mineralogical, and Stable Isotopic Approach" (2013). Graduate Theses and Dissertations. http://scholarcommons.usf.edu/etd/4931
This Dissertation is brought to you for free and open access by the Graduate School at Scholar Commons. It has been accepted for inclusion in Graduate Theses and Dissertations by an authorized administrator of Scholar Commons. For more information, please contact [email protected]. Hypogene Speleogenesis in the Cerna River Basin, SW Romania: A Sedimentological,
Mineralogical, and Stable Isotopic Approach
by
Cristina M. Pușcaș
A dissertation submitted in partial fulfillment of the requirements for the degree of Doctor of Philosophy School of Geosciences College of Arts and Sciences University of South Florida
Major Professor: Bogdan Petroniu Onac, Ph.D. Victor J. Polyak, Ph.D. Jeffrey Ryan, Ph.D. Len H. Vacher, Ph.D. Jonathan G. Wynn, Ph.D.
Date of Approval: November 26, 2013
Keywords: cave, morphology, sulfuric acid speleogenesis, thermomineral water, Băile Herculane
Copyright © 2013, Cristina M. Pușcaș DEDICATION
To my parents and Cosmin.
TABLE OF CONTENTS
List of Tables ...... iii
List of Figures ...... iv
Abstract ...... vi
Chapter One: Introduction ...... 1 Background ...... 1 Aim of the Dissertation ...... 5 Hypotheses to be Tested ...... 7 Hypothesis 1...... 7 Hypothesis 2...... 8 Hypothesis 3...... 8 Dissertation Outline ...... 8 Major Findings and Conclusions ...... 9 Chapter 2 ...... 9 Chapter 3 ...... 10 Chapter 4 ...... 11 Chapter 5 ...... 12 References ...... 13
Chapter Two: Past Surface Conditions and Speleogenesis as Inferred from Cave Sediments in the Great Cave of Șălitrari Mountain ...... 19 Introduction ...... 19 Location and Geologic Setting...... 20 Methods...... 23 Results ...... 25 Discussion and Conclusions ...... 29 References ...... 33
Chapter Three: The Mineral Assemblage of Caves within Șălitrari Mountain: Depositional Environment and Speleogenetic Implications ...... 40 Introduction ...... 40 Geographic and Geologic Setting ...... 42 Cave and Sample Description ...... 43 Analytical Methods ...... 46 Results ...... 47 Discussion and Conclusions ...... 51 Phosphates...... 52
i Nitrates ...... 52 Sulfates ...... 53 References ...... 54
Chapter Four: Tamarugite-bearing Paragenesis formed by Sulfate Acid Alteration in Diana Cave ...... 60 Introduction ...... 60 Local Geology and Hydrology...... 64 Tamarugite Occurrences and Stability ...... 65 Description of Samples and Analytical Methods ...... 66 Results ...... 67 Discussion ...... 70 Conclusions ...... 75 References ...... 76
Chapter Five: Using δ34S in low-grade hydrothermal waters to fingerprint hypogene speleogenesis (Băile Herculane, Romania) ...... 81 Introduction ...... 81 Site Description ...... 82 Hydrogeology and Hydrochemistry ...... 85 Description of the Sampled Springs ...... 86 Results ...... 91 Discussion ...... 97 Hydrochemistry...... 97 Origin of thermomineral water ...... 99 Stable sulfur isotopes ...... 100 Oxygen stable isotopes ...... 102 Dissolved inorganic carbon...... 104 Stable isotopes in cave minerals ...... 105 Conclusions ...... 106 References ...... 108
Appendix 1: Permission to reproduce article from Studia UBB Geologia ...... 116
Appendix 2: Permission to reproduce article from Springer ...... 118
Appendix 3: Permission to reproduce article from E. Schweizerbart'sche Verlagsbuchhandlung ...... 119
ii LIST OF TABLES
Table 3.1: List of minerals discussed in this study ...... 48
Table 3.2: EMPA results for alunite and aluminite from the Great Cave from Șălitrari Mountain ...... 51
Table 4.1: Comparison between the unit-cell parameters of tamarugite from Diana Cave, Romania and Mina de la Compania, Chile ...... 68
Table 4.2: List of cave mineral species identified in Diana Cave and associated formation processes ...... 68
Table 5.1: General information on the monitored thermomineral springs and wells ...... 88
Table 5.2: Summary of water chemistry for the thermomineral and karstic water springs/wells sampled for this study (May 2011-May 2012) ...... 93
Table 5.3: Results of stable isotopic measurements in water samples and resulting precipitates from all the sources (karstic and thermomineral) ...... 94
Table 5.4: Results of stable isotopic analysis of mineral samples from caves within the studied area, as well as gypsum deposits from the thermomineral springs ...... 94
iii LIST OF FIGURES
Figure 2.1: Plan and profile view of the cave, showing the location of the studied sections (red stars) ...... 22
Figure 2.2: Topographic map showing the catchment area of the Presacina Brook and the location of the GCSM ...... 24
Figure 2.3: Hjulström diagram with dashed line showing the flow velocity calculated in ScallopEx ...... 31
Figure 3.1: Topographic map of the studied area ...... 41
Figure 3.2: Map of SC2 and SC3 (modified from Ponta & Solomon, 1982) showing sampling locations...... 44
Figure 3.3: Map of the GCSM (surveyed by ‘‘Prusik’’ Speleo Club Timișoara) showing sampling locations ...... 45
Figure 3.4: Secondary electron image of pseudohexagonal platy taranakite crystals ...... 49
Figure 3.5: Secondary electron image of prismatic crystals of darapskite ...... 50
Figure 4.1: Location of the studied site ...... 61
Figure 4.2: Map of the Diana Cave (modified from Povară et al., 1972); water storage building not at scale ...... 62
Figure 4.3: Photograph of tamarugite-rich aggregates (in paragenesis with native sulfur, halotrichite-group minerals, and gypsum) ...... 63
Figure 4.4: Typical powder XRD pattern of a sample consisting of an admixture of tamarugite, anhydrite, alunite, and native sulfur ...... 69
Figure 4.5: SEM back-scattered electron images of tamarugite aggregates ...... 70
Figure 4.6: Schematic cross-section (not at scale) through Diana Cave ...... 72
Figure 5.1: Simplified geological map of the area (after Năstăseanu, 1980) ...... 84
iv Figure 5.2: Monthly variation in hydrochemical parameters in each of the monitored sources...... 96
Figure 5.3: Average chemical and physical parameters of the studied springs ...... 98
Figure 5.4: Comparison of δ18O and δD of thermomineral springs/wells to δ18O and δD of karstic spring waters ...... 100
Figure 5.5: Monthly variation in δ34S TSV of all sampled sources along a N – S transect ...... 102
18 34 Figure 5.6: Comparison of δ Osulfate and δ Ssulfate (monthly averages) on a N – S transect ...... 103
Figure 5.7: The relationship between δ34S and δ13C of the thermomineral sources, along a N – S transect ...... 105
v ABSTRACT
Ever since it was identified as a speleogenetic process in the Guadalupe Mountains of
New Mexico, USA, hypogene speleogenesis has become the focus of numerous research projects
aimed at discerning between classical epigene caves and sulfuric acid or thermal caves. The first
distinguishing characteristics that were recognized for hypogene caves were passage and cave
morphology. The following step was the identification of rare minerals, specific for processes
associated to hypogene speleogenesis. One other important step was the recognition of the
importance of stable isotopes – mainly of S – in tracing the source of S and the chemical
processes affecting it. Many of the caves now labeled as hypogene are fossil caves, in which
presently the hypogene activity has long died off. Studies comparing stable isotopes from coexisting cave minerals and the waters that generate the cave are rarer. This extensive study
encompasses a description of cave and passage morphologies, cave mineral assemblages, as well
as hydrogeochemistry of thermomineral waters in a peculiar region of Romania.
Băile Herculane (Cerna River Valley, SW Romania) is a spa town known since Roman
times for its numerous thermal springs that were considered to have healing powers. These
springs, along with wells drilled in the past century, are still being used for curative purposes in
several treatment centers in Băile Herculane. The present study is important not only for the
scientific data it produced, but also for economic purposes, as mixing of the thermomineral
waters with meteoric sources is a major concern, due to the dilution it causes.
vi The data presented here is based on multiple investigation methods, each specific to the
analyzed material: powder X-ray diffractions, scanning electron microscope, electron microprobe (for mineral samples), sedimentological investigations (for cave sediments), stable isotope mass spectrometry (for water and mineral samples), field measurements (for water samples).
The results presented here help to clarify the source of dissolved S species in the thermomineral water, the source of the water itself, as well as establish a connection between caves along the Cerna Valley and the thermomineral aquifers.
vii CHAPTER ONE:
INTRODUCTION
Background
Two main categories of dissolution caves have been described based on the source
(surface vs. depth) of acids in the solutions that generate the passages: epigene (surface-derived sources of dissolution) and hypogene (at least one source needed for dissolution derived from depth). The majority of known caves are believed to be epigene caves (Ford & Williams, 2007;
Palmer, 2007). The initial model for epigene cave development was centered on the gravitational seepage of meteoric water carrying biogenic CO2 as the agent of aggressiveness
(Ford & Williams, 2007; Palmer, 2007). In a second model of epigene speleogenesis, called
mixing corrosion, the source of aggressiveness arises from the amalgamation of two distinct
bodies of freshwater, each saturated with respect to calcium carbonate at their ambient pCO2 and temperature, resulting in a solution that is undersaturated under the new conditions (Bögli, 1980;
Gabrovśek & Dreybrodt, 2000). This study focuses on hypogene speleogenesis.
Passage development at or below the water table when at least one of the sources of aggressiveness originates at depth is known as hypogene speleogenesis. A first such model involves solution processes triggered by the rising thermal waters (mineralized or oligomineral).
Well-known examples of hypogene caves created by thermal water – in the so-called hydrothermal karst – are the caves from beneath Budapest (Hungary; Dublyansky, 1995). The second model is a special case of mixing corrosion, evidences along the shorelines of carbonate
1 islands where at halocline, the mixing between meteoric and sea water, generates the so-called
flank margin caves (Mylroie & Carew, 1990). The hypogene component(s) does not originate
from significant depth in this type of speleogenesis. Another model of low temperature
hydrothermal examples having hypogene gases/solutions coming from significant depth involves
rising CO2 and H2S (Klimchouk, 2007).
The history of the third and most important model of hypogene speleogenesis started in
New Mexico, during intensive research carried out in Carlsbad Cavern, Lechuguilla Cave, and
several other caves of the Guadalupe Mountains. Prior to the 1980s, cavers and scientists were
convinced that these were caves formed by the more common epigene model of carbonic acid
dissolution. In the last few decades of the 20th century, however, a new understanding of the speleogenetic processes that were at work here was developed by means of novel analytical
methods (e.g., Jagnow et al., 2000; Polyak & Provencio, 2001). A particular type of hypogene speleogenesis is sulfuric acid speleogenesis (SAS) (Hill, 2000); in this model the source of aggressiveness is provided by ascending H2S – derived from hydrocarbon reservoirs, sulfate
deposits, or sulfide ore bodies – that is oxidized to sulfuric acid, which will react with the
limestone bedrock dissolving it and creating voids (Hill, 1981; Van Everdingen et al., 1985),
following the reaction chain described by Palmer (2007):
+ 2- H2S + 2O2 = 2H + SO4 (1)
+ 2- 2+ 2- 2H + SO4 + CaCO3 = Ca + SO4 + H2O + CO2 (2) The identification of this type of speleogenesis gave way to a worldwide surge of
research in the field (caves from Mexico, France, Italy, Austria, Ukraine, Hungary, Romania,
etc.), connecting SAS to the presence of ore bodies, volcanic activity, and hydrocarbon deposits,
and often mediated by the presence of thermal waters (Hill, 2000; Galdenzi et al., 2008). SAS is
responsible for caves such as Aix-les-Bains (France; Audra et al., 2007), Cova de Pas de 2 Vallgornera (Spain; Fornós et al., 2011), Frasassi (Italy; Galdenzi & Menichetti, 1995), Diana
(Romania; Onac et al., 2013; Pușcaș et al., 2013), Cathedral (Australia; Osborne, 2007), Cueva
de Villa Luz (Mexico; Hose & Pisarowicz, 1999), etc.
In early literature, the determining factors for classifying a cave as hypogene were
restricted to specific cave patterns (maze, network, spongework, etc.) or passage morphologies
(ceiling cupolas, floor feeder tubes, etc.). It was soon recognized that geochemical indicators are
significantly more reliable and as a result, stable isotopes – oxygen (Dublyansky, 1995) and
particularly those of sulfur (S) (e.g., Van Everdingen et al., 1985; Hill, 1987) – became a widely
accepted tool in identifying SAS. Comparative 34S measurements for S-bearing cave minerals
and S source (gypsum or anhydrite in the limestone bedrock, disseminated pyrite, hydrocarbons,
water, etc.) can isolate the source of sulfuric acid that created the cave and deposited the
secondary sulfates. Isotope studies can also indicate the presence of biological activity that leads
to highly depleted 34S values. Hill (1981) and Kirkland (1982) used 34S measurements of
gypsum from Carlsbad Cavern to show that the source of S likely came from hydrocarbons
rather than bedrock. One of the earliest cave studies involving both S and oxygen (O) stable
isotopes is that of Van Everdingen et al. (1985). They measured and compared 34S and 18O from spring water, dissolved sulfates and sulfides, airborne sulfur species, and sulfate minerals from two caves developed in the Cave-and-Basin hot spring travertine deposit (Banff National
Park, Alberta, Canada). The main process they invoke is chemical corrosion, seconded by microbiological activity, through a complex suite of reactions, concluding that the two caves are created mainly through the corrosion of travertine by sulfuric acid aerosols, generated by oxidation of H2S escaped from spring water. This study was closely followed by those of Carol
Hill and her coworkers, all centered on caves from the Guadalupe Mountains, New Mexico (Hill,
3 2000). Bretz (1949) anticipated that deep-seated anhydrite beds from the Castile Formation were
the source of S for the massive gypsum deposits in the caves, while Jagnow (1977) proposed that
pyrite could also be a source for SAS in this region. By the mid 1980s, stable isotope studies
confirmed that the source for strongly depleted 34S in the gypsum deposits from caves other
than Carlsbad Cavern could have been neither marine anhydrite, nor disseminated pyrite from
the limestone bedrock, as they all display very positive values compared to the cave gypsum
(Hill, 1987; Pisarowicz, 1994; Polyak & Güven, 1996). The missing link seems to be the neighboring hydrocarbon reservoirs, which show the same depleted 34S as the cave sulfates
(Hill, 1987). Research based on cave morphology and mineralogy, regional geological history,
and stable isotope data, show that the caves of the Guadalupe Mountains were developed in
several distinct stages of evolution, the final one being SAS.
Another remarkable case is Cueva de Villa Luz, in Tobasco, Mexico. The interest in this
cave comes from its H2S-rich, milky-white colored underground stream that is responsible for
aggressively dissolving the limestone bedrock and forming the cave passages. Later on,
microbiologists developed an interest in the particular cave ecosystem (Hose & Pisarowicz,
1999; Hose et al., 2000). The cave stream mixes with infiltrating freshwater and as a result H2S
is oxidized to form sulfuric acid, which will convert limestone bedrock to gypsum crust. In this
cave, limestone on the cave floor is dissolved by sulfuric acid dripping from the gypsum crust
and “snottites” (mucus-like colonies of single-celled bacteria that hang from the walls and
ceilings of caves) above. Further dissolution of cave walls takes place by acid seepage through
the gypsum crust. In order to discern between several possibilities for the origin of S, Hose et al.
34 (2000) sampled H2S gas from the cave atmosphere as well as S from wall deposits (δ S -11.7‰ for the cave atmosphere, -23.7‰ for sulfur and -23.4‰ for gypsum). Stable isotope data for
4 34 sulfur and H2S of magmatic origin from El Chichón volcano was used for comparison (δ S
values of +4.6‰ and +4.5‰, respectively, ruling it out as the S source for SAS). Negative δ34S
values usually denote bacterially-mediated fractionation, hence the low values from Cueva de
Villa Luz may suggest an origin from bacterial reduction of sulfate. Conversely, no isotope data
for sulfate in the groundwater, nor from evaporite sediments in the area is presented, making further research necessary in order to confidently indicate the source for SAS.
A classic example of hypogene speleogenesis in Europe is the Frasassi cave system
(Italy). Among the peculiarities of this impressive underground system are the massive gypsum
deposits in the dry passages of the cave, and the numerous sulfidic and mildly thermal springs,
some of which are strongly affected by infiltrating meteoric water. The regional geology (mainly
limestones and marl) is complicated by faults and fractures (Galdenzi & Menichetti, 1995;
Galdenzi et al., 2008). Sulfur isotopic data from analyzed sulfate and sulfide samples from both cave deposits and spring water is presented by Galdenzi & Maruoka (2003). Their results show that the cave sulfates are not a result of direct precipitation from calcium sulfate supersaturated solutions. Gypsum 34S in the cave shows values that are either higher or lower than those of the
present-day H2S that rises in the groundwater. This suggests a change in the depositional setting
of sedimentary gypsum over a period of 200 ka.
Aim of the Dissertation
The first SAS cave identified in Romania was Movile, close to the Black Sea. It is a cave
that had no natural entrance for possibly hundreds of thousands of years, until intersected by a
mining shaft. Due to this isolation and deep-seated H2S-rich anoxic water sources, a unique
5 endemic fauna developed in the cave’s ponds (Sârbu et al., 1996). Hypogene caves were also
identified in the NW part of Romania in the Băița metallogenic district (Onac, 2002).
In the western part of the Southern Carpathians (Romania), the presence of numerous
thermomineral springs (identified at surface, in wells, and caves) with varying temperatures and
degrees of mineralization – many rich in H2S – and their association with a number of caves
raised questions about the region’s potentially hypogene speleogenesis (Sumrall, 2009; Onac et
al., 2009a,b; Pușcaș et al., 2010; Wynn et al., 2010; Onac et al., 2011; Pușcaș et al., 2013). The
Cerna River Valley is hosted on its entire 85 km length by a tectonic trough, with a highly
fragmented landscape and vertical displacement of <1000 m. The basement comprises
Precambrian and Cambrian crystalline schists, overlain by lower Paleozoic metamorphic and
igneous rocks, upper Paleozoic metasedimentary and igneous formations, on top of which lay
Mesozoic sediments (Năstăseanu, 1980). Jurassic to Lower Cretaceous limestones flanking the
Cerna Valley host over 100 caves. Karst-related studies are represented by cave surveys (Avram
et al., 1964, 1966; Dancău et al., 1968; Ponta & Solomon, 1982), fauna (Decou et al., 1974;
Decu & Tufescu, 1976), and mineralogical studies (Povară et al., 1972; Diaconu & Medeşan,
1973; Diaconu, 1974; Diaconu & Lascu, 1998-1999; Onac et al., 2009a,b; Onac et al., 2013).
Intense tectonic activity is translated into a wealth of thermal springs, with varying
2- concentrations of dissolved H2S and SO4 , considered as the probable source for the aggressive
solutions that carved numerous caves in the limestone walls. Sumrall (2009) presents 34S values
(in various S species and phases) from a large number of water and mineral samples. The sulfide
34S values range from -21.9‰ to 24.0‰, while sulfate 34S ranges from 16.6‰ to 71.3‰.
Sulfate minerals from several caves cover widely varying values (from -27.9‰ to +20.3‰), demonstrating that various processes are at work in the region. The Cerna cave sulfate δ34S
6 values are evidence for SAS being dependent not only on the S source, but also on the
completeness of sulfate reduction and sulfide oxidation, further demonstrating that a single range
of δ34S values does not necessarily pinpoint SAS processes. The aim of this dissertation is to better constrain the speleogenetic evolution of this region by expanding on the number of springs investigated in earlier research, as well as the methods implemented so far. Identification of cave-mineral assemblages and sedimentological analyses are complemented by a one-year long geochemical study – as a monthly survey – of 19 water sources (both thermomineral and karstic).
Hypotheses to be Tested
Hypothesis 1
The first working hypothesis is that initial epigene phreatic conditions were followed by a short-lived hypogene stage, and a return to vadose conditions. In such conditions, after the development of deep phreatic conduits the caves continued their evolution under vadose conditions following a lowering of the local base level, leading to the accumulation of alluvial sediments, and precipitation of common secondary cave minerals (carbonates, sulfates, and phosphates). Upon further deepening of the regional base level, water of meteoric origin reached the deep faults from where it resurged as heated and mineralized gases/solutions. The result was a SAS flare-up, triggered by the oxidation of hydrogen sulfide (generated by sulfate reduction) and mediated by thermal waters. After the hypogene activity died off, cave evolution under vadose conditions partly overprinted the SAS-related cave morphologies (corrosion niches, ceiling cupolas etc.). If this stands true, along with cave patterns and passage morphologies that are typical for SAS, specific minerals (especially sulfates that precipitate and remain stable under particular pH, relative humidity, and temperature conditions) are considered to be important
7 indicators of hypogene speleogenesis (Polyak & Güven, 1996) and should be present in these caves. The isotopic composition of dissolved sulfur species in the thermomineral water and sulfate minerals in the caves should allow us to retrace the evolution of the system.
Hypothesis 2
The second hypothesis takes into consideration the possibility of a purely hypogene speleogenesis (either SAS or ascending thermal water/steam), at least along the river transect that is affected by thermal water. Thus, all features (morphological and mineralogical) would be linked to a deep-seated aggressive solution. It should be possible to test this hypothesis especially in the case of caves that are at or close to the present day base level and currently affected by thermal water and/or steam, by documenting water and cave mineral geochemistry.
Hypothesis 3
At least some of the caves are purely epigenic and the result of dissolution by surface- generated CO2, thus unaffected by the circulation of thermo-mineral fluids. This would translate into conventional epiphreatic cave morphologies and common minerals (carbonates, phosphates, and sulfates). The stable isotopic signatures in such minerals would be distinctive for near- surface processes.
Dissertation Outline
The work presented in this doctoral dissertation is divided into five major chapters; the first chapter is an introduction to the dissertation, presenting background information on the topic of my research. The following three chapters are already published as individual articles in
8 scientific journals, and the fourth is to be submitted. Although each chapter was conceived as a standalone body of work and handles a specific topic, they have a common thread and are in fact pieces of the same puzzle. In order to fingerprint the speleogenetic processes at work along the
Cerna River we decided to investigate several facets. Passage morphology and cave sediments are relatively easy observable features, but beyond these, investigating cave minerals and speleothem geochemistry are also important for the characterization of different speleogenetic pathways. Dividing these and examining each in a separate chapter allowed us to generate a more accurate model for cave development and subsequent passage transformations in the peculiar karst hosted by the Cerna Valley.
Major Findings and Conclusions
Chapter Two
The second chapter is an analysis of passage morphology and sediments in the longest cave of the Cerna Valley – Peștera Mare din Muntele Șălitrari – and their relationship to speleogenesis (Pușcaș et al., 2010b). The premises for this study were set by the existence near the cave entrance of a >6 m thick alluvial sediment deposit, with grain sizes ranging from boulders to clay and covered by a blanket of dry guano. Sediment and mineral samples collected from six profiles underwent broad analyses to determine their petrological and mineralogical makeup, grain-size distribution, and paleoclimatic significance. The complicated facies alternation suggests frequent changes in the former stream’s hydrological parameters, with frequent flooding, leading to the hypothesis that the climate was somewhat wetter than today.
Both the mineralogical composition of the sediment (ranging from quartz, mica, gypsum, phosphates, and calcite to garnet, zircon, titanite, olivine, serpentine, tourmaline, sphalerite,
9 pyrite/chalcopyrite, and feldspars) and the petrological composition of the larger clasts
(limestone, sandstone, mudstone, granitoids, serpentinite, amphibolite, diorite, gneiss, quartzite, microconglomerate, and schist) ascribe the potential source rocks to an area with contrasting lithologies, such as amphibolites, felsic and basic metaigneous, and metasedimentary rocks,
mixed with a variety of detrital rocks. These rock types are not entirely comprised by the
catchment area of the modern Presacina Brook, thus implying that due either to hydrological
conditions, or to changes in the base level caused by river down cutting or active tectonics, the
former source area was much more extensive. Based on morphological and sedimentological
criteria, the cave started under pipe-full flow conditions (phreatic), and later evolved during a
prolonged and complex vadose phase. Once the underground stream left the cave and most of the
sediment was removed, speleothem precipitation was initiated. In this contribution we put forward evidence that argues for an extra-basin origin of some of the alluvial sediments, as well as extract information about recent paleoclimate changes in the region. Most importantly, evidence to support the existence of a hypogene stage is present in this chapter, in the form of passage morphology and mineralogical assemblage.
Chapter Three
The third chapter is an in-depth investigation of the mineralogy of speleothems and cave deposits in the Șălitrari Mountain cavities. Eighteen minerals belonging to eight chemical groups were identified from three caves within Şălitrari Mountain, in the upper Cerna River basin
(Romania) by means of scanning electron microscopy, electron microprobe analysis, and X-ray
powder diffraction (Pușcaș et al., 2010a). One passage in the Great Cave from Şălitrari
Mountain, the largest cave investigated, exhibits abnormal relative humidity and temperature
10 ranges, allowing for a particular depositional environment. The cave floor is covered by alluvial
sediments (ranging from cobble, sand, and clay to silt and clay-sized material), bear bones, bat
guano, and rubble. These materials reacted with percolating meteoric water and hydrogen
sulfide-rich hypogene hot solutions, precipitating a variety of secondary minerals. Most of these
minerals are common in caves (e.g., calcite, gypsum, brushite), however some of them (e.g.,
alunite, aluminite, nitratine, darapskite, and variscite) require very particular environments in
order to form and persist. Cave passage morphologies suggest a complex speleogenetic history
that includes changes from phreatic to vadose conditions. The latter was punctuated by a sulfuric
acid dissolution/precipitation phase, partly overprinted by present day vadose processes. The
cave morphology and the secondary minerals associated with the alluvial sediments in these
caves are used to unravel the region’s speleogenetic history.
Chapter Four
Continuing the topic from the previous section, this chapter presents an exceptional new
occurrence of the mineral tamarugite, NaAl(SO4)2·6H2O, from a short karst cavity (Diana Cave,
Băile Herculane; SW Romania) described by Pușcaș et al. (2013). It was formed by corrosion of
2- 2- the bedrock (limestone and marls) by a SO4 -rich steam condensate resulting from oxidized S ions escaping from the thermo-mineral water emerging from depth in the cave. Tamarugite forms dull white earthy aggregates. Scanning-electron microscope (SEM) observations reveal tabular subhedral crystals never exceeding 15 μm across. The cell parameters refined from the powder data for the monoclinic space group P21/a are: a = 7.358(6), b = 25.23(2), c = 6.093(5) Å, β =
95.16(5)º, V = 1126.98(1) Å3. The δ34S values of the cave sulfates and the thermal water confirm
marine evaporites as the source of sulfur. The sulfate-acid alteration of limestone with
11 contribution of Al3+ and Na+ from the marls and the thermal water is responsible for the
formation of tamarugite. The steam-condensate alteration paragenesis includes native sulfur,
bassanite, anhydrite, epsomite, pickeringite, halotrichite, apjohnite, and alunite, as well as quartz
and halite, all primary and secondary speleogenetic by-products (Onac et al., 2009a; Pușcaș et al.,
2013).
Chapter Five
The previous chapters show that passage morphology, along with cave sediments and minerals in the Cerna Valley karst support the existence of a hypogene phase in the evolution of these underground voids. A unique attribute of this area is that in some of the lower Cerna basin caves, steam and thermomineral water are still present. The logical next step is to look into the physical and chemical properties of the region’s water sources in order to check if they are consistent with hypogene speleogenesis (SAS or otherwise). Hydrogeochemistry of naturally occurring thermomineral water is used to examine S sources and cycling, and to help distinguish between sulfate reduction pathways. The results are used to fingerprint different speleogenetic pathways in and sulfate mineral formation the surroundings of Băile Herculane, SW Romania.
On-site measurements of chemical and physical parameters (temperature, pH, salinity, total
2- 2- dissolved solids, electric conductivity, redox potential, S and SO4 concentrations) were carried out between May 2011 and May 2012 and confirm previously established grouping of the thermomineral sources (3 aquifers consisting of 5 groups). δ34S values of underground and surface water sources varied between 13.7 and 26.6 ‰ (1 σ) for dissolved sulfide (S2-, HS-, and
2- 18 H2S) and between 2.4 and 71.9 ‰ for dissolved sulfate (SO4 ). δ O values of water ranged from
-11.4 to -8.4 ‰, δD between -80.2 and -68.5 ‰, while δ13C in DIC varied between -34.4 and -
12 7.1 ‰. The δ18O and δD values of the thermomineral water support their meteoric origin, as
proposed by the hydrogeological model proposed by earlier research. Given the high geothermal
gradient, thermochemical reduction is the most likely process for the sulfate in this region, and
based on calculated total sulfur value the source of S is marine evaporites associated with the
limestone bedrock. δ13C values point toward methane (derived from a neighboring coal deposit) as the likely electron donor for sulfate reduction.
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18 CHAPTER TWO:
PAST SURFACE CONDITIONS AND SPELEOGENESIS AS INFERRED FROM CAVE
SEDIMENTS IN THE GREAT CAVE OF ȘĂLITRARI MOUNTAIN
Note to Reader
This chapter has been previously published in Studia UBB Geologia, 2010, 55: 51-57,
and has been reproduced with permission from Studia UBB Geologia. The author of this
dissertation is the first author of the article and the major contributor.
Introduction
The importance of cave sediment studies has been increasingly acknowledged over the
last few decades, especially with the augmented reliability and availability of a handful of
methods that make dating them possible. Their merit resides in their applicability to a variety of
connected research fields: anthropology and archaeology (e.g., Berger et al., 2008), paleoclimatology (e.g., Audra et al., 2007; Polk et al., 2007), paleoenvironment and paleotopography reconstructions, mineralogy (e.g., Polyak & Güven, 2000; Onac et al., 2007), sedimentology and speleogenesis (e.g., Horoi, 1993; Roată, 1993; Häuselmann et al., 2010), as well as several others (e.g., Sasowsky, 2007).
For almost a century the Cerna Valley represented the locus of interest for numerous investigations, initially triggered by the abundant thermal springs and related geological challenges (Povară et al., 1972, 2008; Veliciu et al., 1983; Cosma et al., 1996; etc.). This study
19 continues a series of research papers (Onac et al., 2009a, b; Puşcaş et al., 2010; Wynn et al.,
2010) documenting various aspects of caves from SW Romania, namely the Cerna River Valley.
Our aim is to bring into light new data concerning the caves developed in the massive Upper
Jurassic limestone deposits that flank the Cerna River a few kilometers upstream of the historical spa of Băile Herculane. The purpose of our work is to better constrain the evolution of cave passageways, along with the mode of sediment deposition (directly connected to the direction of the paleoflow and the source area) by combining sediment analysis with observations of passage and cave morphology in the Peștera Mare din Muntele Șălitrari (Great Cave of Șălitrari
Mountain, hereafter GCSM). Sub horizontal passages (common to all caves in the proximity) are symptomatic of caves formed near the base level, meander shortcuts making up suitable traps for alluvial sediments (Quinif, 2006). Alternating lithologies are clues for variations in the hydrological regime, which is directly related to changes in surface climate and topography
(Häuselmann et al., 2008).
Our results constitute a starting point for future studies aimed at dating the cave alluvium and deciphering the paleoenvironmental context of their deposition. For a detailed discussion regarding the mineralogical assemblage and peculiarities of GCSM, the reader is referred to
Diaconu & Lascu (1998), Onac et al. (2009a), Pușcaș et al. (2010).
Location and Geologic Setting
The Cerna Mountains are a geographic unit pertaining to the South Carpathians of
Romania and are characterized by a series of ridges frequently up to 1500 m in altitude, deeply cut by valleys, resulting in a steep relief especially where limestone outcrops. Oriented NE-SW, the Cerna Valley follows a major dextral strike-slip fault (Berza & Drăgănescu, 1988; Kräutner
20 & Krstič, 2002) through its entire 85 km length, locally with canyon walls carved in granite,
schist, and limestone climbing up to 500 m almost vertically (Alexandru et al., 1981). Several smaller strikeslip (both dextral and sinistral) faults striking approximately WNW-ESE, cut the lithologies along the Cerna Valley, while granitic intrusions mapped along the valley are characterized by roughly vertical foliations (Fig. 2.1, inset B). A subdivision of the Cerna
Mountains, Mount Șălitrari flanks the Presacina Brook on the right at its confluence with Cerna, ca. 15 km upstream of Băile Herculane (Fig. 2.1, inset A; Fig. 2.2). Along the Presacina Brook, the intrusion of a small granitoid pluton (and its apophyses) disrupted the original strikes and dips of the beds. However, in the cave area the limestone beds still preserve a roughly N-S strike, while dipping ca. 15° to the E.
The major components of the geological puzzle in the research area are the sedimentary formations belonging to the Alpine Danubian nappe system and the metaigneous formations of
the Alpine Getic-Supragetic nappe system. The Neoproterozoic Neamțu metamorphic series,
outcropping in the SW part of the valley (west of the Băile Herculane Spa) consists of biotite-
bearing migmatitic gneiss, amphibolites, and quartz-bearing micaschists, retromorphosed to
green schist facies. Outcropping in the riverbed of both Presacina and Cerna at their confluence and cross cutting the eastern side of Neamțu Series, the Cerna Granite is comprised of quartz, plagioclase, orthoclase, and various amounts of biotite; its predominant color is gray, but at this particular juncture it is bright red (due to its high orthoclase content). Discordant Jurassic detritic deposits in the area of interest belong to the Presacina facies (conglomerates and sandstone at the base, followed by clays and younger sandstone, topped by limestone) pertaining to the Presacina
Sedimentary Zone (Năstăsteanu, 1980).
21 Figure 2.1. Plan and profile view of the cave, showing the location of the studied sections (red stars). Inset A: Location of Băile Herculane; Inset B: Simplified geologic map and location of GCSM (black star) (modified from Năstăseanu and Bercia, 1968); 1, 2: Upper Anteproterozoic (Sebeș-Lotru Series, Amphibolite Facies); 3: Permian (conglomerates, red clays, sandstones, schists); 4: Paleozoic (granitoids); 5: Mesozoic (ultrabasic rocks); 6: Permian (volcanic-sedimentary formations); 7: Lower Jurassic (conglomerates, clayey schists, sandstones); 8: Upper Jurassic-Aptian (Azuga, Sinaia, and Comarnic strata); 9: Turonian-Senonian (sandstones, conglomerates); 10: Albian-Cenomanian (sandstones, limestones, clays); 11: Tithonian-Aptian (limestones, sandstones, conglomerates, dolerites); 12: Upper Cretaceous (sandstones, conglomerates); the dashed blue frame approximately shows the area presented in Fig. 2.2.
22 A geological cross-section through the Presacina Brook reveals the Bogîltin
(Pliensbachian-Toarcian; comprised of fine to coarse sandstones with alternating levels of conglomerates, microconglomerates and dark clays), Ohaba (same age as Bogîltin; almost exclusively black clayey schists with sandstone lenses), Ciumărna (Aalenian; represented by fine-grained to coarse, hard, massive, quartzitic sandstone), and Iuta Beds (Barremian-Aptian; correspond to limestone grading into black schistous micabearing marly limestones, capped by conglomerate), followed by the wildflysch facies (Upper Cretaceous) and Arjana Beds
(conglomerates, coarse sandstone, calcareous sandstone, clayey sandstone, often with schistose structure) (Năstăseanu, 1980). In the modern catchment area of Presacina Brook none of these sedimentary successions incorporate mafic and ultramafic lithologies such as amphibolites and serpentinites.
Methods
Sediment and mineral samples were collected from four out of the six investigated profiles; thorough lithological columns were drawn at the time of sampling. Profiles were labeled
S (Șălitrari) followed by consecutive numbers, starting from the mid-section of the NP and continuing toward the cave entrance. Samples were labeled PM followed by a number, in a continuous manner. All sediment levels in Profile S1 and S2 were sampled, while in the other 4 profiles samples were taken only from levels that did not correspond to what we encountered in the previous sequences. Samples were oven dried at 55°C and color coded using a Munsell Soil
Color Chart, weighed and sieved for an average of 5 minutes using standard brass sieves and a
W.S. Tyler Rotap RX-29 shaker.
23 Figure 2.2. Topographic map showing the catchment area of the Presacina Brook and the location of the GCSM. Contour interval 30 m. Topographic map from Jarvis et al., 2008. Hydrology modified from data available at Domogled-Valea Cernei National Park (www.domogled-cerna.ro).
The sediment on each sieve was weighed to an accuracy of ±1 g. Separates from the sand sized fraction (φ 2) were washed in DI water to remove clay, and a Franz Magnetic Barrier
Laboratory Separator LB-1 was employed to separate zircon and titanite crystals. Following this treatment, mineral grains were identified and quantified by means of stereoscopic microscopy.
The φ -4 and larger fraction was cut into halves to allow for a better petrological identification of the clasts, in order to ease the identification of the sediment source area. Lithologic columns were created using the software package POLITO 0.3.1 (Stremțan & Tudor, 2010).
Ten samples from the finest material (φ >4) underwent powder X-ray diffraction (XRD) in the X-ray Powder Diffraction and Thermoanalytical Laboratory (University of Miskolc), to
24 further identify sediment mineralogy. Diffraction spectra were obtained on a Bruker D8 Advance diffractometer under the following operating conditions: CuK α1 radiation (1.54056; secondary graphite monochromator, Bragg- Brentano geometry, step-scanning mode, 2º-65º (2θ) range,
0.04º (2θ) step size, and 2 sec/step detecting time (0.6 mm anti-scatter and receiving slit, 0.2 mm detector slit). Identification of components was carried out by Search/Match procedure of Bruker
DiffracPlus EVA evaluation module, based on PDF2 (2005) database.
Measurements of cave wall scallops, passage cross-section, and temperature were input into SCALLOPEX 1.0.1, software intended to facilitate the extrapolation of cave stream paleovelocities (Woodward & Sasowsky, 2009). These features are a proxy for flow velocity, which can be inferred from their length, as well as for flow direction, indicated by the positioning of their steep side. Water viscosity and density, dependent on temperature, as well as passage cross section are the main parameters affecting flow velocity, and thus the resulting scallops.
Results
The succession of the studied profiles heads towards the entrance and their positioning in the cross-section of the cave passage varies. Even though placed within some few meters from each other (Fig. 2.1), the sites display variations in sedimentary facies. While the overall lithological composition is fairly monotonous, the alternation of sediment sequences is indeed complex. None of the profiles contain drip- or flowstone (although a stalagmitic crust frequently covers the sediment deposit), but fragments of mammal bones were occasionally found. Plant material was also absent from the alluvial deposit.
25 Stratigraphic sections are described here from their top to the bottom. Although removed in places, the top layers of the sediment fill from GCSM (outcrops S1 and S2) is very similar to that presented by Foos et al. (2000) in Lechuguilla Cave, New Mexico. Both are topped by dusty gray material, underlain by massive to slightly laminated, dense, brownish-reddish clay, with white deposits filling in a series of crisscrossed fractures within the clay, followed by variably laminated, brown, hard clay. This similarity is supported by the findings of Onac et al. (2009a) and Pușcaș et al. (2010) which, based on stable sulfur isotopes, mineralogy, and passage morphology, argue for the existence of a hypogene stage during this cave’s genesis.
Profile S1 is the least complex, displaying clear limits from one facies to the next, with continuous horizontal beds. Profile S2 is the thickest (ca. 4.65 m) and in its upper third mimics section S1, although transitions are somewhat less explicit and numerous sand lenses are visible.
An interesting feature is the 50 cm thick conglomerate level with heavily weathered, brittle clasts
(sandstone, limestone, mudstone, gneiss) and white, altered carbonate cement. Poorly sorted levels of variously colored sand and gravel prevail in the bottom two thirds of S2. Ten meters down-stream, outcrop S3 is a thick pile of sediment (2.5-3 m) with visible horizontal layering and distinct granulometry (sand and gravel predominate), leaning against the cave wall in the inner curve of a bend. On the opposite wall, S4 is sheltered from erosion in an alcove, and displays sediment levels common to sections S1 through S3. Further downstream S5 comprises alternating levels of sand and gravel admixed with sand. In the bottom third of the alluvial deposit we encountered sediment sequences that were slightly indurated due to the presence of concentrated carbonate solutions, sometimes resulting in poikilitic calcite cement made up of large clear crystals (up to 1 cm). Similar observations were reported by Bosak et al. (2000).
26 The last profile, S6, can be linked to the bottom half of S5, and consists of a mixture of gravel,
sand, and coarse sand, with no trace of clay.
Only 7 out of the 24 sediment samples held φ -4 and larger fraction that allowed for a macroscopic and microscopic petrographic study of the clasts. One or two rock types, with minor additions, dominate each of them. The main participants are dark limestone (subangular to subrounded), light limestone (subangular to rounded), sandstone (angular to rounded), mica-rich
mudstone (generally subrounded), and granitoids (subrounded to rounded). Minor constituents
were found to be serpentinite (always angular), amphibolite (subrounded), diorite (subrounded to
rounded), gneiss (subrounded), quartzite (angular to rounded), microconglomerate (rounded),
and schist (subrounded to rounded). Several of the abovementioned clasts carry pyrite, sphalerite, and chalcopyrite grains, while muscovite, biotite, and quartz are also frequent constituents. Most of these pebbles are flattened and/or elongated and some are weathered and brittle.
The mineralogy of the sediment samples allows for some constraints to be put on their probable source area. All of the twenty analyzed samples showed different concentrations of garnet (light to dark brown and dark pink, probably spessartine and almandine), zircon
(transparent to light yellowish, with some very well preserved elongated prisms, characteristic for magmatic zircons (Pupin, 1980), titanite (transparent to light yellowish-brown flat prism fragments), olivine (dark green and brown fragments), tourmaline (dark brown and green),
sphalerite (dark golden-brown), pyrite/chalcopyrite (dark golden splinters), and pyrite (golden,
with some well preserved cubes).
Powder XRD results come from samples taken from three of the sediment profiles (S1
through S3). Quartz and muscovite are omnipresent, and samples from the top layers of S1 and
S2 are rich in phosphates (hydroxylapatite, taranakite, leucophosphite, and tinsleyite). Feldspars
27 are also common in most samples (sanidine, microcline, albite, and anorthite) giving important
clues about the source rock for these sediments. Diffractions were not targeted to identify clay
minerals, thus our findings are restricted to kaolinite and vermiculite. Gypsum is a regular
occurrence in fine-grained sand and clay lenses. A remarkable presence is that of nordstrandite, a
common weathering product in bauxitic soils derived from limestone (Anthony et al., 1997), but
an unusual cave mineral (Polyak & Provencio, 2001; Merino &., 2009).
The fact that a local fault guided the dissolution of the Main Passage is visible in several
locations, and could be a clue as to why the cave passages are linear. The only place along the
Main Passage where scallops could be clearly observed is at the entrance of the cave, on a
passage length of approximately 8 m, where the ceiling is less than 8 m high. Successive
generations of scallops can be observed on the ceiling and upper half of both walls, the more
recent being superimposed on the older ones, and considerably smaller (<1 cm), all indicating that the paleoflow direction was towards the valley of the modern Presacina Brook. Based on 20
measurements of the older generation of large scallops, we calculated the paleovelocity of the
subterranean stream to have been 0.32 m/s. The fact that current markings are present in such an
odd location, where they are usually destroyed by frost shattering, is rather difficult to explain.
One possibility is that this is the only sector in the cave where they were formed, due to
hydrological peculiarities. Other plausible explanations are that they do exist elsewhere in
GCSM but are too high up on the walls to be observed, or that they were erased by erosion, as signs of heavy mechanical weathering are visible throughout the cave. Also, the limestone at the cave entrance seems to belong to a different sedimentary facies than that in the rest of the cave, although we did not carry out thorough facies analysis.
28 Discussion and Conclusions
Sediments in GCSM are directly connected to conditions in the catchment area
(petrography, vegetation cover, rainfall, uplift and stream down-cutting) and are of both
allochthonous and autochthonous origin. The identified processes at work range from massive
collapse to fines transported as suspended load. Along the entire length of the Main Passage it is
clear that at least during one stage the void was almost completely filled with sediment (patches
of detritus on the walls, meandering ceiling channels, and pendants). Unfortunately, there is no
hard evidence to allow an estimation of the amount of sediment removed through stream erosion
as compared to that extracted for manufacturing gunpowder in historic times. The cross-section
of the passages suggests that pipe-full flow was a major genetic stage, while the shallow canyon
upstream of the NP (Fig. 2.1) implies a lowering of the base level after the sediment was
removed. A fairly simple cave pattern, with limited branching, no meandering or significant side
passages, and the steeply inclined floor of the Final Passage might advocate an organized
recharge (ponor-fed type cave) (the profile of the Speotimiș Passage is rather flat). The fact that
the general passage trend is almost perpendicular to the flow of the present surface stream
(Presacina Brook; Fig. 2.2), together with the presence of current markings (indicating that the
cave was discharging water into the brook), and the tilting of the cave passages toward Presacina
(profile view in Fig. 2.1) suggest that GCSM was carved by one of the springs feeding this brook,
and further on, both brook and cave must have been at approximately the same elevation at the time. Auler et al. (2009) found that the lack of dripstones, while flowstone is virtually omnipresent, may be explained through repeated cycles of sediment deposition and removal. We believe that our similar findings in the NP may be due to analogous conditions, as well as to the present dryness of this passage.
29 The cobbles made of sedimentary rocks (various types of limestone, sandstone, mudstone,
conglomerate and microconglomerate) are most probably sourced from the Presacina
Sedimentary Zone. The mineralogical composition of the fine-grained sediments is consistent
with felsic (granitoids, ortho- and paragneiss) as well as mafic (amphibolites and mafic
metaigneous rocks) source rocks. The presence of serpentine in sample PM 83 and olivine in
samples PM 83 and PM 72 is a strong evidence that the source area contained mafic lithologies,
such as the Sebeş - Lotru amphibolites that outcrop NE from our study site. The felsic lithologies
that supplied the sediments were most probably ascribed to Lainici - Păiuş type metasediments
(Liégeois et al., 1996; Iancu et al., 2005). Thus, we conclude that the original catchment area was probably larger than the modern one, although we are unable to set more precise confines.
An essential conditioning factor in the evolution of the Cerna River (and its tributaries) is the continuous activity of Cerna fault. Reactivated during Alpine orogeny (middle Cretaceous), the dextral strike-slip fault has been active (with subsequent lull periods) since and based on field
and map studies, a minimum separation of 30 km is assumed. The tectonic activity of a fault has
an important impact on the evolution of sedimentary basins situated along its strike and can
dramatically change the sedimentation rate, as well as the area that supplies sediments. It worth
mentioning that a separation of 30 km assumed for Cerna fault would have greatly impacted
Presacina Brook as well, thus modifying its catchment area and imprinting great heterogeneities in the sediment sources. Although calculated based on present temperatures rather than paleotemperatures and taking into account only the older generation of current markings, plotting paleoflow velocities calculated using ScallopEx 1.0.1 on a Hjulström diagram (Fig. 2.3) shows that at the point of exit from the cave the current was still strong enough to erode and transport fine grained material, but it would have already deposited coarse alluvium.
30
Figure 2.3. Hjulström diagram with dashed line showing the flow velocity calculated in ScallopEx.
The mineralogical and petrologic study of the sediment deposit indicates that mostly allogenic waters carved the cave. As pointed out by Palmer (2001), in the case of allogenic streams most of the cave growth takes place during flood events, which are usually restricted to certain periods in the hydrologic cycle. During peak levels debris carried by the floodwaters can dam the underground stream, resulting in slack water. Normal flow and flood conditions, as well as slack waters, can be deducted in GCSM from the sediment pile deposited in the NP.
Alternating levels of gravel, sandy gravel, fine to coarse sand, and clay imply frequent variations in flow energy (repeated cycles of normal flow interrupted by flooding, damming, and ponding) inherent to ponor-related caves, which are quick to respond to even relatively small discharge fluctuations. Most pebbles are sub-rounded to well rounded, elongated and/or flattened, and
31 habitually positioned on their flat side, parallel to the paleocurrent, indicating regular high-
energy flow conditions. However, some sediment levels show no clast orientation, clearly
depicting flood conditions.
On the basis of morphological clues, sedimentary facies and mineralogical and petrographic makeup of the alluvium we argue that the major speleogenetic phases involved were as follows: 1. epiphreatic regime (suggested by the quasi tubular cross-section of the passages); 2. vadose flow with repeated depositing and removal of significant volumes of sediment (implied by the ceiling pockets, ceiling channels and residual patches of fine sediment on the walls); 3. removal of the sediment from the entire cave, with the exception of the NP (our initial working hypothesis –the rerouting of the cave stream toward underlying cave passages, following the base level lowering- although supported by the presence of breakdown that could mask the existence of passages capable of removing important volumes of sediment and the occurrence of cave entrances a few tens of meters below the openings of GCSM, was contradicted by the meager size of the aforementioned caves, and the complete absence of sediment therein. We believe that the preservation of the alluvium in the NP, could be due to a particularity of the bedrock morphology –a concave segment- in this sector of GCSM); 4. deepening of the local base level followed by massive collapse; 5. dripstone growth in the final sector of the Main Passage, and flowstone throughout the cave.
The main processes at work in GCSM (sediment input, sediment erosion and speleothem precipitation) hold paleoenvironmental value; the large mass of alluvium deposited here suggesting a prolonged period of wet climate, punctuated by frequent flooding. Dating this sediment would enable us to establish whether the floods were more common than in today’s climate, as well as the time interval during which the sediment was deposited. The presence of
32 nordstrandite, which occurs mostly in regions that benefit from a warm and moist type of weather (Hathaway & Schlanger, 1965; Dani et al., 2001), identified in a sandy gravel stratum could point to such climate conditions. Nordstrandite was also reported to precipitate within caves under hypogene conditions (Onac et al., 2009c; Spilde et al., 2009). As of now we are unable to discriminate between a primary or secondary origin for this unusual mineral. The subterranean stream left this cave (making room for drip- and flowstones) either because of tectonics (valley down-cutting as a result of uplift), a change in climatic conditions leading to a somewhat cooler and dryer stage, or a combination of both. The absence of wood debris in the sediment profiles is probably due to their incompatibility on a long term with the cave environment. Although not yet completed, our study opens the door for future research in the area, including multiple dating possibilities (pollen, quartz pebbles, flowstone, dripstone, sediment) and assessments of valley down-cutting rates in the region.
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39 CHAPTER TREE:
THE MINERAL ASSEMBLAGE OF CAVES WITHIN ŞĂLITRARI MOUNTAIN
(CERNA VALLEY, SW ROMANIA): DEPOSITIONAL ENVIRONMENT AND
SPELEOGENETIC IMPLICATIONS
Note to Reader
This chapter has been previously published in Carbonates and Evaporites, 2010, 25: 107-
115, and has been reproduced with permission from Springer. The author of this dissertation is the first author of the article and its major contributor.
Introduction
Loose sediment deposits within caves have proved time and again that their study can be applied to a wide array of research fields (hydrogeology, paleoclimatology, paleontology, anthropology, mineralogy, etc.) (Sasowsky & Mylroie 2004; White 2007). Interactions between bedrock, alluvium, and percolating or seeping solutions, under restrictive cave conditions produce minerals indicative of their environment of formation (Hill & Forti 1997; Onac 2005).
Sulfuric acid speleogenesis (SAS) is responsible for typical cave morphologies and features
(Audra et al., 2002, 2007a; Sancho et al., 2004; Klimchouk 2007), cave sediments (Foos et al.,
2000; Audra et al., 2002), specific ecosystems (Sârbu et al., 1996; Engel 2007), and mineral associations (Hill 1987; Polyak & Güven 1996; Audra et al., 2007b; Onac et al., 2007, 2009a, b).
40 Figure 3.1. Topographic map of the studied area. Modified after Ponta & Solomon, 1982. Inset map of Romania with the location of Băile Herculane.
Hypogene-related cave morphologies, particular types of sediment, unusual mineral
assemblages, and their location within the cave might hold the key to cave genesis.
Thermo-mineral water activity around the Roman town of Herculaneum (today Băile
Herculane, inset in Fig. 3.1) (Cristescu 1978), together with evidence of saltpeter mining from cave deposits for the production of gunpowder in the times of the Ottoman and Austro-
Hungarian empires (Ponta & Solomon 1982; Diaconu & Lascu 1998-1999) raised a relatively recent interest in the mineral assemblages of caves in the region. The caves from Muntele
Şălitrari (Saltpeter Mountain), located on the right bank of the Presacina Brook, a right-hand
41 side tributary of the Cerna (Fig. 3.1), host large deposits of alluvial sediments that were altered
by percolating nitrogen- and phosphorous-rich solutions produced by bat colonies. The limestone
bedrock was equally affected by these guano leachates. This article, part of a more extensive
research concerning cave morphologies and mineral associations along the Cerna Valley,
describes the mineral assemblages of three such caves, assesses the role of sulfuric acid in their
genesis and attempts to distinguish between several overlapping speleogenetic processes.
Geographic and Geologic Settings
Located in the southwestern extremity of the Southern Carpathians, the Cerna Valley is
hosted on its entire 85 km length by a tectonic trough. Its 555 km2 hydrologic basin drains into
the Danube to the south. The significant drop in altitude from springs to confluence (ca. 1950 m
of relief) results in a highly fragmented landscape, carved mainly in schist, granite, and
limestone, with canyon cliffs dropping almost vertically, up to 500 m (Alexandru et al., 1981).
The Cerna region experienced a complex and sustained geological evolution. The
basement is represented by Precambrian and Cambrian crystalline schists, overlain by lower
Paleozoic metamorphic and igneous rocks, upper Paleozoic meta-sedimentary and igneous formations (Bojar et al., 1998; Kräutner & Krstic 2002), on top of which lay several Mesozoic sedimentary cycles and finally Tertiary deposits, present mostly in the surrounding depressions.
The Cerna Valley limestones belong to the Danubian sedimentary unit; based on their fossil contents, their ages are Jurassic to Lower Cretaceous (Năstăseanu 1980; Iancu et al,. 2005).
Avram et al. (1964, 1966), and Dancău et al. (1968) carried out extensive cave survey and inventory in the region, whereas Ponta & Solomon (1982) focused solely on the Presacina
Brook karst area. Several papers describe the cave mineral associations along the Cerna Valley
42 (Povară et al., 1972; Diaconu & Medeşan 1973; Diaconu 1974; Onac et al., 2009a, b), with two
others specifically considering caves within Şălitrari Mountain (Diaconu & Lascu 1998-1999;
Onac et al., 2009a). Onac et al. (2009a) and Sumrall (2009) also studied sulfur isotope ratios in minerals from the Peştera Mare din Muntele Şălitrari (Great Cave from Şălitrari Mountain,
(GCSM)).
Cave and Sample Description
Thirty-seven samples have been collected from three of the ten caves developed within
the Şălitrari Mountain, on the right-hand side of the Presacina Brook, namely Şălitrari caves no.
2 and 3 (Fig. 3.2), and GCSM (Fig. 3.3). All three cave entrances are located between 400 and
500 m above sea level (asl). The general orientation of the major cave passages in the Şălitrari
Mountain is approximately perpendicular to the Presacina Brook (Fig. 3.1). Şălitrari Cave no. 2
(SC2), 440 m long, is located approximately 0.5 km upstream of the confluence of Cerna and
Presacina (Fig. 3.1), at an altitude of 420 m asl. It is accessible through two large entrances and
contains wide passages and large chambers, with floors covered by sediment, clay or breakdown.
Speleothems of any kind are lacking. Several locations along the main passage hold evidence of
early mining of saltpeter (Ponta & Solomon 1982). Unfortunately, looters seeking intact cave
bear bones meant for contraband are carrying out more recent diggings, disturbing the cave
sediment. We collected 7 samples from various locations within the cave (Fig. 3.2), in the form
of crusts, nodules, aggregates or earthy masses, ranging in color from white to dark brown.
Şălitrari Cave no. 3 (SC3) situated ca. 30 m east from the second entrance in SC2, at an altitude
of 455 m asl, is the smallest of the caves presented here (only 42 m long; Ponta & Solomon
1982).
43 Figure 3.2. Map of SC2 and SC3 (modified from Ponta & Solomon, 1982) showing sampling locations.
44 Figure 3.3. a. Map of the GCSM (mapped by ‘‘Prusik’’ Timișoara) showing sampling locations. b. stratigraphic column of the sediment deposit, showing minerals identified in each level (C calcite, G gypsum, Q quartz, K kaolinite, M montmorillonite, Ar ardealite, Am aluminite, Al alunite, D darapskite, T taranakite, V variscite, B brushite). c. The alluvial sequence described in inset b.
45 The longest cavity in the investigated area, a few hundreds of meters east of SC2 and at
an altitude of 480 m asl, is GCSM. The total length of its known passages is 1,500 m. Ponta &
Solomon (1982) divided the cave into two sectors according to their general direction and size of passage cross-section. A third sector was discovered in the 90s by the Speotimiş Caving Club
(Fig. 3.3). The first sector, which is the only one lacking speleothems, contains the so-called
Nitrate Passage, where temperature is higher and relative humidity is lower than in the rest of the cave (~11.7ºC and <75%, as compared to <10ºC and 96-100%) (Diaconu & Lascu 1998-
1999; Onac et al., 2009a). This passage is approximately 30 m long, 20 m wide, and 10 m high.
Sediment and limestone boulders cover the limestone floor. Cave stream erosion and saltpeter mining activities exposed a thick alluvial deposit that once filled most of the passage in this section of the cave. The sequence consist of sediments ranging from sand and pebbles at the base
(~1 m), to an intermediate clay level (crisscrossed by white veins and lenses rich in phosphate minerals; ~0.85 m) and a 1.10 m thick silt-sized material containing nitrates, sulfates, and some phosphates atop of the sequence (Fig. 3.3, insets). This sector also holds testimony that more recently, looters have heavily degraded the sediment searching for cave bear bones. The relatively short Nitrate Passage (along with some remote areas of the cave) is the source for most of our samples (Fig. 3.3), represented by nodules, crusts, aggregates, acicular or fibrous crystals, highly weathered limestone fragments or earthy masses.
Analytical Methods
Preliminary observations and sample selection were conducted with the aid of a Nikon
SMZ 1500 stereomicroscope, equipped with a Nikon NI-150 High Intensity Illuminator and a
Nikon Digital Sight DS-U1 camera.
46 Powder X-ray Diffraction (XRD) patterns were obtained for 37 samples on a Rigaku
MiniFlex diffractometer (University of South Florida (USF), Department of Geology), working conditions being 30 kV and 15 mA, with a Ni Kβ-filtered CuKα radiation. A fixed 1.25° scattering slit and a 0.3 mm receiving slit were used to collect patterns. Samples were continuously scanned (at a speed of 0.5 or 1º/s) from 5 to 70º 2θ, using a fixed step size of 0.02º
2θ/s. The internal standard was Silicon (NBS-640b).
Electron microprobe (EMPA) was used to acquire further chemical data on more
challenging minerals. We employed a JEOL 8900 R Superprobe (Florida Center for Analytical
Electron Microscopy, Miami). Kaersutite (SiO2, K2O, FeO), apatite (CaO, P2O5), anorthite
(Na2O, Al2O3), diopside (MgO), and anhydrite (SO3) served as standards. Scanning electron
microscopy (SEM) was carried out on a JEOL 6490 LV in the Lisa Muma Weitz Imaging Core
Laboratories, at the University of South Florida, samples being sputter-coated with Au.
Results
The mineralogical association in the three caves from the Şălitrari Mountain, first
described by Diaconu & Lascu (1998-1999) and later expanded by Onac et al. (2009a),
comprises the ubiquitous calcite and gypsum, the somewhat less frequent brushite, taranakite, apatite, and quartz, and the rare darapskite, aluminite, and alunite. The complete mineral
assemblage of these caves is presented in Table 3.1. All three caves have passages that display a
variety of common calcite speleothems, the most impressive being GCSM. Calcite also occurs in
the form of white, yellowish or brown crusts, twinned aggregates, or transparent acicular
crystals. The samples were collected at the base of cave walls and as weathered limestone
fragments that sit on top or were found buried in layers of sediment or guano.
47 Traces of pyrite (FeS2) along with its weathering products, hematite (Fe2O3), and goethite
[Fe3+O(OH)] were all found in a rust-colored, hard nodule, in the wall of a secondary passage in
GCSM (identified by XRD and petrographic microscope).
Table 3.1. List of minerals discussed in this study.
Gypsum [CaSO4·2H2O] forms crusts of various colors, white damp nodules, translucent
aggregates, and colorless, transparent, acicular crystals. Samples were collected from walls,
highly weathered limestone blocks, and at various levels within the alluvial deposits. Identified
through all of the methods presented in the previous section, gypsum was found in various
settings in each of the three caves.
Apatite-(CaOH) and apatite-(CaF) are also present in all three caves, as brown, brittle
material associated with bone deposits, black or brown crusts on limestone boulders, and white
48 brittle aggregates within various levels of the alluvial deposits.
Brushite [CaHPO42H2O] was collected from the Nitrate Passage in GCSM as white, soft crusts lining veins of taranakite and variscite in the upper clay level and as white crumbly nodules at the base of the wall, and from SC3 as gray hydrated nodules covering cave popcorn.
In both cases it was identified by XRD.
Soft, white nodules collected from the top of the sediment sequence, immediately below the guano layer, in the Nitrate Passage of GCSM, were identified by XRD as ardealite
[Ca2(SO4)(HPO4)·4H2O].
Taranakite [K3(Al,Fe)5(HPO4)6(PO4)218H2O], identified by XRD and SEM (Fig. 3.4), was collected from various levels of the sediment sequence as small, white, brittle nodules, and white veins, associated with variscite in the lower part of the sediment deposit, and in some cases with darapskite, aluminite and alunite, in the upper level.
Figure 3.4. Secondary electron image of pseudohexagonal platy taranakite crystals.
49 Variscite [AlPO4·2H2O] was found in the form of white veins or nodules, near the contact with a thick bed of desiccated clay, which probably supplies the necessary Al, while P is leached from decomposing guano.
Darapskite [Na3(SO4)(NO3)H2O] was collected from the very top of the sediment sequence (insets B and C in Fig. 3.3) as a fine, compact, grayish material with small white nodules dispersed within it, and identified by means of XRD and SEM (Fig. 3.5). It is associated with gypsum, calcite, quartz, and aluminite.
Aluminite [Al2(SO4)(OH)47H2O] was found in association with gypsum, forming white crumbly nodules above the clay layer. It was documented by SEM, EMPA (Table 3.2), and XRD.
Alunite [KAl3(SO4)(OH)6] forms nodules of white powdery material and was identified by means of XRD, SEM, and EMPA (Table 3.2).
Figure 3.5. Secondary electron image of prismatic crystals of darapskite.
50 In GCSM, alunite occurs within the silt and clay-sized sediment horizon in the upper part of the alluvial profile, similar to one of the several depositional environments described by
Polyak & Güven (1996) from Carlsbad, Spider, and Lechuguilla caves. Quartz, kaolinite, montmorillonite, and illite, identified in XRD spectra, are allochthonous minerals, derived from the alluvial sediment deposit.
Table 3.2. EMPA results for alunite and aluminite from the Great Cave from Șălitrari Mountain.
Discussion and Conclusions
Seventeen cave minerals belonging to 8 chemical groups were identified from three caves
along Presacina Brook, a tributary of the Cerna River, by means of XRD, EMPA, and SEM.
These mineral assemblages are by-products of complex interactions taking place at different
moments at the interface between limestone bedrock, cave sediment, organic material,
descending meteoric water, and formerly ascending thermo-mineral water/steam.
51 Phosphates
The reaction between the phosphate ion (leached from guano and bones), the
allochthonous siliciclastic sediments, and limestone is ultimately responsible for the precipitation
of a variety of phosphate minerals within the cave environment (Hill & Forti 1997; Onac 2005).
The phosphate association in SC2, SC3, and GCSM is linked to the presence of bat colonies and
cave bear remains within these caves. In the GCSM, in particular along the Nitrate Passage, we
explain the precipitation of phosphates throughout the entire sediment sequence through leaching
of mobile elements from the top guano layer towards the bottom of the alluvial deposit.
Both brushite and taranakite require damp environments and acidic solutions at the time
3- of their precipitation (Hill & Forti 1997). The PO4 is derived from solutions percolating from the top guano layer, whereas limestone supplies Ca2+ and clay supplies Al3+, K+, and Fe3+. The thick (~1 m) clay level where they precipitated in the Nitrate Passage probably prevents them from dehydrating.
The micro-climatic settings prevailing along the Nitrate Passage, where the ardealite sample was collected, closely match the ones earlier reported in the literature (Hill & Forti 1997;
Anthony et al., 2000) for the typical occurrences of this mineral. The mineral is a result of guano breakdown in this dry passage, by reactions with carbonate material. Such reactions are often mediated by the presence of sulfate with δ34S values ~6.5‰ (Onac et al., 2009a), indicating that
the presence of this mineral could be linked to SAS.
Nitrates
Nitrates source nitrogen from organic matter within the cave and from organic-rich
meteoric water percolating from the surface. Nitrate cave deposits are linked to the activity of
nitrifying bacteria in the caves and/or the soil above them, to the presence of clastic sediment
52 within the caves, providing the alkali, and require low humidity values (Hill, 1981). Darapskite in particular is stable between 13 and 74ºC (Ericksen & Mrose 1970) and 50 to 84% relative humidity values (Frazen & Mirwald 2009), whereas nitratine, previously reported by Diaconu &
Lascu (1999), has a stability field that spans the entire range of cave temperatures (4-22ºC) and from 71 to 81% relative humidity (Hill & Forti 1997). These particular environment conditions are met in the Nitrate Passage from GCSM.
Sulfates
Gypsum in the Nitrate Passage is a result of the interaction of limestone and sulfuric acid resulted from the oxidation of ascending sulfide-rich thermo-mineral solutions. A positive proof of this genetic pathway is the sulfur isotope signature (δ34S) of gypsum and other sulfate minerals in the Nitrate Passage that shows values between 0.5 and 6.5‰ (Onac et al., 2009a).
Alunite forms under a wide temperature range (15-400C) generally indicating acidic conditions (Anthony et al., 2003).
Aluminite and alunite are confined to the upper half of the sediment deposit in the Nitrate
Passage, above the thick clay bed, indicating prolonged dry and warm conditions that preserved them. Their position within the clastic sequence suggests that sulfidic solutions reacted with the alluvium at some point during the vadose phase, but after the sediments were set in place.
Coupled with their sulfur isotopic signature, these minerals might be indicative of SAS
(Cunningham et al., 1995; Polyak & Güven 1996; Polyak & Provencio 2001).
Pyrite from the Speotimiş Gallery in GCSM is diagenetic in origin. Its weak XRD patterns along with the well-defined ones for hematite and goethite suggest that the two oxides formed in a low-temperature environment by the oxidation of primary pyrite by seeping meteoric water.
53 The warm and relatively dry cave environment exhibited by the Nitrate Passage in
GCSM is host to the richest and most exotic mineral assemblage (Table 3.1). Several of these
minerals are sensitive to temperature and relative humidity variations and easily undergo phase changes (dehydration, phase transformation) towards more stable species (Howie, 1992; White,
1997). This suggests that either this very cave is the only one to provide appropriate conditions for such minerals to precipitate (relatively low humidity and higher than normal temperatures), or that such minerals existed in most of these caves, but that GCSM is the only one whose topoclimate allowed their preservation.
After a phreatic phase (suggested by elliptical cross-section of some passages and wall notches), the water table dropped and vadose flow was initiated, as indicated by the thick alluvial sediment deposits. We find further evidence for vadose conditions in the orientation of cave passages that are approximately perpendicular to Presacina Brook (Fig. 3.1), suggesting the underground streams were its tributaries. Alluvial deposits are thicker towards the entrance of the caves. Streams that undercut cave walls caused the significant limestone breakdown visible today.
Once the vadose regime was attained, speleothems started filling some of the passages.
Typical morphological features (e.g., ceiling cupolas, blind ascending passages), together with the mineral assemblage and its S isotope signature (Onac et al., 2009a; Sumrall, 2009) point towards the existence of a SAS phase, mediated by sulfide-rich thermal fluids that enter the cave from depth along fractures and joints. Vadose features subsequently overprinted part of this SAS stage.
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59 CHAPTER FOUR:
TAMARUGITE-BEARING PARAGENESIS FORMED BY SULFATE ACID
ALTERATION IN DIANA CAVE, ROMANIA
Note to Reader
This chapter has been previously published in European Journal of Mineralogy, 2013,
25: 479-486, and has been reproduced with permission from E. Schweizerbart'sche
Verlagsbuchhandlung. The author of this dissertation is the first author of the article and the
major contributor.
Introduction
The Diana Cave (22 m long) in Băile Herculane (SW Romania; Fig. 4.1A) has an
entrance passage of ~8 m that develops along the Diana Fault, which causes the nodular
limestones (J3) and the overlaying Lower Cretaceous (K1) marls of Iuta strata to come in direct contact (Năstăseanu, 1980). The fault also facilitates the emergence of two thermo-mineral H2S- rich springs (Diana 1 and 2; hereafter Diana springs) at the inner end of this passage (Povară et al., 1972). From this point, the cave passage continues for another 14 m on a NW-SE direction
(Fig. 4.2). In the early 1970s this second passage was enlarged to drain the thermo-mineral water into a pumping station located just outside the cave.
60
Figure 4.1. A. Location of the studied site. B. Geological map of the region (modified from Năstăseanu, 1980). 1. crystalline schists (Pre2); 2. wildflysch with olistoliths (K2); 3 Urgonian limestones (K1); 4. Iuta strata (K1); 5. bedded limestones (K1); 6. nodular limestones (J3); 7. conglomerates (J1); 8. crystalline basement (Pre2); 9. Diana Cave; 10. Diana springs; 11. normal fault; 12. thrust fault (Cerna Fault).
Therefore the entrance passage was also enlarged, the cave walls were reinforced with cement,
and the water drain was covered with concrete slabs. The major ions dissolved in the Diana
+ + 2+ - 2- thermo-mineral springs are Na , K , Ca , Cl , and SO4 (Diaconu, 1974; Povară et al., 2010); water temperatures average 51.64 °C, pH of 7.42, summer cave air temperature is ca. 22 °C, and relative humidity > 95%. Due to the existence of only one cave entrance the general air circulation is fairly simple: cold air from the outside sweeps into the cave along the floor, heats up at the contact with the thermo-mineral water, ascends, and exits the cave along the ceiling
(Povară et al., 1972). Currently, the concrete covering the cave walls is heavily corroded by the sulfuric acid resulting from the oxidation of H2S outgassing from the Diana springs. The lower part of the cave walls as well as the concrete slabs that cover the water drain display abundant gypsum crusts, soggy aggregates of native S, and a variety of unusual cave sulfates (Figs. 4.2,
4.3).
61 Figure 4.2. Map of the Diana Cave (modified from Povară et al., 1972); water storage building not at scale. Inset: Photograph of the weathered concrete slabs covering the water and the minerals deposited on top of them (field of view ca. 1 m).
62 Figure 4.3. Photograph of tamarugite-rich aggregates (in paragenesis with native sulfur, halotrichite-group minerals, and gypsum).
Among them is tamarugite, NaAl(SO4)2·6H2O, a mineral that has been previously identified only from non-limestone caves associated with secondary volcanic activities, i.e., thermal springs or fumaroles (Mackenzie et al., 1995; Rodgers et al., 2000; Demartin et al., 2010; Onac & Forti,
2011a).
Onac et al. (2009) first identified tamarugite in Diana Cave, a typical limestone cavity with no relationship to any past or present volcanic activities; this paper gives a detailed description of the mineral and its paragenetic assemblages. The latter is of interest because shed light on cave-forming environments in which sulfuric, rather then carbonic acid is the driving the dissolution/minerogenetic processes. By-product minerals (gypsum, native sulfur, Al-sulfates, etc.) produced during sulfuric acid speleogenesis (Polyak & Provencio, 2001) may provide evidence of past steam-condensate alteration conditions in caves that no longer host active H2S
63 systems or in which subsequent processes removed part of the morphological and mineral evidence.
Local Geology and Hydrology
The geology of the Cerna Valley is complicated by the occurrence of several nappe
complexes, such as Getic, Danubian, etc. (Balintoni et al., 2010). Diana Cave develops at the
contact between nodular limestones (J3) and dark grey Iuta marls (K1), both part of the
sedimentary cover of the Danubian nappe. Initially referred as Nadanova strata, the name for the
marl deposits along the Cerna Valley was later changed to Iuta strata, to help distinguish between units with similar lithologies but different ages (Năstăseanu, 1980). The major clay minerals within the Iuta marls (forming the floor and lower part of the cave walls) are illite, kaolinite, and chlorite (Diaconu & Medeșan, 1975). Their importance resides in the ions released to the cave geochemical system upon reacting with the H2SO4 that condenses on cave walls.
Along most of its length Cerna’s streambed overlaps the Cerna Fault, a dextral strike-slip fault that is part of an intricate fault system forming a major graben (ca. 1000 m of vertical displacement; Iancu, 1976). Meteoric water descends to significant depth seeping through karst units and down some fault zones; it becomes heated and mineralized resurfacing as thermo- mineral springs (with temperatures varying between 23 to 55 °C) along tectonic deformations or by means of deep artesian wells (Povară et al., 2008). The heating process is due to the high geothermal gradient associated with deep crustal fractures (Gheorghe & Crăciun, 1993).
Based on hydrogeological and chemical differences, the thermo-mineral aquifers in the
Cerna River basin have been divided into the Northern Aquifer (mostly carbonate waters, 9.8 –
2- 34 °C, low SO4 concentration and low total dissolved solids –TDS) and Southern Aquifer (38 –
64 2- 60 °C, high TDS, H2S, and dissolved SO4 ) (Povară et al., 2008; Wynn et al., 2010). The thermal water flowing through Diana Cave belongs to the latter aquifer.
Tamarugite Occurrences and Stability
Tamarugite was first described from Mina de la Compania, Sierra Gorda, Chile (Gordon,
1940). It has also been identified in cavities within the Vulcanello Crater (Vulcano Island, Sicily;
Lombardi & Sposato, 1981), Grotta dello Zolfo (Italy; Zambonini, 1907; Forti et al., 1996), and
Ruatapu Cave (New Zealand; Mackenzie et al., 1995; Rodgers et al., 2000). In all these locations, the origin of tamarugite is linked to post-volcanic processes. Excluding the study of Onac et al.
(2009), the only other non-volcanic occurrence of tamarugite was recently reported by Lazaridis et al. (2011) from the Aghia Paraskevi caves (Greece).
2- In surface conditions, tamarugite has been found to precipitate in the presence of SO4 ions, the source of which is either related to volcanic activities (e.g., fumaroles; Delmelle &
Bernard, 2000; Bortnikova et al., 2005; Vaughan et al., 2005; Balić-Žunić et al., 2009) or oxidized pyrite in sedimentary (mostly carbonatic) deposits (Keller, 1935; Segnit, 1976; King,
1998). Although it is reported mainly from arid regions, coastal processes have also been shown to create tamarugite as a weathering product of pyrite and Al3+-rich clays from the host-rock and
Na+ from sea spray (Garvie, 1999 and references therein). The mineral is also indicative for acid
sulfate soils (Raven et al., 2010).
Three secondary, water soluble, hydrated sulfate minerals are known with the general
formula NaAl(SO4)2·nH2O (with n ranging from 6 to 12): tamarugite (n = 6), mendozite (n = 11),
and Na-alum (n = 12). One other Na-Al sulfate, but with a different stoichiometry, is natroalunite
NaAl3(SO4)2(OH)6. Mendozite easily dehydrates to tamarugite under atmospheric conditions and
65 as a result is extremely rare (Keller, 1935; Fang & Robinson, 1972). So far mendozite has not
been reported from any caves, indicating that tamarugite is the first species to precipitate, or that
cave conditions are such as to immediately convert it to tamarugite, the thermodynamically
stable species under cave settings. Tamarugite was reported from a wide range of ambient humidity (up to 90 %; Sabelli et al., 1987) and temperature values (Lombardi & Sposato, 1981), but the pH in all occurrences is moderately to highly acidic. Sodium alum and natroalunite in caves are also indicative of acidic conditions (Forti et al., 1996; Polyak & Provencio, 2001) and
appear stable under a wide range of cave atmosphere temperatures and relative humidity.
According to Fang & Robinson (1972) sodium alum can dehydrate directly to tamarugite,
without passing through the intermediate mendozite phase.
Description of Samples and Analytical Methods
Samples of tamarugite were collected from the walls and floors of both natural and
reinforced cave passages (sampling locations are shown in Fig. 4.2). They form white, brittle nodules and paper-thin crusts, pasty, white, yellow, or orange colored masses, and dark yellow,
well-crystallized crusts (as described by Hill & Forti, 1997 and Keller, 1935, from other
localities). Water samples were analyzed in parallel for stable isotopes as well as physical and
chemical properties.
Mineralogical investigations were carried out on a Rigaku Miniflex (University of South
Florida-USF, Tampa, USA) and a Philips X’Pert theta-2theta type powder X-ray diffractometer
with scintillation counter (Institute of Mineralogy and Crystallography, Vienna, Austria). Topas3
software (Bruker AXS) was applied for Rietveld refinements. Despite our efforts, no individual
66 tamarugite crystals of sufficient size were available for single-crystal X-ray diffraction
experiments.
Samples containing tamarugite were investigated by means of a JEOL JSM 6490 SEM
equipped with secondary (SE) and backscattered electron (BSE) detectors operated in low
vacuum mode (Lisa Muma Weitz Laboratory, USF).
Sulfur stable isotopic measurements of gypsum, bulk sulfates, and precipitates from
thermal-water samples were carried out using a Costech Elemental Analyser (EA), connected to
a Thermo Finnigan Delta V continuous flow Isotope Ratio Mass Spectrometer (IRMS) at USF.
Reference materials used were IAEA SO-5 and IAEA SO-6 for sulfate and IAEA S-2 and IAEA
2- S-3 for the dissolved sulfide precipitated from the spring water. Dissolved SO4 was analyzed as
- BaSO4, whereas dissolved HS was measured as AgS. The reproducibility between replicate
standards in each run was estimated to be better than ± 0.1 ‰ (1σ). We chose to analyze gypsum,
as it was impossible to separate a sufficient amount of pure tamarugite.
Further data on the Diana springs’ chemistry and stable isotope composition of thermal
waters and cave sulfates are obtained from, or compared to other publications (Diaconu, 1974;
Diaconu & Medeşan, 1975; Povară et al., 2008; Onac et al., 2009, 2011; Wynn et al., 2010).
Results
Tamarugite is monoclinic, space group P21/a, with unit-cell parameters refined from
powder data: a 7.358(6), b 25.23(2), c 6.093(5) Å, β 95.16(5)º, V 1126.98(1) Å3, and Z 4 (Table
4.1). Mineral species identified by Rietveld refinements of powder XRD patterns are compared
to previous studies in Table 4.2. One sample showed the presence of small amounts of halite, the
origin of which is most likely the thermal water (Diaconu, 1974; Povară et al., 2010). A
67 characteristic pattern in which tamarugite (about 20 wt%) occurs along with alunite, anhydrite,
and native sulfur (roughly 30, 40, and 10 wt%, respectively) is given in Fig. 4.4.
Table 4.1. Comparison between the unit-cell parameters of tamarugite from Diana Cave, Romania and Mina de la Compania, Chile. After Robinson & Fang, 1969.
High-magnification, BSE images of tamarugite from Diana Cave acquired via SEM (Fig.
4.5) clearly indicate that tamarugite forms tabular (parallel to {010}) euhedral to subhedral
crystals, showing comparable morphologies to those published by Lombardi & Sposato (1981),
Sabelli & Santucci (1987), Rodgers et al. (2000), and Balić-Žunić et al. (2009). Individual crystals rarely show signs of dissolution and they are significantly smaller (up to 15 µm across) when compared to those reported in literature.
Table 4.2. List of cave mineral species identified in Diana Cave and associated formation processes.
68
Figure 4.4. Typical powder XRD pattern of a sample consisting of an admixture of tamarugite, anhydrite, alunite, and native sulfur. The phase identification is confirmed by a Rietveld refinement (program Topas, Brucker, 1999).
69 The average stable isotope δ34S values of gypsum and other bulk sulfate samples from the
2- 2- Diana Cave is 18.74 ‰, whereas in the thermo-mineral water (both in SO4 and S ) it is
23.88 ‰, both values suggesting a marine evaporite source for sulfur. However, neither gypsum nor other evaporites occur in the lithologies surrounding the cave; therefore, precipitation of sulfates in cave from such deposits is ruled out. In accordance with Wynn et al. (2010), who reported the presence of a complete sulfur cycle in the thermal aquifers of Cerna Valley, in the
Diana Cave sulfur is related to the thermal water. Furthermore, Onac et al. (2011) documented a clear relationship between the S isotopic composition of cave sulfate minerals in this region and the δ34S values of sulfates in the thermal waters from which they precipitated under various
geochemical conditions.
Figure 4.5. SEM back-scattered electron images of tamarugite aggregates.
Discussion
Minerogenetic mechanisms (Onac & Forti, 2011b) responsible for the peculiar mineral
assemblage in Diana Cave are CO2 diffusion, oxidation, dehydration, double replacement, and steam-condensation alteration. The geochemical settings and their key chemical reactions
70 responsible for subaqueous and subaerial aggressive dissolution of the limestone/marl bedrock, gypsum replacement of limestone, subaerial precipitation of extensive sulfate deposits, and deposition of native sulfur in the cave are illustrated in Fig. 4.6 and discussed below. Zones A and B (Fig. 4.6) host products of some particular geochemical reactions. When entering the oxygen-enriched cave waters (subaqueous part of zone A) dissolved sulfides mobilized from deep anoxic reservoirs are rapidly oxidized to elemental sulfur through reaction 1 or 2 (Palmer,
2007).
- + 0 HS + ½O2 + H S + H2O (1)
0 2H2S + O2 2S + 2H2O (2)
These reactions are intermediate phases as under oxic conditions native sulfur further
oxidizes to produce sulfuric acid (reaction 3; Palmer, 2007).
0 2- + 2S + 2H2O + 3O2 2SO4 + 4H (3)
Moving upward into the subaerial part of zone A and in zone B, the hydrogen sulfide
outgassing from the hot thermal water is converted into extremely aggressive sulfuric acid (4),
causing the partial replacement of the limestone bedrock by gypsum (5; reactions 4 and 5 from
Palmer & Palmer, 2000):
+ 2- H2S + 2O2 2H + SO4 (4)
+ 2- 2H +SO4 + CaCO3 + 2H2O CaSO4·2H2O + H2O + CO2 (5)
This aggressive alteration of the carbonate rock (known as sulfuric acid speleogenesis)
extends from the ceiling downward only to, or slightly below, the water table (zones A and B)
and is responsible for further enlargement of cave passages. Due to the presence of abundant H+ in both zones (A and B), limestone weathers according to the reaction below:
+ 2+ CaCO3 + 2H Ca + H2O + CO2 (6)
71 Figure 4.6. Schematic cross-section (not at scale) through Diana Cave. showing The primary geochemical reaction zones (A and B; separated by the dashed line), the source of ions for the mineral assemblage described, and the typical occurrence of the main mineral species are pictured. (a) Steam-condensate alteration film; (b) gypsum crusts; (c) halotrichite-group minerals; (d) tamarugite aggregates, (e) native sulfur; (f) gypsum rafts.
72 2- 2+ Given the availability of SO4 and Ca , gypsum also forms in the following way:
2+ 2- Ca + 2H2O + SO4 CaSO4·2H2O (7)
Reaction 7 is responsible for the precipitation of thin gypsum rafts on the water surface, as well as gypsum aggregates on the cave floor and walls (Fig. 4.6b and f). Tamarugite and
minerals of the halotrichite group are products of the steam-condensate alteration occurring at or below the limestone/marl boundary (Fig. 4.6).
The presence of tamarugite indicates sufficiently strong acidic conditions in the steam condensate to release Na+ and Al3+ from the marls (Fig. 4.6a and c). In addition to any Na+ released from the bedrock, an important contribution of this ion must come from the Na-Cl-type thermal water, which is periodically in contact with the marls located in lower part of the cave passage. Therefore, we consider the chemical reaction proposed by Rodgers et al. (2000) for tamarugite formation in environments affected by steam condensate alteration, to also apply to
Diana Cave:
+ 2- + Al2(OH)4[Si2O5] (kaolinite) + 2Na + 4SO4 + 6H + 11H2O
2NaAl(SO4)2·6H2O (tamarugite) + 2H4SiO4 (8)
Upon cooling of the solutions and when evaporative conditions prevail in the cave
atmosphere, the silicic acid formed through reaction 8 will dehydrate to form silica (reaction 9).
H4SiO4 SiO2 + 2H2O (9)
Quartz was identified along with tamarugite in several of the powder patterns, thus
confirming our above hypothesis. However, given the circumstances (human impact on the cave
passages) we do not completely exclude the possibility that part of the quartz identified is not
actually authigenic but rather released by weathering from bedrock or concrete.
73 As compared to alunogen, tamarugite is far less common in carbonate cave
environments; its formation in the Diana Cave is due to higher sodium content in the system. The
small size of the crystals may indicate either pulsatory growth conditions (probably dictated by
frequent fluctuations in the springs’ physicochemical parameters) or an insufficient concentration of the necessary ions in the system.
Even though early reports of this highly soluble Na-Al sulfate are from arid environments
(e.g., Palache et al., 1951), the majority of the more recent sites are located in the vicinity of rather high-temperature brines or coastal areas. Delmelle & Bernard (2000) identified tamarugite
-1 2- crusts in volcanic waters with pH < 0.4, temperature > 40° C, and ≈ 70,000 mgL SO4 .
Bortnikova et al. (2005) report on the presence of the same mineral precipitated due to supersaturation upon cooling and pH increase of extremely hot (close to boiling point) and acidic
-1 2- (pH of 1 – 3.5) fumarole-associated thermal brines (9,310 mgL SO4 ). Keller (1935) identified
the mineral in cm-thick crusts on dolomite and attributed its presence to the redox reactions
taking place between meteoric water and carbonate-associated pyrite (pyrite being the source for
both sulfate and H2SO4 necessary to weather the dolomite and release the Na and Al from within
clays). Sea spray has been identified as the culprit in several coastal occurrences of tamarugite
(Hutton, 1970; Segnit, 1976; King, 1998), all connected to pyrite-bearing sedimentary (mostly
carbonaceous) rocks (thus the same general geochemical processes as described by Keller in
1935). The Aghia Paraskevi occurrence (Lazaridis et al., 2011) is indeed the most similar to the
Diana Cave in terms of overall lithology, thermal-water temperature and pH, although sulfide is
not present.
74 Conclusions
Little information exists on the morphology of Diana Cave before it was enlarged and parts of its walls reinforced with concrete; however, the anthropic impact has had a major impact on the cave environment. Thus, tamarugite occurs in two distinct settings: 1) natural, as a result of the steam condensate alteration of the marl bedrock and 2) anthropogenic, following a similar
2- process that affected the concrete reinforced sections of the cave walls. The SO4 ions are a result of the oxidation of hydrogen sulfide escaping from the Diana springs. The steam- condensate alteration causes vigorous weathering of the bedrock (Onac et al., 2009) and plays a major role in the precipitation of the tamarugite-bearing paragenesis by releasing ions from the bedrock and making them available for further chemical reactions.
Although relatively common in mineral assemblages at the surface and within caves affected by past or present volcanic activities, tamarugite is extremely rare in caves developed in carbonate rocks. The only other known occurrence of this mineral is from Aghia Paraskevi caves
(Lazaridis et al., 2011). Based on our observations from Diana Cave, conditions conducive to tamarugite precipitation in carbonate caves seem to be the presence of: 1) thermo-mineral water which through surface evaporation and subsequent condensation on the cold cave walls gives rise to steam-condensate alteration processes; 2) availability of sulfuric acid; and 3) a constant source of Al3+ and Na+.
Diana Cave provides no evidences that tamarugite is a dehydration product of Na-alum or mendozite. The minerogenetic processes that take place in this cave are continuous; consequently, if tamarugite is derived from another mineral species, this precursor should have been identified in at least one sample. Since this is not the case, we consider that tamarugite is a primary by-product of the sulfate acid alteration process.
75 What we find to be indeed special about the Diana Cave tamarugite occurrence is that it uniquely combines elements that are found separately in the above-mentioned sites. That is, the bedrock is carbonaceous (limestone and marl) and provides the Al and part of the Na ions, similarly to the coastal sites, but the sulfate component in the system derives from the conversion of H2S degassing from a Na-rich thermo-mineral water (comparable to those from volcanic areas) of meteoric origin. The δ34S values in the cave sulfates confirm the marine-derived
(evaporites) source of S.
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80 CHAPTER FIVE:
USING δ34S IN LOW-GRADE HYDROTHERMAL WATERS TO FINGERPRINT
HYPOGENE SPELEOGENESIS (BĂILE HERCULANE, ROMANIA)
Introduction
Băile Herculane (Herculean Baths; inset in Fig. 5.1) became a highly regarded spa soon after the Roman conquest of Dacia (ca. 107 AD; Povară, 2011) and continues to function as such in the present. Although surfacing on a relatively short transect along the Cerna River (ca. 40 km), the chemical and isotopic composition of the thermomineral springs varies significantly
(Papiu, 1960; Marin, 1984; Povară et al., 2008). Modern scientific research of the thermal aquifers beneath the lowermost segment of the Cerna River was initiated in the early 1970s and was threefold: understanding the hydrogeology of the region, ensuring the sustainability of the resort, and assessing the potential to harness the thermal energy offered by these waters
(Oncescu, 1953; Pascu, 1968; Povară, 1973; Povară & Lascu, 1978; Mastan et al., 1982; Simion,
1982; Marin, 1984).
In addition to the numerous thermal springs, one other peculiarities of the resort is the presence of hypogene (or hypogene-imprint) caves dotting the surroundings – one of the few places in Romania where this type of speleogenesis has been identified (Sârbu, 1996; Onac,
2005; Onac et al., 2009; Onac et al., 2013; Pușcaș et al., 2013). Some of the caves (e.g., Peștera
Haiducilor and Peștera cu Aburi) in Băile Herculane were inhabited by humans during the
81 Bronze Age and Roman rule, due to the heat of the thermal water and associated steam
(Boroneanț, 2000).
Sulfur (S) – one of the major components of hydrothermal waters – is most often found in
2- - 2- natural waters as sulfate (SO4 ) or sulfide (H2S, HS , and S ), depending on the system’s pH, redox potential, and temperature. The most important reactions involving S in thermal waters are considered to be oxidation of sulfides, hydrolysis of native S (S0), and reduction of sulfate
(Kaasalainen & Stefánsson, 2011). There are several studies that use δ34S in water as a tracer of various natural and anthropogenic processes (Rightmire et al., 1974; Schwarcz & Cortecci, 1974;
Krothe & Libra, 1983; Kester et al., 2003; Yuan & Mayer, 2012). In the Cerna Valley watershed preliminary work was carried out in the Summer of 2008 (Sumrall, 2009; Wynn et al., 2010;
Onac et al., 2011) with the purpose of identifying S sources, the chemical processes it undergoes from the deep seated reservoir to the surface, and to gain further insight into the hydrogeology of the Băile Herculane area. This particular study aims at expanding the field area, number of thermomineral water sources, and duration of field measurements, offering a more in-depth understanding of the intricate hydrogeological and geochemical processes that take place underneath the lower sector of the Cerna River and their connection to hypogene speleogenesis.
Site Description
The geology of this segment of the Southern Carpathians has received special attention over the past few decades, with focus mainly on surface mapping and numerous hydrological assay boreholes, in order to elucidate its tectonic history and magmatic and metamorphic evolution (Kräutner & Krstić, 2002; Iancu et al., 2005; Balintoni et al., 2010). The main geologic formations related to the Băile Herculane thermomineral reservoir pertain to two important
82 domains, namely the Danubian terrane and the Getic/Supragetic terrane (Povară, 2008), comprising pre-Mesozoic high-grade metamorphic and magmatic basement rocks and discontinuous Upper Carboniferous-Permian continental deposits covered by Jurassic-
Cretaceous sedimentary strata (Schmid et al., 1998; Iancu et al., 2005). The sedimentary units consist of alternating conglomerates, sandstones, shales, and mainly thick carbonate sequences
(Năstăseanu, 1980). A significant positive crustal geothermal anomaly (ca. 600 °C at 20 km depth) was identified here and is thought to be the result of the region’s complicated geologic history (Demetrescu & Andreescu, 1994). Next to petrography, the region’s tectonic features are of great importance to the manifestation of the peculiar thermomineral hydrology of the Băile
Herculane area. In our research area (Fig. 5.1) the Cerna River, a tributary of the Danube, flows
NE – SW along the Cerna Graben, a structural feature with a vertical displacement of up to ca.
1000 m. The second major tectonic feature affecting the local hydrology is the Cerna Syncline
(ca. 25 km long), running along the west slope of the Cerna River and intersecting the Cerna
Graben in the vicinity of Băile Herculane. The Cerna Anticline, interposed between the syncline and the graben acts as a hydrologic barrier (Povară et al., 2008). Associated with the three major features, secondary faults are also of great importance, allowing the migration of water and dissolved constituents from neighboring basins (Cosma et al., 2008). These 3 major structural features control the flow of the 3 main categories of water bodies: the Cerna River and its surface tributaries, karstic springs, and the thermal aquifer (Marin, 1984; Povară, 2012). Also, research revealed a significant radioactivity – related to the presence of granites and granite-derived arkoses – which is inversely correlated with the concentration of dissolved solids. The concentration of the dissolved constituents increases from N to S, while radioactivity decreases along the same transect (Michailescu, 1923; Athanasiu, 1928; Papiu, 1960; Cosma et al., 1996).
83 Figure 5.1. Simplified geological map of the area (after Năstăseanu, 1980). Inset A: location within Europe; inset B: site location within Romania. Google map showing the location of the sampling sites along the Cerna River, SW Romania. Southern and Northern aquifers; drainage boundaries of the Cerna River.
84 Hydrogeology and Hydrochemistry
Due to the fame and peculiarity of the thermal springs in the vicinity of Băile Herculane, the regional hydrology, especially the origin of the waters, has been a subject of research since the second half of the 19th century (e.g., Koch, 1872; Pamfil, 1921; Popescu-Voitești, 1921;
Athanasiu, 1939). Initially a juvenile, fumarole-type origin – related to the emplacement of either
the Cerna granite or a Tertiary volcanic stage – was attributed to these springs, based on their
temperature, presumed chemical composition, significant chlorine content and amount of
released gasses (H2S, CO2, CH4, and N2) and water vapor (Koch, 1872; Schafarzik, 1901;
Popescu-Voitești, 1921). In the following decades a second theory (vadose origin) was
elaborated, proposing that meteoric water descends along faults and fractures (down to ca. 2 km)
and became heated due to the local geothermal gradient, later resurfacing along transverse faults
(Athanasiu, 1927; Athanasiu, 1939; Papiu, 1960). The S content of the waters was attributed to
either lagoon-derived sulfate deposits being reduced to H2S in the presence of organic carbon
(Oncescu, 1953), or dissolution of pyrite contained by the sedimentary formations (Papiu, 1960).
Finally, a third model (mixed origin) – accepted today with some amendments – was developed
by Simion (1987), proposing a mixture of local, hot juvenile source, cold meteoric seepage
waters, and an allochthonous mineralized water component (Mitrofan & Povară, 2000; Povară et al., 2008).
More recent research has revealed the presence of 3 distinct aquifers: Northern Complex,
Granite Sill (the correct term we prefer to use is pluton, as the Cerna Granite is not a volcanic body; Iancu, 1976; Iancu et al., 1994), and Southern Complex (Povară et al., 2008). Additionally, the Southern Complex was divided into 4 groups: 1). Hercules (Hercules I, Despicătura, Hygeea,
Apollo I, and II), 2). Diana (Diana I+II, III, IV, and Hebe), 3). Neptun (Neptun I+IV, II, and III,
85 Venera I and II), and finally 4). Southern Group (Traian, Decebal, Fabrica de Var, Vicol, Sera de flori, and Stadion). The grouping of these water sources (Figure 5.1 and Table 5.1) was suggested by Simion (1986) based on the type of occurrence (natural springs and wells), their position along the Cerna River (from North to South), and hydrochemical characteristics. An important feature of the thermomineral sources is the significant amount of gas they carry to the surface – atmospheric, N2, H2S, CH4, and He (Mastan et al., 1982).
Description of the Sampled Springs
The sampled springs and wells belong to all three of the following aquifers: Northern
Aquifer – Pișătoarele, Piatra Pușcată 1, Piatra Pușcată 2, and Crucea Ghizelei; Granite Pluton
Aquifer – 7 Izvoare Calde Dreapta, 7 Izvoare Calde Stânga, and Scorillo; and the Southern
Aquifer – Hercules I, Diana III, Diana IV, Domogled karstic spring, Venera I and II, Neptun II and III, Sera de flori, and Fabrica de Var.
Both Pișătoarele and Domogled are cold karstic springs located on the east bank of Cerna.
The first of these cold karstic springs is surfacing as a diffuse waterfall, whereas the second is collected through a mine gallery within the Domogled Mountain. Piatra Pușcată 1 (PP1) and
Piatra Pușcată 2 (PP2) surface less than 0.5 km from each other on the right bank of Cerna, at riverbed level; therefore, during floods they become submerged.
Crucea Ghizelei is the northernmost borehole thermal source to be used for balneotherapeutic purposes. 7 Izvoare Calde Dreapta (7 ICD) and 7 Izvoare Calde Stânga (7
ICS) are located on the right and left bank, respectively, of the river, in close proximity of the riverbank. 7 ICD is a group of several outlets currently collected through a common pumping station, whereas 7 ICS consists of a group of spring outlets clumped together. Scorillo is also a
86 borehole on the left bank of Cerna. Hercules I is a natural spring, surfacing through a cave; based on its location, geochemistry, and physical parameters it is considered to be the main outlet for the Cerna Syncline aquifer, with an important meteoric component (Povară & Marin, 1984). As is the case for several of the other springs and wells in Băile Herculane, the Diana III well has been fitted with a small concrete basin and a metal spout, giving people access to the water for their treatment. A few feet away is the pumping station and collector basin for spring Diana I+II, surfacing from the Diana Cave. Diana IV is a small pool of milky-green thermal water – perennial and of fairly constant size; ca. 3 m wide by 4 m long – a few meters away from the riverbank on the left side, just across the river from Diana III. Venera I and Venera II are natural springs, the first of which surfaces in the basement of a building, and the second appears on
Cerna’s left bank. The Neptun I+IV well, and Neptun II and III springs are located across the river from the previous two. Sera de flori is also a borehole (identified as the Traian borehole by
Sumrall, 2009; Wynn et al., 2010), whose pumping station is located on the east bank of the river, in the vicinity of two other wells (Traian and Decebal). The furthest south among the springs we sampled is Fabrica de Var (Lime Factory; FV), on the east side of the river.
Materials and methods
Monthly water samples were collected between May 2011 and May 2012 from 18 sources: 15 thermomineral wells and springs, 2 karstic springs, and 1 river; one-time measurements were carried out for two additional thermomineral sources (Fig 5.1; Table 5.1).
87 Table 5.1. General information on the monitored thermomineral springs and wells. For a complete list, see Povară, 2011.
88 On-site measurements were carried out using a Hanna HI 9828 Multiparameter meter
(calibrated before each field session) and included pH, temperature, dissolved oxygen (DO),
conductivity, salinity, oxido-reduction potential (ORP), and total dissolved solids (TDS). In
addition, a Hach DR 2700 spectrophotometer was used on-site to measure the concentrations of
2- - 2- dissolved sulfate (SO4 ) and total sulfide (as H2S, HS , and S ), following Hach’s methods 8131
and 8051 (Hach Company, 2002); SulfaVer 4 powder pillows and Metylene Blue were used as
reagents for dissolved sulfate and sulfide, respectively.
Water samples were collected for stable isotope measurements, namely δ13C (as
18 34 2- 2- dissolved inorganic carbon -DIC), δ O, δD, and δ S (in both SO4 and S ). DIC samples were
filtered through 25 μm Polysulfone GD/X syringe filters (Whatman) and collected in 20 mL clear EPA VOC vials with open-top caps lined with hand-punched black butyl-rubber septa.
Approximately 10 mg of CuSO4 was added to each vial prior to sample collection to act as a
bactericide. Samples were kept in a refrigerator until analyzed. For many of the thermomineral
sources the amount of CuSO4 had to be significantly increased, as sulfide minerals instantly
precipitated out of solution.
2- 2- Based on concentration measurements, between 1 L (for S ) and up to 4 L (for SO4 ) of water were collected with a luer-lock valve polyethylene syringe and prepared – following the method of Mayer & Krouse, 2004 – for filtration of precipitated sulfur species. Filtering was carried out within a few hours of sampling, using NALGENE filters with receivers and a vacuum pump. We used MCE filters with a pore size of 0.45 μm, which ensured an easy recovery of even the minutest amount of precipitate. Resulting precipitates were dried for several hours at 60 °C in a drying oven.
89 Analyses of both waters and solids were carried out by elemental analyzer (EA) and
34 isotope ratio mass spectrometry (IRMS) as follows: for δ S samples (as Ag2S and BaSO4) were
converted to sulfur dioxide gas following the method of Yanagisawa & Sakai (1983). Thermal
conversion (TC/EA) was used to measure δ18O in the solid samples. The DIC, δ18O, and δD
analyses followed CO2, respectively H2 equilibration (see Epstein & Mayeda, 1953; Torres et al.,
2005; Assayag et al., 2006). A number of reference materials were used to check for accuracy
and precision: IAEA-S2 and IAEA-S3 for sulfide, IAEA-SO5 and IAEA-SO6 for sulfate (0.2 ‰ precision; reported to Cañon Diablo Troilite -CDT), NBS-18 for DIC (0.1 ‰ precision; vs.
Vienna Pee Dee Belemnite -VPDB), and 2 in-house standards (VEEN and HTAMP) for δ18O and δD (0.05 and 0.6 ‰ precision, respectively; reported to Vienna Standard Mean Ocean Water
-VSMOW).
Cave sulfates and gypsum precipitated around the thermomineral springs’ spouts were also sampled. Following X-ray diffractometry to identify the mineral species, the samples were analyzed by EA to verify their δ34S and to check against dissolved S species in the
thermomineral waters. Some of the gypsum samples were powdered and dehydrated at 350 °C
for ca. 3 hours and prepared for δ18O measurements by combustion with excess C at 1350 °C in a
TC/EA, precision being 1σ.
Monthly precipitation average amounts obtained from the Băile Herculane
Meteorological Station (obtained from the Romanian National Meteorology Administration;
ANM) were obtained for comparison to the monthly hydrochemical measurements that we carried out in the field.
90 Results
A summary of the monthly chemical and physical parameter measurements carried out
during the 1-year long field campaign is presented in table 5.2 and figure 5.2. Temperature
variations are negligible in all the wells, indicating a lack of infiltration from cold meteoric
waters (Figure 5.3). The Cerna River shows temperature changes that approximate ambient
variations; the karstic springs, Pișătoarele and Domogled, are also relatively stable, as indicated
by measurements that were carried out very close to their discharge points. Among the thermal
springs, Hercules I has the largest scatter, indicating that the system is easily disturbed by cold
meteoric input (Povară & Marin, 1984), Venera II and Diana IV are both located in the proximity
of Cerna’s thalweg and experience various degrees of mixing with river water, as evidenced from temperature measurements. The rest of the thermomineral springs exhibit an amplitude of
<5 °C over the 14 months of measurements.
The river and 2 karstic springs exhibit the most significant change in pH over the period of the study and are also the most alkaline (pH 8 – 9.5); their high alkalinity and pH is explained by buffering due to the limestone bedrock they are in contact with. Among the thermal water sources, those from the Southern Aquifer Complex are overall the most acidic due to the high sulfide concentration (Table 5.2). Those belonging to the Northern Aquifer Complex are intermediate due to buffering by limestone and lower dissolved sulfide, while those on the
Granite Pluton exhibit pH values only slightly below those in the cold waters (Cerna, Domogled, and Pișătoarele), probably also buffered to some extent by dissolution of minerals within the fractured and weathered Cerna granite.
Because the curves for salinity, electric conductivity (EC), and total dissolved solids
(TDS) follow similar trends, figure 5.3 only plots monthly TDS measurements. TDS values
91 below 1000 ppm were measured in the river, karstic springs, the Northern aquifer and Granite
Pluton thermomineral sources. Except for Hercules I, which exhibits a significant scatter (ca.
1000 to 3000 ppm) probably due to its sensitivity to meteoric input, all other thermal waters were
relatively stable over the course of the study. As an exception, Diana III and Venera I exhibit a
few spikes in TDS that we attribute to anthropic activity (construction at the spring site). The
grouping of thermomineral sources as described in Simion (1986) is also evident (Figure 5.3).
Dissolved sulfide concentrations display a very strong variability in the thermal sources.
Over the hydrologic year they are insignificant to nil in the river and karstic springs, as well as
springs and wells in the Northern Aquifer, Granite Pluton (with the exception of 7 ICS and
Scorillo), and the Hercules group. Meanwhile, the other thermal sources in the Southern Aquifer
Complex have significant dissolved sulfide concentrations (generally between 35 and 60 mg/L,
but exceptionally up to 107 mg/L). Regarding dissolved sulfate, springs do not seem to follow
the established groups (i.e., Simion, 1986), instead they cluster in 3 major categories: 1) under 20
mg/L (Pișătoarele, CG, Domogled, Neptun II, Venera I & II, and FV, Cerna); 2) between 20 and
80 mg/L (PP1, PP2, Diana III & IV, Neptun III); and 3) >80 mg/L (7 ICD, 7 ICS, Scorillo,
Hercules, Sera de flori).
13 18 18 34 Detailed stable isotope measurements (δ CDIC, δD and δ O in water, and δ O and δ S in precipitates) were carried out on all of the water samples and cave minerals (Tables 5.3 and
5.4). The only similar data for this region is presented in Wynn et al. (2010) and Onac et al.
(2011) and while there is some variation due to the length of this study and the changes induced by the alternation of seasons (i.e., precipitation amount and ambient temperature), results presented here generally agree with those reported in the above mentioned studies.
92 Table 5.2. Summary of water chemistry for the thermomineral and karstic water springs/wells sampled for this study (May 2011-May 2012).
93 Table 5.3. Results of stable isotopic measurements in water samples and resulting precipitates from all the sources (karstic and thermomineral).
Table 5.4. Results of stable isotopic analysis of mineral samples from caves within the studied area, as well as gypsum deposits from the thermomineral springs.
94 Oxygen and hydrogen stable isotopic ratios are slightly more variable than data from
Wynn et al. (2010), as would be expected given the significantly larger dataset resulted from
sampling 19 water sources over a period of 12 months. δD varies between -80.2 and -68.5 ‰ in
thermal waters (springs and wells) and -74.6 and -55.6 ‰ in the karstic springs and Cerna River.
δ18O values fluctuate between -8.4 ‰ and -11.4 ‰ in the thermal sources, and between -11 and -
8.7 ‰ for karst springs and the Cerna River. Dissolved sulfide concentrations are extremely low
to nil in the Northern Aquifer Complex and the Granite Pluton, as well as in all cold water
34 sources; δ Ssulfide varies from 13.7 to 26.6 ‰. Dissolved sulfate in the thermal waters where
34 mixing with cold meteoric water is not obvious has δ Ssulfate isotopic ratios varying from 17.9 to
39.9 ‰; in karstic waters the range is between 2.4 and 10.3 ‰. There is a more significant
13 variation in δ CDIC between the various sources, with values between -34.4 and -7.1 ‰.
18 δ Osulfate in the mineral samples (mainly gypsum) are generally comparable to those of
dissolved sulfur in the water samples (δ34S ≈ 20 ‰) (δ18O and δ34S; Table 5.4). With the
exception of the gypsum sample from Despicatura Cave (depleted δ18O ratios compared to other
gypsum samples) δ18O values in sulfate minerals are significantly enriched compared to those in
dissolved sulfate or in the water itself for this small sample set of oxygen isotopes in sulfate
minerals.
95
Figure 5.2. Monthly variation in hydrochemical parameters in each of the monitored sources. Monthly rainfall averages for Băile Herculane (ANM, Romania).
96 Discussion
Hydrochemistry
All the thermomineral waters fall into the long-established categories (aquifers and groups) when it comes to hydrochemical parameters (Figure 5.3) (spring averages), with the
exception of dissolved sulfate that tends to be more scattered. As expected, the two karstic
springs (Pișătoarele and Domogled) and the river (Cerna) are markedly different from the
thermomineral waters, with significantly lower temperature, electric conductivity, salinity, total
dissolved solids, dissolved S species, and the highest pH and redox potential. Because of the
depth from which water emerges and the fact that most boreholes are cased, temperature
variations in most wells are exceptionally small. Except for the river, which shows a delayed and
dampened response to seasonal changes, most natural springs (thermal or karstic) also show little
temperature changes throughout the year. However, significant temperature fluctuations are
observed in waters that undergo mixing (with meteoric; e.g., Hercules I, CG, PP1, Diana IV, and
Venera II). pH variability throughout the hydrologic year is generally moderate, and
dissimilarities between sources can be easily explained through a combination of water
chemistry (dissolved species) and the lithology (buffering capacity) traversed by each
spring/well. While TDS variations between aquifers and spring groups are obvious, there is little
change in each individual thermal source within a year – with the exception of Hercules I, which
is susceptible to substantial mixing with surface water. The relative stability – with few
exceptions – of these physical and chemical parameters within each thermomineral spring/well
and within an aquifer or group suggests a homogeneous source and comparatively little
interference from hydrochemically distinct aquifers.
97 Figure 5.3. Average chemical and physical parameters of the studied springs. Thermomineral sources in red, cold springs in blue. Shaded areas represent the spring groups, while the dashed perimeters encompass the aquifer complexes. Numbers on the X axis are consistent with those assigned to the water sources in Table 5.2.
98 PP1 and 2, Diana IV and Venera II surface along the riverbed and as a result are
periodically either flooded by the river or at least undergo partial mixing. Hercules I, although
surfacing several hundreds of meters away from the river, is the spring that shows the most rapid
response to precipitation (within hours; Povară & Marin, 1984) with drastic changes in both
temperature and dilution (Figure 5.2).
Diffuse flow is a second possible explanation for the relative hydrochemical stability of
the Cerna Valley natural springs; according to Shuster & White (1971) the behavior of a
carbonate aquifer is largely determined by regional geology, the two end-members being: 1)
diffuse flow (through joints, fractures, bedding planes) and 2) conduit flow (through solution
passages). Water parameters (temperature, chemistry, and discharge) vary considerably in a conduit spring over a hydrologic year, while they remain relatively constant in springs fed by
diffuse flow – as appears to be the case in the Cerna Valley.
Origin of thermomineral water
18 18 As a result of redox reactions, δ Osulfate is considerably higher than δ Owater.
Thermomineral water δD (ca. -70‰) is close to the global meteoric water line (GMWL), as suggested by karstic water δD; δ18O of thermomineral waters is slightly more enriched (by ca.
2‰) as compared to meteoric δ18O (Figure 5.4). Most of the high-temperature thermomineral sources (e.g., CG, FV, Diana I+II, Diana IV) plot to the right of the GMWL, pointing to evaporative effects.
99 Figure 5.4. Comparison of δ18O and δD of thermomineral springs/wells to δ18O and δD of karstic spring waters. The clustering of the majority of thermomineral waters along the GMWL supports their meteoric origin.
Stable sulfur isotopes
Sumrall (2009), Wynn et al. (2010), and Onac et al. (2011) have shown that the source of
dissolved S in the waters is derived from marine evaporites associated with the limestone in the
aquifer, and that the variations in the S species in the thermal water and cave sulfates is the result
of the reduction – in the presence of methane – of these sedimentary evaporites. The stable
18 34 isotopic ratios we have measured (δ Owater, dissolved sulfate, cave minerals, δ Sdissolved sulfate & sulfide, minerals,
13 and δ CDIC) can be used to identify the source of S and trace the processes that affected it,
allowing us to produce a model that is as close to reality as possible.
Sulfate is reduced – with concomitant oxidation of the electron donor – in the presence of
electron donors following one of two distinct pathways (i.e., thermochemical sulfate reduction –
TSR – and bacterial sulfate reduction – BSR), each governed by a distinct temperature domain.
The temperature range for BSR is from 0 to approximately 60–80 °C, while TSR occurs in the
range of 100–140 °C, (can go up to 160–180 °C in certain settings) each producing specific S
100 isotopic fractionation factors (-30 to -15 ‰ for BSR and -20 to -10 ‰ for TSR, depending on the
temperature) (Machel, 2001; Machel et al., 2005). While TSR occurs inorganically and has
relatively low reaction rates, BSR has fast rates and is organically mediated. The robust
geochemical tool for discriminating between these two processes are comparisons of δ34S, δ13C,
and δ18O values in the waters. Both TSR and BSR rates are limited mainly by the availability of
the chief reactants, sulfate and reactive organic matter (Machel, 2001). If minimal isotopic
fractionation occurs (i.e., S conversions under aerobic conditions) dissolved sulfate may retain
the S isotopic signature of the source rocks. Conversely, significant isotopic fractionation occurs
during bacterially mediated (i.e., dissimilatory) sulfate reduction (Nakai & Jensen, 1964); sulfur
isotopic fractionation that occurs during the oxidation of sulfides or organic sulfur to sulfate is
negligible (Mayer et al., 1995; Schiff et al., 2005).
Given the presence of the significant geothermal anomaly beneath the Cerna Graben,
TSR is a more plausible process than BSR (Wynn et al., 2010). Total δ34S sulfur value (TSV) calculated as an isotopic mass balance (Rajchel et al., 2002; Wynn et al., 2010) is a useful diagnostic tool for identifying (or at least narrowing down) the source of S (e.g., marine evaporites, sedimentary sulfides, igneous sulfur, or anthropic sources) assuming a closed system and that the sources have different δ34S values. δ34S TSV for the Cerna Valley thermomineral
springs varies between 18 and 30 ‰ (slightly lower for the mesothermal springs with a
significant meteoric input); for karstic springs it varies from 3 to 10 ‰, and for the Cerna River
δ34S TSV varies from 3 to 16 ‰ (Figure 5.5). The TSV observed in karstic springs and the Cerna
River suggest that their S is most likely derived from atmospheric input or from sulfides (e.g.,
disseminated pyrite) or sulfates (e.g., disseminated gypsum) within the sediment or bedrock
(Cortecci & Longinelli, 1970; Mayer et al., 2010); this surface input is periodically also
101 perceptible in mesothermal springs, especially after heavy rainfalls (e.g., PP1, PP2, Hercules I).
Conversely, the TSV range of thermomineral water samples advocate for a deep seated input,
namely of marine sedimentary origin; typical δ34S values in Jurassic and Cretaceous sulfate in
the ocean is in the range of +17 to +20 ‰ (Bottrell & Newton, 2006). However, an additional
explanation is needed for the significantly heavier (+20 to +30 ‰) TSV measured during certain months in waters of the Southern and Granite Pluton aquifers. A potential explanation is TSR limited by the availability of S, which would lead to enriched δ34S as compared to the source sulfate, unless TSR goes to full completion.
Figure 5.5. Monthly variation in δ34S TSV of all sampled sources along a N – S transect.
Oxygen stable isotopes
In addition to measuring δ34S of sulfate (dissolved or solid) – which can pinpoint the
source of S, analyzing the δ18O of sulfate will contribute to identifying the geochemical
102 processes that affected the sulfate during its transit toward the surface. The δ18O of the resulting
sulfate is a function of reaction pathways, the relative contribution from water, air, and
environmental conditions (pH, temperature, oxidizing agent, etc.) (Van Stempvoort & Krouse,
18 34 1994). Changes in sulfide oxidation pathways will affect δ Osulfate, leaving δ Ssulfate virtually unaffected, resulting in a lack of covariation between the two (Turchyn et al., 2009). Figure 5.6
34 18 shows a plot of δ S and δ O of dissolved sulfate as a function of spring location from N to S.
18 34 Figure 5.6. Comparison of δ Osulfate and δ Ssulfate (monthly averages) on a N – S transect. Red ellipses represent thermomineral springs, the blue one groups the shallow cold waters.
There is a significantly larger scatter in δ34S as compared to δ18O, suggesting different extents of
34 TSR. The Cerna River and the karstic springs display distinctive δ Ssulfate signatures (+4 to
+7 ‰), suggesting surface S sources (e.g., atmospheric, anthropic, etc.) or significant mixing
18 (especially for the river); δ Osulfate of karstic waters is ca. +4 ‰, advocating for a mixed
103 atmospheric and meteoric origin. Two groups of thermomineral springs are evident
(differentiated mostly by the range of δ18O) and two outliers: Diana III (highest δ34S) and Diana
IV (lowest δ18O).
According to Van Stempvoort & Krouse (1994) in the case of BSR, δ18O in resulting sulfate will be 10 to 20 ‰ heavier than the water where it as oxidized, while in the case of TSR
18 fractionation factors are lower, producing δ Osulfate close to that of the water. One other potential
18 explanation for the observed heavier δ Osulfate is a contribution of atmospheric O during
oxidation (Seal, 2003).
Dissolved inorganic carbon
Based on δ13C values of DIC, methane – present in many of the thermomineral sources –
is a plausible electron donor for TSR in the Cerna Valley, plotting the relation between δ13C of
DIC and δ34S of sulfate (5.7) can be useful in tracking the extent of TSR completion in these
waters. The Northern Aquifer (thermomineral springs PP1, PP2, and CG) appears to be a
separate group, apparently unaffected by TSR, with δ13C very close to that measured in karstic
springs; thermal sources in this aquifer – including CG – must represent (as suggested by tracer
studies; Povară, 2008) heated karstic waters, with a low mineralization –including extremely low
S.
Thus, waters in the Granite Pluton and Southern aquifers are the only ones undergoing
TSR; 7 ICS and Hercules I appear to witness the inception of TSR, with progressively more negative δ13C values than that of the karstic waters (-10.8 ‰); in the case of Hercules I TSR may
be periodically masked by significant meteoric input, resulting in dilution of the thermomineral
component and mixing with waters carrying a diverse isotopic signature. Diana III shows the
104 most extreme values, in both δ13C and δ34S, indicating TSR close to completion. While figure 7
suggests that the development of TSR does not increase progressively from N to S, it must be
mentioned that tracer studies showed meteoric mixing in almost all of the thermomineral springs
(PP1 and 2, 7 ICD, Hercules I, Diana I+II and IV, Neptun II and III; Povară, 2011) which might
affect or mask other processes. Our results – covering between 5 and 13 months – are in general
agreement with data from Wynn et al. (2011).
Figure 5.7. The relationship between δ34S and δ13C of the thermomineral sources, along a N – S transect. The progression of TSR increases from an incipient stage in Hercules I toward Diana III, where it approaches completion.
Stable isotopes in cave minerals
In many of the caves affected by SAS, δ34S of sulfate minerals was found to be significantly δ34S depleted compared to the δ34S of the sources of sulfur (Hill et al., 1987).
However, cave minerals with negative δ34S are not enough to discriminate between TSR and
BSR or prove that the cave was impacted by sulfuric acid. Cave sulfates from two of the caves
presently affected by thermomineral water/steam and gypsum precipitated around some of the
105 thermomineral springs’ spouts (Table 4) are used here for comparison with the stable isotopic
ratios we measured in the waters. δ34S in all but two of the samples is relatively close to the
values measured in dissolved sulfide from the corresponding water. These small variations are
due to either changes in the stable isotopic composition of dissolved S, or different fractionation
factors for the sulfate minerals. The two outliers are a gypsum sample (δ34S +13.5 ‰) from
Despicătura Cave – where thermomineral waters frequently mix with meteoric input – and one
34 18 from the concrete basin of Sera de flori well (δ S +5.8 ‰). The δ Osulfate signature in the
18 gypsum sample from Neptun II spring is just slightly more enriched than the water δ Osulfate.
This indicates that the gypsum from Depicatura Cave resulted from mixing with a meteoric source, which periodically discharges overflow from Hercules I. Our results support those of
Onac et al., (2011) and make the connection between coexisting thermomineral waters and
34 18 resulting sulfate minerals. Onac et al., (2011) also report δ Ssulfate and δ Osulfate for minerals in
caves further N of Băile Herculane and at higher altitudes, but we found no correspondence
between present day thermomineral waters in the same areas and stable isotopes in the minerals.
Conclusions
In addition to being geologically intriguing, these thermomineral sources are
economically important for the region, attracting large numbers of tourists. Their curative
properties are acknowledged by physicians and pumping stations deliver the water from several
of the wells directly to some of the treatment centers. As a result, there is an understandable
concern for maintaining their therapeutic qualities by identifying and quantifying mixing with
cold meteoric waters, as well as finding engineering solutions for preventing it. Using stable
106 isotopes in the waters in the Cerna Valley emerges as a useful new tool for quantifying the extent of dilution of thermomineral waters by meteoric input.
Here we offer improved results via more modern δ34S, δ13C, δ18O, and other analyses for spring and well waters from Băile Herculane. Although intense research was carried out around
Băile Herculane for over half a century, earlier work is somewhat scattered and noncontinuous with very frequent analyses during certain periods, followed by years of no measurements for some of the less important (i.e. low discharge) springs. Significant amount of data has either remained unpublished, or was published in obscure reports. One other issue in comparing our results to previous work is that measurements were carried out using various methods (not always described and sometimes empirical) and precision is not always reported in literature.
Nevertheless, the major findings of the monthly measurements of the physical and chemical parameters presented here are in agreement with results from similar measurements in earlier literature (summarized in Povară, 2011). Sources belonging to the Northern aquifer complex display low TDS, salinity, EC, and temperature, while those from the Southern aquifer complex have significantly higher values for the same parameters (Fig. 5.2). Our monthly measurements
(in accordance with previous studies) indicate that several of the natural springs and wells are impacted by mixing with surface water (i.e., meteoric input), resulting in a decrease in temperature, dissolved S, TDS, EC, and also affecting the stable isotopic values.
The data presented here outline the link between thermomineral waters and resulting minerals in caves situated close to the present day regional base level. The ability of the thermomineral waters to precipitate gypsum replacement crusts in surface (spring spouts) and cave (Diana Cave) settings and thick gypsum deposits (Despicătura Cave) is an indication that they are capable of partially or entirely creating the caves as well. The S stable isotopic
107 signatures in thermomineral waters paired with their physical and chemical characteristics can be
used to reconstruct the processes that affected S from reservoir to the final product. Stable
isotope geochemistry coupled with hydrochemical data provides a useful tool for the identification of S sources and cycling in complex aquifers. Additional comparison with cave mineral stable isotopic ratios allows for a better discrimination between speleogenetic pathways, as well as fingerprinting SAS. A hydrologic connection between the N and S aquifers is probable and the Granite Pluton acts as a somewhat permeable boundary due to the highly fractured and weathered rock. Water mineralization appears stronger in springs/wells from the Southern
Aquifer and average temperatures are also higher compared to the Northern Aquifer, raising the
possibility of a southward migration of the geothermal anomaly.
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-----Original Message----- From: Dr. A. Naegele [mailto:[email protected]] Sent: Friday, September 27, 2013 2:41 AM To: Burke, Merilyn Subject: Re: [www] Permission to use an article in a dissertation
Dear Ms. Burke,
thanks for your inquiry. Your student may of course put this article in his dissertation. Nevertheless we only grant permission to use the final submitted manuscript version, the original publishers' pdfs may not be used.
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I have a student who has published an article in European Journal of Mineralogy: Pușcaș;, C.M., Onac, B.P., Effenberger, H.S. & Povară, I., 2013. Tamarugite formed by acid-sulphate alteration in a typical karst setting: case study from Diana Cave, Romania. European Journal of Mineralogy. DOI: 10.1127/0935-1221/2013/0025-2294. She now has to put his article in his dissertation and understands that he may need to obtain permission in order to use his article. Do you have a policy concerning this issue, and is there a policy as to the correct citation that your journal may request to be used in citing the article if permission is granted?
Thank you, Merilyn Burke USF Tampa Library --Sent to: Schweizerbart publishers Company: University of South Florida Language: English Referrer: http://www.schweizerbart.de/home/contact?af=footer ______Dr. A. Nägele E.Schweizerbart'sche Verlagsbuchhandlung OHG (Naegele u. Obermiller) Science Publishers Johannesstr. 3A 70176 Stuttgart, Germany Tel. +49-711-351456-0 / FAX +49-711-351456-99 www.schweizerbart.de [email protected] Amtsgericht Stuttgart, HRA 1864
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