Field Course on Fluid-Rock Interaction in Shallow Hydrothermal Systems

E N e R A G Field Course

on Fluid-Rock Interaction in Shallow Hydrothermal Systems Program and ield guide

Lead by: Co-organised by: Hosted by:

József & E rzsé bet Tóth

Endowed Hydrogeology Chair Foundation

József & Erzsébet Tóth Geological Survey of Finland Endowed Hydrogeology Chair Foundation Eötvös Loránd University Department of Geology Hungary 27-31 May 2019

Horizon-2020 Excellency Network Building for Comprehensive Research - ENeRAG

Field Course on Fluid-Rock Interaction in Shallow Hydrothermal Systems – Program and field guide

Ferenc Molnár, Gabriella B. Kiss, Judit Mádl-Szőnyi, Anita Erőss, Andrea Mindszenty

Edited by: Gabriella B. Kiss Cover by: Ildikó Erhardt, Ádám Tóth, Tímea Havril Publisher: József and Erzsébet Tóth Endowed Hydrogeology Chair and Foundation

ISBN 978-615-81297-1-8

ENeRAG project has received funding from the European Union’s Horizon 2020 research and innovation programme under Grant Agreement No. 810980

Horizon-2020 Excellency Network Building for Comprehensive Research - ENeRAG Fluid-rock interaction in shallow hydrothermal systems Field course program May 27-31, 2019 - Hungary

27 May, Monday

Place: Room no. 0-817 (Dudich Endre room), Eötvös Loránd University (ELTE), Budapest, Pázmány Péter s. 1/C

8.00 - 9.00 Registration 9.00 – 9.10 Welcome and Introduction Assoc. Prof. Judit Mádl-Szőnyi, ELTE, scientific coordinator of the ENeRAG project 9.10 – 9.20 Introduction to the field course – practical, technical and safety issues Res. Prof. Ferenc Molnár, Geological Survey of Finland (GTK) 9.20 – 10.00 Introduction to the geology and structural evolution of the Pannonian Basin and its surroundings Res. Prof. János Haas, Hungarian Academy of Sciences 10.00 – 10.20 Mesozoic submarine volcanism and related hydrothermal systems in North Eastern Hungary Assist. Prof. Gabriella B. Kiss, ELTE 10.20 – 10.40 Coffee break 10.40 – 11.20 Paleogene and Neogene volcanism and related hydrothermal systems in Hungary Res. Prof. Ferenc Molnár, GTK 11.20 – 12.00 Geological evolution and fluid-rock interaction in the Buda Thermal Karst Prof. Andrea Mindszenty, ELTE 12.00-13.00 Lunch break 13.00 – 14.30 Introduction to fluid-rock interaction in terrestrial volcanic-hydrothermal systems: mineralogy and geochemistry Res. Prof. Ferenc Molnár, GTK 14.30 – 15.00 Introduction to fluid-rock interaction in submarine volcanic-hydrothermal systems: mineralogy and geochemistry Assist. Prof. Gabriella B. Kiss, ELTE 15.00 – 15.30 Coffee break

15.30 – 16.00 Numerical simulation of hydrothermal fluid flow in the Buda Hill area over geological time scales Post. Doc. Timea Havril, ELTE 16.00 – 16.30 Introduction to fluid rock interaction and fluid flow patterns in the Buda Thermal Karst Senior Researcher Anita Erőss, ELTE 16.30 - 17.00 Concluding remarks, preparation for the field trip Instructors, ELTE & GTK (Ferenc Molnár, Anikó Lovász)

28 May, Tuesday

7.30 - 08.00 Gathering at the ELTE Campus, Danube entrance at the ELTE Pázmány Péter s. 1/C building 8.00 Departure to the Velence Mts. 09.00 – 10.30 Field stop 1: Nadap, quarry: observations on hydrothermal alteration of a Variscan granite intrusion and Paleogene andesite necks 10.30 – 11.00 Travel to Pázmánd 11.00 – 11.45 Field stop 2: Pázmánd Zsidó Hill: observations on advanced argillic alteration with pyrophyllite in Paleogene andesite 11.45 Departure to the Buda Hill with 20 minutes technical break at a parking lot on the motorway M7, lunch on the bus 14.00- 14.20 Arrival to the Gellért Hill Field stop 3: Panoramic view from the top of Gellért Hill to the geology of the Buda Thermal Karst area 14.30 – 17.30 Visit of the springs, wells and caves in the Gellért Hill area Field stop 4: Gellért Spa building: overview of the recent hydrogeology of the Buda Thermal Karst area Field stop 5: Ősforrás Spring of the Gellért Spa Field stop 6: Gellért Tunnel: Wells and in situ experimental study of the chemistry of discharging thermal water and related precipitates Field stop 7: Aragonite Cave: Ancient thermal water related phenomena in the geological record

29 May, Wednesday

7.30 - 08.00 Gathering at the ELTE Campus, Danube entrance at the ELTE Pázmány Péter s. 1/C building 8.00 Departure to Recsk 10.30 – 10.45 Technical break at the gas station at Parádfürdő 10.45 - 11.10 travel to Recsk, Lahóca Hill 11.10 - 13.10 Field stop 1: Recsk, Lahóca Hill, northern gulley, advanced argillic alteration in crystal tuff 13.10- 13.30 travel to Parádfürdő 13.30 -14.30 Field stop 2: advanced argillic alteration and silicification in the Recsk intrusive- volcanic complex, Etelka pit, Parád 14.30 – 14.45 travel to Parádfürdő, Ilona Valley, Timsós pit 14.45 – 15.30 Field stop 3: Ilona Valley, Timsós pit, hydrothermal breccia 15.30 – 15.45 travel to Darnó Hill, Nagyrézoldal quarry

15.45 – 16.45 Field stop 4: Darnó Hill, Nagyrézoldal quarry, submarine hydrothermal alteration of Mesozoic pillow basalt 16.45-17.45 travel to accommodation, Eger Facultative program: thermal spa, Eger

30 May, Thursday

7.30 – 8.00 Gathering for departure 8.00 – 11.00 Travel to Vizsoly, tuff quarry, technical break at Szikszó 11.00 – 11.30 Field stop 1: Vizsoly, tuff quarry, fresh, poorly welded ignimbrite with gas-emanation channels 11.30 – 12.15 travel to Mád, Kerektölgyes 12.15 – 13.45 Field stop 2: Mád, Kerektölgyes, hot-spring fed lake environment with silica pods and argillites 13.45 – 14.30 travel to Mád, Király Hill 14.30 – 16.00 Field stop 3: Mád, Király Hill, zonal hydrothermal alteration with alunite in a steam heated hydrothermal alteration zone hosted by ignimbrite 16.00 – 17.30 Travel to accommodation, Telkibánya

31 May, Friday

7.30 - 8.00 Gathering for departure (Telkibánya) 8.00 - 9.00 travel to Pálháza, Perlite quarry, with technical break before entering to the quarry 9.00 – 11.30 Field stop 4: Pálháza, Perlite quarry, perlite and peperitic perlite (hot rhyolite lava interaction with water-saturated sediments) 11.30 – 12.15 travel to Rudabányácska, Bányi Hill 12.15 -13.30 Field stop5: Medieval gold mining pits and K-feldspar alteration in acidic tuff, Rudabányácska, Bányi Hill 13.30 – 13.45 travel to Sárospatak, Botkő 13.45 – 14.30 Field stop 6: Sárospatak, Botkő, silicification with cinnabar mineralisation 14.30 – 18.00 travel to Budapest, arrival to the ELTE

Location of field stops and accommodations of the field course (source of the base map: tudasbazis.sulinet.hu).

Field guide table of contents

Day 1 stops 1-2: Hydrothermal alteration zones in the Variscan granite and Palaeogene volcanic rocks in the Velence Mts., Western Hungary ...... 7 1. Geology of the Velence Mountains ...... 7 2. Hydrothermal alteration zones in the Velence Mts...... 8 2.1. Mineralisation related to the emplacement of Variscan monzogranite ...... 9 2.2. Mineralisation in the Variscan granite related to fluid mobilisation processes of Triassic age ...... 10 2.3. Mineralisation related to the andesitic volcanism of Palaeogene age in the Velence Mts...... 11 3. Field stops ...... 12 3.1. Field stop 1: Tertiary hydrothermal alteration zones in the Variscan monzogranite, quarry at the village of Nadap ...... 12 3.2. Field stop 2: Hydrothermal alteration in the Palaleogene intrusive-volcanic complex of the Velence Mts., Zsidó Hill, near the village of Pázmánd ...... 15 4. References ...... 18 Day 1 stops 3-7: Hydrogeology and thermal water related phenomena in the Gellért Hill area – Introduction to the field trip in the Buda Thermal Karst area ...... 20 1. The hydrogeological environment of the Buda Thermal Karst area ...... 20 1.1 Geological settings ...... 20 1.2. Flow regimes ...... 23 2. Field stops in the Gellért Hill area ...... 27 2.1. Field stop 3: Panoramic view from Gellért Hill ...... 27 2.1.1. Geological build-up of the Gellért Hill ...... 27 2.1.2. Field stop 4: Gellért Spa building, overview of the flow regime ...... 29 2.2. Field stop 5: Ősforrás (Spring of the Gellért Spa) ...... 32 2.3. Field stop 6: Gellért Tunnel: Wells and in situ experimental study of the chemistry of discharging thermal water and related precipitates ...... 33 2.4. Field stop 7: Aragonite Cave ...... 35 3. References ...... 37 Day 2 stops 1-3: Hydrothermal alteration zones in outcrops of the porphyry-epithermal system of Palaeogene age at Recsk...... 43 1. Introduction – regional geology and types of mineralisation ...... 43 2. Geology and mineralisation of the Recsk Ore Complex ...... 45 3. Field stops in the epithermal zones of the Recsk Ore Complex ...... 49 3.1. Field stop 1: advanced argillic alteration zone in the HS type epithermal system on the Lahóca Hill, northern gulley ...... 49

3.2. Field stop 2: Etelka pit and adit, northern slope of the Veresvár Hill ...... 50 3.3. Field stop 3: hydrothermal breccia dike, Timsós pit, Ilona Valley ...... 51 4. References ...... 53 Day 2 stop 4: Darnó Hill, Nagy-Rézoldal quarry: submarine hydrothermal alteration of Triassic pillow basalt ...... 55 1. Regional geology ...... 55 2. Proximal and distal volcanological facies in a submarine basaltic lava flow ...... 56 3. Submarine hydrothermal processes in the Darnó Unit ...... 57 4. Field stop 4: the Nagy-Rézoldal quarry ...... 58 5. References ...... 59 Days 3-4: Field stops in the Tokaj Mts...... 62 1. Geological setting of the Tokaj Mts. in the frame of the Neogene-Quaternary Carpathian Volcanic Arc ...... 62 2. General geology and mineral deposits of the Tokaj Mts...... 63 2.1. Geology and volcanism ...... 63 2.2. Hydrothermal systems, fluid-rock interaction, and mineral deposits in the Tokaj Mts...... 69 3. Field stops in the Tokaj Mts...... 71 3.1. Field stop 1: Fresh ignimbrite with gas-emanation channels: quarry at Vizsoly village .. 71 3.2. Field stop 2: Steam-heated alteration zone in ignimbrite: Király Hill north of the village of Mád ...... 73 3.3. Field stop 3: Silica and clay deposits in a hot-springs fed lacustrine basin west of Mád ...... 78 3.4. Field stop 4: Perlite pit, Gyöngykö Hill at the village of Pálháza ...... 79 3.5. Field stop 5: Adularia-sericite alteration in pumiceous rhyolite/crystal tuff: historic gold mining site at the village of Rudabányácska ...... 82 3.6. Field stop 6: Strongly silicified steam-heated alteration zone on the Botkö Hill, northwest of the town of Sárospatak ...... 84 4. References ...... 88

Day 1 stops 1-2: Hydrothermal alteration zones in the Variscan granite and Palaeogene volcanic rocks in the Velence Mts., Western Hungary1

Ferenc Molnár (Geological Survey of Finland)

1. Geology of the Velence Mountains

The Velence Mountains consist of about 200 km2 of outcrops of Paleozoic granitoid and Paleogene intermediate volcanic rocks (Fig. 1.) in the central part of the Transdanube region of Hungary, 50 km southwest of Budapest. The S-type Variscan (280-300 Ma) postcollision biotitic monzogranite intrusion (Buda, 1985) and its granite-porphyry and aplite dikes in the western part of the Velence Mts. is hosted by Palaeozoic shales that show contact metamorphism along the igneous body (Jantsky, 1957). Recent studies suggest a Permian age (283±5 Ma) of magmatism, based on zircon U-Pb dating method (Uher et al., 2010). The monzogranite crystallised at about 520-700°C from a water-saturated eutectic melt at an estimated pressure of about 2 kbars (Buda, 1985, 1993). Very few and thin monchiquite-spessartite and carbonatite dikes of Cretaceous age also occur in the granitoid intrusion (Horváth & Ódor, 1984). Calc- alkaline andesitic rocks (lava flows and pyroclastic deposits) of Palaeogene age (Upper Eocene- Lower Oligocene) crop out in the eastern part of the Velence Mts. (Fig. 1.). Exposures of the highly altered andesitic rocks form a half circle which probably represents an eroded southern remnant of a caldera-like palaeovolcanic centre (see inset structural sketch map on Fig. 1.). Under the remnants of the volcanic structures, deep drillholes proved the presence of a subvolcanic diorite intrusions (Dudko et al., 1989). The area of Palaeogene volcanic rocks is separated from the Variscan granitoid intrusion along a NW-striking fault (Nadap-Lovasberény Fault); however, dikes and small stocks with andesitic composition of Palaeogene age also transect the granitoid body (Fig. 1.). Paleogene volcanic units in Hungary are scattered in a ca. 300 km long and 30-50 km wide NE-SW trending belt along the Periadriatic-Balaton Line (Fig. 2.). Intrusive-volcanic complexes crop out in the Velence Mts. and near Recsk in NE-Hungary, but they are covered elsewhere. Other pyroclastic deposits of similar age are also known from deep hydrocarbon exploration drillings. Basaltic-andesite, andesite, dacite and diorite of intrusive-volcanic complexes have medium- to high-K calk-alkaline character. REE-patterns, Sr-, Nd-Sm- and Pb-isotope ratios are compatible with subduction-related origin (Salters et al., 1988; Benedek, 2002). Biostratigraphy and K-Ar dating (44-27 Ma) indicate that magmatism took place in Late Eocene - Early Oligocene The Paleogene volcanic belt of Hungary can be correlated with the syn- to post-collisional Periadriatic igneous belt in the Alps and volcanic-intrusive units in the north-western part of the Sava-Vardar belt (Benedek 2002, Pamić, 2002). The Palaeogene Volcanic Belt of Hungary and its pre-Tertiary basement is in an allochtonous position. The present day geological structure of Hungary is dominated by the Pannonian Basin which is located in the back-arc zone of the oblique subduction-collision system of the Carpathians (Fig. 2.). The basement consists of two major composite terranes with contrasting Triassic-Middle Cretaceous sedimentary sequences and they are separated by the Mid- Hungarian Line (MHL) and the Periadriatic-Balaton Line (PBL) strike slip faults (Fig. 2A). The ALCAPA block in the north includes the Transdanubian Unit (TU) which has comparable Mesozoic sequences with the Southern Alps, whereas the Tisza Megaunit in the south is

1 This chapter of the guidebook is based mostly on Molnár et al. (2003), Molnár (2004), Bajnóczi (2004) and Benkó et al. (2014). 7 characterised by Mesozoic sequences deposited in the northern (European) margin of the Neotethys Ocean (Kázmér & Kovács, 1985; Csontos, 1995). The strongly tectonised zone between MHL and PBL also contains blocks of Dinaridic origin (e.g. Bükkium). The pre-Tertiary continental units had been juxtaposed along MHL during the Late Oligocene and Miocene as a result of collision between Africa (Adriatic plate) and Europe. TU was located between the Eastern and the Southern Alps at about 30 Ma ago; it has been displaced along PBL to northeast in the order of 400 to 500 km and rotated westward as part of ALCAPA (Kázmér & Kovács 1985, Csontos 1995). The ‘continental escape’ of ALCAPA and Tisza from the Alpean collision zone initiated the Carpathian subduction and collision between the northeastern and eastern margins of escaping micro-continents and down-going Europe (Fig. 2B).

Fig. 1. Outcrop map and structural sketch map of the eastern part of the Velence Mts. Locations of field stops are also marked on this map.

2. Hydrothermal alteration zones in the Velence Mts.

The monzogranite blocks and granite porphyry dikes are characterised by strong hydrothermal alteration in several zones of the Velence Mts. Quartz veins with Pb-Zn ore and fluorite-quartz veins also occur in the central and northern parts of the granite mass and were subjected to mining during the 1950`s. In the eastern part of the Velence Mts., mineral exploration by the Geological Institute of Hungary revealed that the whole intrusive-volcanic complex is strongly altered and carries showings of porphyry copper and gold mineralisation.

Fig. 2. Tectonic setting of the Hungarian Palaleogene Volvanic Belt (A) and tectonic evolution of the basement of the Pannonian Basin during the Alp-Carpathian orogeny (B). The intrusive-volcanic complexes in the Hungarian Palaeogene Volcanic Belt are relatively well preserved because the belt “escaped” from the between collision zone between the Adria (African) and European plate by northeastward extrusion of basement units. That parts of the post-collisional (slab breakoff related) igneous belt which remained in the Alpean collision zone are much deeply eroded, well below the intrusive-volcanic level, due to the intense uplift in that zone.

2.1. Mineralisation related to the emplacement of Variscan monzogranite

Small pegmatite bodies with maximum a few tens of m3 volume occur as pocket- and vein- like formations in all parts of the Variscan monzogranite intrusion (Jantsky, 1957). Their primary mineral assemblage mostly consists of quartz, K-feldspar, plagioclase, biotite and muscovite. Widely distributed miarolitic cavities im the granite intrusion are also characterised by similar mineralogy, though occurrences of tourmaline, garnet, molybdenite, apatite, epidote, fayalite and grünerite accessories have also been reported at some localities (Nagy, 1967, Buda, 1993). Quartz-tourmaline metasomatic-stockwork mineralisation occurs along the contact zone between the intrusion and Palaeozoic shales (Vendl, 1923, Jantsky, 1957). In those tourmaline- rich zones, veinlets also contain native gold, molybdenite, as well as wolframite and scheelite showings at some places (Böjtös-Varrók, 1966, 1967).

Quartz-molybdenite stockwork mineralisation occurs along the eastern margin of the monzogranite body (south of the Field stop 1, Fig. 1.) and it is hosted partly by the intrusion and partly by the contact-metamorphosed shale (Jantsky, 1957). The ore was explored in an inclined adit and results proved that it is non-economic with an average grade of 0.03 % Mo (Jantsky, 1957). In the siliceous veinlets, molybdenite is associated with pyrite, pyrrhotite, galena, sphalerite, Zn-rich tetrahedrite, chalcopyrite and marcasite (Tokody, 1944, Jantsky, 1957, Molnár, 1997). Host rocks show weak silicic, kaolinitic, sericitic and chloritic alteration along the veinlets.

2.2. Mineralisation in the Variscan granite related to fluid mobilisation processes of Triassic age

Quartz-fluorite-barite-sulphide veins with variable proportions of the major minerals (e.g. from fluorite-poor quartz-sulphide veins to sulphide poor fluorite-quartz veins) are widely distributed mostly in the western part of the monzogranite intrusion. The width of veins is up to several metres and they can be followed for a few hundred metres along their most typical northeast-southwest and north-south strike-directions. Orientation of veins largely parallels with the strike-direction of granite-porphyry and aplite dikes. Outcrops of these veins are leached due to weathering and show vuggy-cellular structure of fine-grained quartz with numerous secondary minerals (Koch, 1985). Below the oxidation zone, the most common sulphide minerals of collomorph-banded and brecciated veins consist of sphalerite, galena, pyrite, chalcopyrite, grey ore minerals and stibnite. Monzogranite along the veins has been subjected to beresitic alteration (intense silicification with sericitic K-feldspar and chloritic biotite; Kiss, 1954). Sulphide ores from the veins were mined near Pátka in the northern parts of the Velence Mts. (Fig. 1.) where the grade of ore was 1.23 % Pb and 4.81 % Zn. Barite is a common gangue mineral in some veins only (southern part of the Meleg Hill, Fig. 1., Erdélyi, 1939, 1952), and mostly in their shallow parts. Fluorite forms abundant masses of coarse grained pinkish and greenish aggregates in some veins especially north of Pákozd (Fig. 1.). Total production was

8040 tonnes of high grade (92-95 % CaF2) and 10 573 tonnes of low grade (10-20 % CaF2) fluorite (Jantsky, 1966). On the basis of REE content of fluorite Horváth et al. (1989) suggested that formation of veins may also be related to a hydrothermal activity during emplacement of monchiquite-spessartite and carbonatite dikes of Cretaceous age. However, those dikes are not widespread and are very thin with no alteration around them in the granite; thus it is unlikely that the weak Mesosoic magmatic event may have generated important hydrothermal activity resulting in formation of several metres thick hydrothermal veins. Current systematic, regional scale studies on clay mineralogy in the alteration zones of granite along the hydrothermal veins revealed that most commonly R3 (highly crystalline) illite is associated with the mineralised hydrothermal veins (Benkó et al., 2014). The crystallinity of illite indicates around 250 °C maximum temperature of the hydrothermal fluids. K/Ar dating of these illite samples provided ages of 209—232 Ma, supporting the Mid- to Late-Triassic age of the hydrothermal fluid flow. Fluid inclusion data from the vein quartz and secondary fluid inclusions from the rock forming quartz in the alteration zones around the veins suggest that mixing of low salinity (<12 NaCl equiv. wt%) NaCl-H2O type fluids and high salinity (10-

26 CaCl2 equiv. wt%) Ca-rich fluids took place during this hydrothermal acivity. Regional distribution of densities of fluid inclusion planes in rock forming quartz suggest hydrothermal circulation was regional in the granite, but more intensive around the mineralised zones. Lead isotope signatures of hydrothermal veins in the Velence Mts. (206Pb/204Pb = 18.278—18.363,

207Pb/204Pb = 15.622—15.690 and 208Pb/204Pb = 38.439—38.587) indicating upper crustal source of lead (and other metals). The nature of mineralising fluids, age of the fluid flow, as well as lead isotopic signatures of ore minerals point towards a genetic link between epigenetic carbonate- hosted stratiform-stratabound Alpine-type lead-zinc-fluorite deposits in the Southern and Eastern Alps and the studied deposits in the Velence Mts. In spite of the differences in host rocks and the depth of the ore precipitation, it is suggested that the studied deposits along the Periadriatic-Balaton Lineament in the Pannonian Basin and in the Alps belong to the same regional scale fluid flow system, which developed during the advanced stage of the opening of the Neotethys Ocean. The common origin and ore formation process is more evident considering results of large-scale palinspastic reconstructions. These suggest, that the studied deposits in the central part of the Pannonian Basin were located in a zone between the Eastern and Southern Alps until the Late Paleogene and were emplaced to their current location due to northeastward escape of large crustal blocks from the Alpine collision zone (Fig. 2.).

2.3. Mineralisation related to the andesitic volcanism of Palaeogene age in the Velence Mts.

Porphyry-copper type mineralisation occurs at 880 to 1070 m depth in the Pd-2 drillhole (located near of the Field stop 2, see Fig. 1.) in a subvolcanic intrusion of diorite composition. The host diorite is characterised by potassium-silicate (K-feldpsar, phlogopite, sericite) and propylitic (epidote, chlorite) alteration (Darida-Tichy, 1987). Quartz stockwork contains pyrite, chalcopyrite and sphalerite. Disseminated ore consists of chalcopyrite, bornite, pyrite, magnetite and ilmenite in strongly silicified zones. Locally, at 670 m depth, minor skarn with andradite, epidote, quartz and pyrite is also present. High sulphidation type epithermal mineralisation is widespread in the outcrops of the the eroded volcanic sequence in the eastern part of the Velence Mts. (Fig. 1.). Hydrothermal centres are characterised by irregular and mushroom-shaped massive and vuggy silica bodies surrounded by alunite-kaolinite and pyrophyllite-dickite/kaolinite-diaspore-alunite alteration (Darida-Tichy, 1987; Bajnóczi, 2004). Occurrences of topaz, zunyite and native sulphur are also known in the alteration zones (Nemecz, 1979). Silica bodies are brecciated and contain abundant secondary iron minerals. Non-weathered silica masses are characterised by abundant pyrite dissemination. Alunite crystals from the alteration zones are sodium-rich and contain phosphorus minerals (schwanbergite-woodhausite) in their cores (Vendl, 1931; Bajnóczi et al., 2002). Siliceous breccia bodies with copper mineralisation also occur along an E-W striking contact zone of granite and shale in the eastern part of the Velence Mts., close to the Tertiary volcanic field (Fig. 1.). The breccias occur as several metres wide dike-like bodies on the surface, as well as few centimetre thick veinlets in drillcores. Typical sulphides in the siliceous matrix of the breccia are pyrite, enargite, chalcopyrite and grey ore minerals. Breccia fragments consist of silicified granite and shale with variable size up to 10-20 centimetres and angular shapes. The occurrence of enargite, that is a typomorphic mineral in high sulphidation type epithermal systems (Hedenquist & Arribas, 1999) and the proximity of the mineralisation to the Tertiary hydrothermal fields (Fig. 1.), as well as fluid inclusion data (Molnár, 2004) suggest that the hydrothermal breccias along the granite-shale contact in the eastern part of the Velence Mts. were also formed in the hydrothermal system of the Palaeogene volcanism.

3. Field stops

3.1. Field stop 1: Tertiary hydrothermal alteration zones in the Variscan monzogranite, quarry at the village of Nadap

The quarry at Nadap village is located in the easternmost outcrop of the Variscan monzogranite intrusion, along the Nadap fault separating the Paleozoic granite-shale field form the remnants of the Paleogene volcanic structure to the East (Fig. 1.). The abandoned and partly recultivated quarry on the Gécsi Hill apparently exposes granite only, however, the excavated rock volume was a subvolcanic stock of a Paleogene andesite (Fig. 3.). The quarry is kept in good and safe conditions due to geological interest.

Fig. 3. Geological sketch map of the Gécsi Hill (Gécsi hegy), the area of Field stop 1.

The soutwestern wall of the quarry exposes hydrothermally unaltered, but weathered monzogranite. It is well seen here that weathering acted laterally and caused some vertical zoning of weak smectite-Fe-oxide alteration. However, the southern corner of the quarry shows intense silicification and advanced argillic alteration. The silicified zone also contains pyrite dissemination and several generations of quartz veining. In the less silicified zones the original magmatic texture is well preserved, but K-feldspars and biotite are altered to sericite/illite. The quartz veining is also present along the southeastern wall of the quarry, but the intensity of silicification is lower, whereas the argillic alteration is more intense at some places. This zone corresponds to a NE-SW oriented hydrothermal alteration zone which can also be followed to the southeast outside of the quarry.

Theoretically, the silicification and argillic alteration in the NE-SW oriented zone could have been formed during the granite-related post-magmatic hydrothermal processes, especially, because the orientation of the altered zone is parallel with the most common strike-direction of granite porphyry dikes hosted by the monzogranite. However, K-Ar dating of well crystalline illite provided 30.0 and 32.7 Ma ages which are exactly in the range of ages for alunite and illite from the hydrothermal alteration zones hosted by the Palaeogene andesite in the eastern part of the Velence Mts. (Fig. 4.). The origin of hydrothermal fluids in the argillic alteration zone hosted by the Variscan granite is further constrained by the results of fluid inclusion data. Regional scale fluid inclusion studies in the Velence Mts. revealed significant differences among fluids related to the granite intrusion, to the Triassic hydrothermal processes and to the Palaeogene volcanic activity (Molnár, 1997; Molnár, 2004; Benkó et al. 2014). Fluid inclusions entrapped in hydrothermal quartz and in secondary fluid inclusion planes of rock forming quartz of granite have significantly different homogenisation temperature and salinity data from the granite related pegmatite and molibdenite mineralisation, from the Triassic base metal and fluorite rich veins and from the Tertiary volcanic environments (Fig. 5.). The different entrapment conditions are also reflected by the compositions of fluid inclusion assemblages (Fig. 6.). In the NE-SW oriented alteration zone in the quarry the density (e.g. number*length/mm2) of secondary fluid inclusion planes (FIP) in rock forming quartz is exceptionally high (Fig. 7.). The reason of the increase of the density of fluid inclusion planes is the presence and abundance of those fluid inclusion assemblages which characterise the hydrothermal system of the Paleaogene volcanism (Fig. 6.).Thus, the fluids of the shallow porphyry-epithermal system penetrated and interacted with the Paleozoic monzogranite causing intense alteration.

Fig. 4. Mineralogy and K-Ar ages for the argillic alteration zones in the easternmost part of the Variscan monzogranite intrusion and in the Palaeogene volcanic structure in the Velence Mts.

The presence of the siliceaous-argillic alteration zone of Paleogene age in the old Variscan monzogranite calls on the importance of evaluation not only the spatial connection between hydrothermal alteration zones and magmatic formations but also their temporal relationships.

Fig. 5. Fluid inclusion homogenisation and salinity data for the Variscan, Triassic and Palaeogene hydrothermal systems in the Velence Mts.

Fig. 6. Summary of chracteristics of fluid inclusion assemblages in the Variscan, Triassic and Palaeogene hydrothermal systems of the Velence Mts.

Fig. 7. Zoning in the densities of fluid inclusion planes along the argillic alteration zone of Palaeogene age hosted by the Variscan monzogranite on the Gécsi Hill (Field stop 1). The high densities of fluid inclusion planes mark the zones of the most intense interaction of the Palaeogene hydrothermal system with the Variscan granite..

3.2. Field stop 2: Hydrothermal alteration in the Palaleogene intrusive-volcanic complex of the Velence Mts., Zsidó Hill, near the village of Pázmánd

The small quarry on the Zsidó Hill near the village of Pázmánd exposes strongly silicified and argillitised hydrothermal alteration zones in the easternmost outcrops of the Palaeogene andesitic volcanic structure (Fig. 1.). The argillic zones here consist of pyrophyllite (Nemecz, 1979). All of the outcrops of the eroded remnants of the stratovolcnic structure in the eastern part of the Velence Mts. are characterised by intense and zonal hydrothermal alteration with massive silica bodies in the centre and advances argillic alteration mantle around them. However, most of the massive silica bodies are not the result of silicification (e.g. adding silica to the rock), but represent residual silica bodies (e.g. apart of SiO2, almost all of alkali, alkali earth and other metal oxides were leached out from the rock). The residual silica bodies are characterised by vuggy texture because they contain empty places of dissolved rock forming minerals. Therefore they are called “vuggy silica” bodies. Formation of residual silica bodies requires extremely acidic

(pH<2) conditions which is characteristic to hydrothermal systems with mixing of volcanic SO2 and meteoric water. Thus the residual silica bodies also marks the major upflowing zones of magmatic fluids. The advanced argillic alteration zones around the residual silica bodies of the eroded stratovolcano of the Velence Mts. contain variable mineral assemblages (e.g. pyrophillite- dickite; topaz-zunyite-kaolinite, kaolinite only etc. with or without alunite; Fig. 4.). This variability reflect different temperature and acidity conditions (Fig. 8.) suggesting that these alteration

15 zones at the approximately same elevations at their current positions were formed at different depths at the time of the hydrothermal activity.

Fig. 8. Variation in the clay mineralogy of the hydrothermal alteration zones as a function of temperature and fluid composition in the Palaeogene hydrothermal system of the Velence Mts.

Near to the entrance of the quarry at the Zsido Hill, the 1200 m deep Pázmánd (Pd) – 2 hole was drilled during a mineral exploration project in the 1970`s. This drillhole found a porphyric diorite intrusion at 700 m depth below the current surface, with showings of porphyry-copper type mineralisation under the altered volcanic rocks (Fig. 9.). Vertical zoning of hydrothermal alteration in this drillhole is characterised by advanced argillic alteration at shallow levels, phyllic alteration (sericite/illite) at intermediate depths, and K-feldspar alteration in the intrusion at depth. This vertical zoning is typical in porphyry-epithermal systems (Sillitoe, 2010). The vertical and horizontal zoning of hydrothermal alteration in the Palaeogene intrusive- volcanic complex of the Velence Mts. is the result of the interaction of mixed magmatic- meteoric fluids with igneous rocks. The degree of mixing between the SO2 rich high temperature acidic magmatic and cold-neutral meteoric components of hydrothermal fluids varied as a function of distance of major upflow zones. This mixing process together with the fluid/rock ratio is well expressed by the variation in the O and H isotope composition of hydrous silicate alteration minerals (Fig. 10.). The very ligh H isotope composition of rock forming hydrosilicate minerals in the less altered zones of the diorite intrusion reflects fractionation of isotopes during segregation of magmatic fluids from the parent melt.

Fig. 9. Vertical zoning of hydrothermal alteration in the Pd-2 drillhole. The vertical aleration zoning found in this drillhole is typical for porphyry-epithermal systems.

Fig. 10. O and H isotope data for hydrothermal and rock forming hydrous silicate minerals from the Palaeogene intrusive-volcanic complex of the Velence Mts.

4. References

Bajnóczi, B (2004): A Velencei-hegység paleogén hidrotermális folyamatai (Palaeogene hydrothermal processes of the Velence Mts.). Unpublished PhD dissertation, Budapest (in Hungarian) Bajnóczi, B, Molnár, F, Maeda, K, Nagy, G, Vennemann, T (2002): Mineralogy and genesis of primary alunites from epithermal systems of Hungary. Acta Geol. Hung., 45: 101–118 Benedek, K (2002): Palaeogene igneous activity along the eastern-most segment of the Periadriatic-Balaton Lineament. Acta Geol. Hung., 45/4: 359-371 Benkó, Zs, Molnár, F, Lespinasse, M, Billström, K, Pécskay, Z, Németh, T (2014): Triassic fluid mobilization and epigenetic lead-zinc sulphide mineralization in the Transdanubian Shear Zone (Pannonian Basin, Hungary). Geol. Carpath., 65/3: 177-194 Böjtös-Varrók, K (1966): Wolfrám a Velencei-hegységben. Bull. Hung. Geol. Inst. 1964., 293–300 (in Hungarian) Böjtös-Varrók, K (1967): A palaköpeny hidrotermás ércesedése a Velencei hegységben. Bull. Hung. Geol. Inst. 1965, 293–299 (in Hungarian) Buda, Gy (1985): Origin of collision-type Variscan granitoids inHungary, West Carpathian and Central Bohemian Pluton. Unpublished PhD thesis, Budapest, 148 p. Buda, Gy (1993): Enclaves and fayalite-bearing pegmatitic "nests" in the upper part of the granite intrusion of the Velence Mts., Hungary. Geol. Carpath., 44/3: 143-153 Csontos, L (1995): Tertiary tectonic evolution of the Intra-Carpathianarea: a review. Acta Vulc., 7: 1-13 Darida-Tichy, M (1987): Paleogene andesite volcanism and associated rock alteration (Velence Mountains, Hungary). Geologicky Zbornik – Geol. Carpath., 38: 19-34 Dudko, A, Darida-Tichy, M, Majkuth, T, Stomfai, R, (1989): A Kelet-Velencei paleovulkán szerkezete (The structure of the E-Velence palaeovolcano). Ált. Földt. Szemle, 24: 135-148 (in Hungarian) Erdélyi, J (1939): A nadapi barit és hematit (Barite and hematite from Nadap). Földt. Közl., 69: 290–296 (in Hungarian) Erdélyi, J (1952): Baryt von Sukoró. Acta Geol. Hung., 1: 2–9 Hedenquist, JW & Arribas, A (1999): Epithermal gold deposits: I. Hydrothermal activity and processes, and II. Characteristics and genesis of epithermal deposits. In: Molnár, F, Lexa, J, Hedenquist, JW (eds.): Epithermal mineralization of the Western Carpathians. Soc. Econ. Geol. Guidebook Ser., 31: 13-64 Horváth, I & Ódor, L (1984): Alkaline ultrabasic rocks and associated silicocarbonatites in the NE part of the Transdanubian Mts, Hungary. Mineralia Slovaca, 16: 115–119 Jantsky, B (1957): A Velencei-hegység földtana (The geology of Velence Mts.). Geol. Hung. Ser. Geol., 10, 166 p. Kázmér, M & Kovács, S (1985): Permian-Palaeogene paleogeographyalong the eastern part of the Insubric-Periadriatic Lineament system: evidence for continental escape of the Bakony- Drauzug Unit. Acta Geol. Hung., 28/1-2: 617-648 Kiss, J (1954): Hydrothermal mineralization on the northern part of the Velence Mts. Bull. Hung. Geol. Inst. 1953, 111-127 Koch, S. (1985): Magyarország Ásványai (Minerals of Hungary). 2nd edition (Ed.: Mezősi, J) Akadémiai Kiadó, Budapest. 562 p. (in Hungarian) Molnár, F (2004): Characteristics of Variscan and Palaeogene fluid mobilization and ore forming processes int he Velence Mts., Hungary: A comperative fluid inclusin study. Acta Mineral. – Petrogr., Szeged, 45/1: 55–63

Molnár, F (1997): Újabb adatok a Velencei-hegység molibdenitjének genetikájához: ásványtani és folyadékzárvány vizsgálatok a Retezi-lejtakna ércesedésén (New data regarding the genesis of molybdenite from Velence Mts: mineralogical and fluid inclusion studies on ores from the Retezi-lejtakna). Földt. Közl., 127/1-2: 1–17 (in Hungarian) Nagy, B (1967): A sukorói turmalinos pegmatit előfordulás ásvány-kőzettani, geokémiai vizsgálata (Mineralogical, petrological and geochemical study of the tourmaline-bearing pegmatite occurrence of Sukoró). Bull. Hung. Geol. Inst. 1965, 507–515 (in Hungarian) Nemecz, E (1979): A pázmándi Pd.-1.sz. pirofillit kutatófúrás anyagvizsgálata (Study of the Pd-1 pyrophyllite borehole of Pázmánd). Manuscript. Data repository of the Hungarian Geological Institute (in Hungarian) Pamić, J (2002): The Sava-Vardar Zone of the Dinarides and Hellenides versus the Vardar Ocean. Eclogae Geologicae Helvetiae, 95: 99-114 Salters, JM, Hart, SR, Pantó, Gy (1988): Origin of Late Cenozoic volcanic rocks of the Carpathian Arc, Hungary. In: Royden, LH & Horváth, F (eds.): The Pannonian Basin, a study in basin evolution. Am. Assoc. Petroleum Geol., Geol. Memoirs, 45: 279-292 Sillitoe, RH (2010): Porphyry copper systems. Econ. Geol., 105: 3–41 Uher, P, Putis, M, Kohut, M, (2010): Permian A-type granites of the Western Carpathians: results of zircon SHRIMP dating, DATING 2010 - Slovak Geol. Soc. Vendl, A (1923): Újabb adatok a Velencei-hegység kőzeteinek ismeretéhez (New data about the rocks of Velence Mts). Ann. Mus. Nat. Hung., 20: 81 (in Hungarian) Vendl, A (1931): A nadapi alunit (Alunite from Nadap). Math. Term. Tud. Ért., 31: 95–101 (in Hungarian)

Notes:

Day 1 stops 3-7: Hydrogeology and thermal water related phenomena in the Gellért Hill area – Introduction to the field trip in the Buda Thermal Karst area

Anita Erőss, Petra Kovács-Bodor, Ádám Tóth, Tímea Havril, Andrea Mindszenty, Judit Mádl-Szőnyi (Eötvös Loránd University, Department of Physical and Applied Geology)

1. The hydrogeological environment of the Buda Thermal Karst area

1.1 Geological settings

Geologically, the Buda Thermal Karst system (BTK) is the NE extreme of the Transdanubian Range (TR) (Fig. 11.). The hydrogeologically complex thick carbonate system of the TR is located in the Central Pannonian Basin, Hungary, and is bound by major tectonic lineaments. The Pannonian Basin is a Neogene structure formed as a result of extension-related attenuation of the lithosphere in late Early to Late Miocene times (Royden & Horváth, 1988). From the latest Miocene on, inversion of the basin has been continuous and contributed to the uplift of certain basement-blocks. The TR is one of these uplifted blocks. Due to the thin lithosphere, the whole Basin is characterised, even now, by elevated heat flux (~100 mW/m2) as compared to the surroundings (Lenkey et al., 2002). Geological boundaries of the TR are tectonic lineaments, namely: Rába Line in the NW, Diósjenő-Ógyalla Line in the N and Balaton Line in the S.

Fig. 11. Transdanubian Range Unit: boundaries of the range and the Paleozoic and Mesozoic formations with the indication of main structural features and the area of unconfined carbonates (modified after Fülöp in Haas, 2001 and Fodor, 2010) and also the morphologic ally delineated catchment area of Buda Thermal Karst (BTK) (after Mádl-Szőnyi et al., 1999) (after Mádl-Szőnyi et al., 2017a)

The TR is built up by thick Paleozoic metamorphic and Permian to Cretaceous sedimentary sequences; however, the Triassic–Early Jurassic carbonates form the main karst system of the area. They are dissected by major faults and deformed also by folding. The carbonate range is of 300 km x 100 km in extent (Csepregi, 2007) and is partially confined by thick Neogene sediments mainly at the edges (Fig. 11.). In the area of Budapest, the Mesozoic carbonate suite of the TR is downfaulted and continues deep in the basin to the East of the Danube, under a thick Neogene cover. The basement of the BTK is not known, but like in the whole TR, Permian siliciclastics and shallow marine evaporitic carbonates are hypothesised also here (Majoros, 1980; Haas et al.; 1989). The schematic stratigraphic column of the Buda Hills is shown by Fig. 12. The oldest known strata in the BTK area are Triassic (Ladinian to Norian) platform carbonates (dolomites and

20 limestones) that constitute the bulk of the area with a thickness of 1500 to 1800 m (Wein, 1977; Haas, 1988). During this time the TR was part of the southern shelf of Neotethys located between the South Alpine and East Alpine Units (Haas et al., 1995). Jurassic and Cretaceous sediments are missing in the BTK area, however, since from the Early Jurassic to Mid-Cretaceous sedimentation had been continuous all over the TR, it is supposed that these strata were present also in the BTK area, but were removed by subsequent erosion. Triassic carbonates are separated from the overlying Eocene strata by a long-lasting period of tectonically controlled subaerial exposure that spanned probably from Late Cretaceous to Late Eocene (Wein, 1977). The Palaeogene transgressional sequence starts with basal conglomerate and breccia, with erosional remnants of a once continuous bauxitic blanket at its base. This is overlain by Upper Eocene to Lower Oligocene limestone, marls and clays. From the Mid-Eocene to Mid-Miocene times two different depositional environments have developed in the Buda Hills. The dividing line between the two, the Buda Line, was an important paleogeographic boundary (Báldi & Nagymarosy, 1976), interpreted as a paleotopographic high formed above a blind thrust according to Fodor et al. (1994). SE of this antiform, sedimentation continued in the pelagic marly-clayey facies from Late Eocene well into the Oligocene with extensive deposits of the Buda Marl, the anoxic Tard Clay and several hundreds of metres of the Kiscell Clay (Báldi et al., 1976; Báldi, 1983). The area NW of the Buda Line became subaerially exposed in latest Eocene earliest Oligocene times. Sedimentation over the eroded surface of Mesozoic and older Tertiary formations resumed still within the Early Oligocene with coarse-grained coastal siliciclastics (Hárshegy Sandstone Formation). Fine-grained sand and silt deposition took place during the Late Oligocene. The area was gradually uplifted along NW-SE oriented normal faults, and became subaerially exposed again by the end of the Oligocene and Early Miocene. From the Miocene on the sedimentation in the area became associated with the evolution of the Pannonian Basin. Along the margins of the “island” of the Buda Hills coastal and shallow marine siliciclastics and also carbonate sediments were deposited, while the uplifting interior became subject to erosion. Since the basal strata of the early Neogene shallow marine sediments around the Buda Hills do not contain any clasts from either the Triassic or the Eocene carbonates, it is assumed that at this time, these carbonates were still covered by thick Oligocene clays (Wein, 1977). From the latest Miocene on deposition of lacustrine limestones and travertines began and continued with changing intensity up to recent times. They were tentatively associated with thermal waters discharging along the foothills of the uplifting terrains as already suggested by Schréter (1912). This was proved by detailed geomorphological, sedimentological, biostratigraphical and geochemical studies by Scheuer & Schweitzer (1988), Müller & Magyar (2008), Korpás (2003), Kele (2009) and by many others. According to recent results of Kele (2009), travertine formation was most intense in the Middle Pleistocene, suggesting that deep, regional-scale water circulation was already active at that time. However, his latest U/Th datings show that from Late Pleistocene to Holocene times there was no appreciable travertine deposition in the Buda Hills. In order to improve the understanding of the Pleistocene uplift history of the area Szanyi et al. (2012) tried to calculate the uplift rate by speleotheme dating in several former thermal caves. Based on U/Th dated cave rafts they suggest that from 500 to 280 ka BP the area was characterised by rather slow uplift (0.06-0.3 mm/y), however, in the 280-70 ka interval uplift (and/or concomitant river incision) had been accelerated to 0.16 mm/y on the average. Latest results of Virág (2019) have shown that this accelerated rate of incision in several cave of the BTK was accompanied by substantial (several metres) increase of the karst watertable and also the temperature of the thermal water pointing to some climatic event (a pluvial?) interacting with the supposedly monotonous, tectonically controlled uplift. Focussing on karst features it has to be pointed out, that all the Late Triassic to Early Tertiary formations

21 of the Buda Hills are carbonates, highly prone to karstification (Fig. 12.). In addition to the limestones and dolomites, the calcareous Buda Marl with its rather high carbonate content (50- 80 %; Kleb, 1993) also belongs to the karstic category that is evidenced by the presence of cave passages in it (Leél-Őssy, 1995; Leél-Őssy et al., 2011).

Fig. 12. Stratigraphic column of the Buda Hills (after Poros, 2011)

There were four major tectonic phases to affect the carbonate succession of the study area (Fodor et al., 1994). During the Cretaceous, compressional tectonics was predominant, due to the SE-NW compression, anticlines were formed (with folds and smaller thrusts) and the rigid dolomites became heavily fractured (Wein, 1977). From the Late Eocene to the Early Miocene the area was intermittently affected by large scale strike-slip activity (WNW-ESE directed compression and perpendicular tension). The stress field changed in the Middle Miocene, and until the Quaternary new N-S to NE-SW trending normal faults were formed by ESE-WNW

22 extension. Pleistocene tectonics is characterised by N-S (NE-SW) extension and perpendicular compression. Due to the complex tectonic evolution of the area, the generated fractures, faults and folds play crucial role in canalizing groundwater flow, and controlling dissolution and precipitation processes.

1.2. Flow regimes

The lithological continuity of the Triassic carbonate suite provides for the hydraulic continuity of the large groundwater body of the TR. The hydraulic continuity of the TR was inferred on the basis of the results of long-term (1950–1990) mine dewatering (Alföldi & Kapolyi, 2007) which has resulted in drying up of springs at the highest elevations and caused changes in chemical composition, temperature and spring discharge at lower elevations. In addition, the hydraulic head decrease could be followed on a regional scale for the whole reservoir (Csepregi, 2007). Precise delineation of the BTK system in the hydraulic sense is, therefore difficult within this more than 300 km long and about 100 km wide body of the TR. Because the aquifer system of the TR is characterised mainly by gravity-driven groundwater flow (Alföldi, 1979), the topographic elevation was taken into account at the first delineation of the surface catchment area of BTK by Mádl-Szőnyi et al. in Mindszenty (1999) (Fig.11.). This delineated surface catchment area includes areas also to the East of the Danube, where the exposed Mesozoic carbonate hills and topographically elevated non-carbonate areas (Gödöllő Hills) contribute to the discharge of the Buda Thermal Karst (Mádl-Szőnyi et al., 2015; Mádl- Szőnyi et al., 2017b). The bare karstic surfaces represent 15% of the whole TR (Lorberer, 1986). Based on recent hydrogeological studies the new conceptual model of the BTK system is illustrated on Fig. 13.

Fig. 13. Conceptual Tóthian-type GDRGF pattern for the boundary of unconfined and confined sub-basins. The consequences on flow-related manifestations imply that the shallow karst aquifer (inset; modified after Goldscheider & Drew, 2007) is embedded into the regional flow pattern as a local system (after Mádl-Szőnyi & Tóth, 2015) 23

Fig. 14. a Surface manifestation of GDRGF for the TR. Springs are related to local (cold), intermediate (lukewarm) and regional (thermal) flow systems based on the classification of Mádl-Szőnyi & Tóth (2015). Location of epigene and hypogene caves is based on the data from the Hungarian Ministry of Agriculture and Rural Developments, Hungary (A. Gazda, 2014, personal communication). The figure is complemented by the morphologically delineated catchment area of Buda Thermal Karst (BTK) (after Mádl-Szőnyi et al. in Mindszenty, 1999). b Spring groups and epigene and hypogene caves of Budapest and surroundings (EOV is the Hungarian National Grid which is a transverse Mercator projection— positive X is pointed to north and positive Y is pointed to east. Coordinates refer to metres)

The BTK area receives fluid from several sources (meteoric and basinal), in agreement with the conceptual model (Fig. 13.). In addition to meteoric fluids, the contribution of basinal fluids at the discharge area of the BTK was first suspected based on water chemical analyses by Alföldi (1979) and was proved based on recent analyses by Erőss (2010), Erőss et al. (2012a) and on mineralogical and fluid inclusion studies by Poros et al. (2012). The origin of the basinal fluid was identified as being related to the confining layers of the carbonates. The mechanisms of its contribution to spring discharge were also revealed in the form of predominant vertical downward leakage from the confining layer to carbonates and upward flow to the springs (Mádl-Szőnyi et al., 2015; Mádl-Szőnyi & Tóth, 2015; Mádl-Szőnyi et al., 2017b). The discharging fluids from different origins (meteoric and basinal) with different flow systems (intermediate and regional) result in a wide range of discharge features including springs, caves and also mineral precipitates. Extensive hypogene cave systems have developed (e.g., Takács-Bolner & Kraus, 1989; Leél-Őssy, 1995; Leél-Őssy & Surányi, 2003; Leél-Őssy, 2017) and are actively forming even at present due to the interaction between fluids and the carbonate rocks (Erőss, 2010; Erőss et al., 2011). Therefore, the BTK system can be considered as a type example of an active hypogene karst (Erőss, 2010). The natural discharge of the system is manifested mainly in the form of springs along and in the riverbed of the River Danube forming three distinct discharge areas and are strongly influenced by the tectonic pattern (Alföldi et al., 1968; Ötvös et al., 2013; Erhardt et al., 2017) (Fig. 14.). The Northern discharge zone of the BTK is characterised by lukewarm springs (18–24 °C, TDS < 1000 mg/l); in the Central zone both lukewarm (21–27 °C, TDS < 1000 mg/l) and thermal springs (53–63 °C, TDS > 1000 mg/l) occur, while in the Southern discharge area of the system temporally and spatially uniform thermal water discharge (33– 45 °C, TDS 1500–1700 mg/l) is characteristic (Papp, 1942; Alföldi et al., 1968; Erőss et al., 2008).

The lukewarm springs were evaluated as belonging to intermediate flow systems and the thermal springs as the discharge of regional flow systems based on cluster analysis (Bodor et al., 2014) and numerical simulations (Mádl-Szőnyi & Tóth, 2015). Since the second part of the nineteenth century, deep wells have been increasingly used in addition to natural springs. Déri- Takács et al. (2015) and Kovács & Erőss (2017) evaluated the waters in Budapest by multivariate exploratory techniques. Based on geochemical parametres and temperature data, seven groups as optimal grouping were found by combined cluster and discriminant analysis (CCDA) (Kovács et al., 2014). The seven groups found by CCDA are in accordance with the previously established hydrogeological conditions and flow system characteristics of the area. All the groups represent different parts of the gravitationally driven flow system of the area. Déri-Takács et al. (2015) showed that the influence of the temperature and chloride content was the strongest in grouping springs according to their characteristics. This is in good agreement with the contribution of basinal fluids proved in the discharging water. A comprehensive hydrogeological study was carried out for the characterisation of processes acting presently and their resulting parametres at the discharge zone of the BTK (Erőss, 2010; Erőss et al., 2011, 2012a, b). Studying the attributes of springs and wells, caves and mineral precipitates, that is, the entire range of phenomena related to discharge (Tóth, 1971), information can be obtained related to the parent flow system, its hydrogeological environment and the processes taking place throughout the entire length of the flow system and in the close vicinity of the discharge zone. Moreover, the identification and understanding of recently active processes and their manifestations in this hypogene karst area will help to identify and understand paleo-phenomena both in the BTK and in other hypogene karst areas of similar settings. The relationships between flow systems and hypogene karstification were examined for the Central and Southern discharge zone, with further input from the deep wells for the characterisation of the confined part of the system in

Pest. Two distinct conceptual flow- and cave development models were elaborated for the Central and Southern discharge areas (Mádl-Szőnyi et al., 2017a).

The evolution of the BTK started in the Late Miocene with the structural inversion of the Pannonian Basin, which led to the uplift of the western “Buda” part and simultaneous subsidence of the eastern “Pest” part of the system. Due to the uplift of the Buda Hills (Kele et al., 2011; Ruszkiczay-Rüdiger et al., 2005; Wein, 1977), the erosion of Palaeogene siliciclastic cover formations was initiated from Pliocene times. Simultaneously, due to the NW-SE trending normal faulting between the uplifting and subsiding blocks (Fodor et al., 1999), sedimentation continued in the subsiding basinal sectors. The structural inversion of the area east of the BTK (the so called Gödöllő Hills) has been ongoing for the past 4 million years (Ruszkiczay-Rüdiger et al., 2006; Ruszkiczay-Rüdiger et al., 2007). Due to this landscape evolution the recent surface watershed of the BTK has been evolving (Ruszkiczay-Rüdiger et al., 2007). During the large scale changes caused by the uplift and erosion, relative importance of fluid driving forces was changed, which led to the alteration in the dominance of the system fluid components. In the fully confined initial system state, the flow and temperature field was clearly dominated by natural thermal convection (Havril et al., 2016). In agreement with the regional distribution of vein-filling minerals and connected paragenesis described by Poros (2011), a uniform groundwater flow system can be assumed within the BTK until the start of uplift in the Late Miocene. Development of the current porosity of the BTK was assumed to start in Miocene times (Poros, 2011). Porosity and permeability enhancement in that buoyancy-dominated stage may be linked to the increasing dissolution capacity of cooler water raised by convection cells (Andre & Rajaram, 2005; Dublyansky, 2000; Mádl-Szőnyi et al., 2018). Indeed, retrograde solubility of calcite is known to sustain dissolution along a flow path with decreasing temperature, as discussed e.g. by Palmer (1991). Initially, the pore space was assumed to be filled with antecedent seawater (Mádl-Szőnyi et al., 2018). As the area became subaerially exposed, and uplift of the western part of the system began, meteoric water started to infiltrate and dilute the original pore water, and simultaneously modified the redox conditions within the system, in agreement with vein-filling paragenesis which reflects a more oxidative environment during system evolution (Poros et al., 2012). After 4 Ma, due to the initiated uplift of the eastern sub-system (Ruszkiczay-Rüdiger et al., 2007), basinal fluid flow toward the main discharge point began, which could have contributed to the formation of hypogene caves (Poros et al., 2012; Dublyansky et al., 2014). Such cave formation would not only be due to the mixing of different fluids (Gray & Engel, 2013; Mádl-Szőnyi & Tóth, 2015) but also due to the dissolution-enhancing effect of aggressive gases (e.g. CO2 and H2S) (Palmer, 2007) transported by groundwater from the covered sub-basin (Poros et al., 2012; Erőss et al., 2012a). With progressing uplift and erosion the area of exposed carbonates increased and provided increasing direct recharge to the system. This resulted in increasing proportion of karst water at the expense of basinal fluids. Uplift resulted in growing hydraulic potential, i.e. in continuously strengthened gravitational driving force. As the recharge area of the gravitationally driven (karstic) flow system increased, the karstic component became more and more dominant in the system. This is reflected by the increasing occurrence of carbonate mineral precipitates. However, the basinal fluid component is still present. Today, we can see a “snap-shot” of this transient system.

These processes were also evaluated by fluid flow, heat and mass transport simulations and the change of the rate of different processes from the late Miocene to recent time were revealed (Havril et al., 2016; Havril, 2018; Szijártó et al., 2019; Galsa et al., 2019). For instance, it was established that the pure advective heat transport due to steady-state forced thermal convection causes a large hot upwelling with a surface temperature of 50−70 °C beneath the

26 regional discharge area (i.e. beneath the River Danube) in agreement with appearances of thermal springs. However, additional effect of time-dependent natural convection caused by buoyancy forces also facilitates hot upwellings in Mesozoic carbonates, which might elucidate the unexplained heat anomalies in temperature maps and profiles (Mádl-Szőnyi, 2019). Direct volcanic contribution to the hydrothermal fluids could not be confirmed. Instead, upwellings were suggested as the heat source of the thermal fluids due to the elevated heat flux of the Pannonian Basin as well as the insulating role of the cover siliciclastic formations (Mádl-Szőnyi et al. 2018; Havril, 2018).

2. Field stops in the Gellért Hill area

2.1. Field stop 3: Panoramic view from Gellért Hill

2.1.1. Geological build-up of the Gellért Hill

The Gellért Hill (235 m asl) is situated right on the riverside (Fig. 15.). It is an almost bare karst area, bound by major fault lines and is characterised by steep slopes – almost vertical cliffs – on the Danube-side (Fig. 16.).

Fig. 15. Location map of field stops in the Gellért Hill area: Stop 3 Panoramic view from Gellért Hill; Stop 4 Gellért Spa building; Stop 5 Ősforrás (Spring of the Gellért Spa); Stop 6 Gellért Tunnel; Stop 7 Aragonite Cave (modified after Kovács-Bodor et al., 2019a)

The bulk of it is built up by Upper Triassic dolomite (Fig. 16.) containing also organic-rich marl and clay-marl intercalations. It is unconformably overlain by a Late Eocene transgression sequence beginning with cherty basal conglomerates, breccia or sandstone. The marly matrix of the basal clastics increases upwards and very soon it gives way to the Buda Marl, with some sandstone and nummulitic limestone intercalations. Upwards the Buda Marl is passing into pelagic (shallow bathyal) clays (Oligocene Tard- and Kiscell Clay) and then by sandstones (Upper Oligocene Törökbálint Sandstone). Neither the Miocene nor the Pliocene is represented in the stratigraphic column of the Hill. In the Quaternary travertines, loess, slope scree, fluvial gravel and sand were deposited (Korpás et al., 2002). Fig. 17. shows the schematic stratigraphic column of the Gellért Hill area. The geological map on Fig. 16. and the section on Fig. 18 clearly show that the area was also strongly influenced by tectonics. According to Fodor in Korpás et al. (2002) three different Tertiary stress fields could be identified, and one Triassic event is also supposed. The Tertiary stress fields can be assigned to four deformational events. The oldest, NNE-SSW

27 tension created the E-W trending “Citadella fault”. Due to its dextral to normal synsedimentary displacement, the SE part of the hill subsided ca. 100 m. The fault deformed then the only partly consolidated sediments and induced various soft-sediment deformation features (creeps, slides, etc.) within the Upper Eocene-lowermost Oligocene clastics.

Fig. 16. Location of the springs at the Gellért Hill discharge zone (geological map is based on Fodor in Korpás et al., 2002)

Fig. 17. Simplified lithostratigraphic column of the Gellért Hill area (modified after Korpás et al., 2002) (Tr: Triassic, E: Eocene, O: Oligocene, Q: Quaternary; 1: Main Dolomite, 2: Basal breccia and conglomerate, 3: Szépvölgy Limestone, 4: Buda Marl, 5: Tard Clay, 6: Kiscell Clay, 7: Törökbálint Sandstone, 8: Quaternary sediments: gravel, sand, travertine)

Fig. 18. Geological profile, across the river right to the South of the uplifted Gellért Hill (by L. Fodor in Korpás et al., 2002) (For trace of the profile refer to Fig. 16. A-B)

Faulting resulted in steeply dipping sedimentary dykes with sandy marl and limestone fillings. Most of the dykes and the marl itself were silicified presumably due to hydrothermal fluid circulation along the main fault. After the Early Oligocene, the western segment of the “Citadella fault” was reactivated and connected to another dextral normal fault south of the Gellért Hill. The Ottnangian-Middle Miocene phase induced a large normal slip on the north-eastern boundary fault of the Gellért Hill. The fault crosses the Danube petering out to several fault strands. This fault represents the eastern boundary fault of the whole Buda Hills and might have accumulated up to 1000 m of displacement. Reactivation took place during the Late Miocene to Quaternary due to ESE-WNW tension.

2.1.2. Field stop 4: Gellért Spa building, overview of the flow regime

Since the current hydrogeological system is characterised by an artificially influenced groundwater discharge the primary discharge features – for instance location of the upwelling; temperature and chemical composition of the springs – were summarised based on the evaluation of archive hydrogeological data by Erőss et al. (2008). In the Gellért Hill discharge zone the springs could be characterised by temporally and spatially uniform temperature (33.5- 43.5°C) and chemical composition (1450-1700 mg/l TDS). This study revealed strong structural control on the springs’ locations as well (Fig. 16.). The discharge rate of the springs has varied between 2390-4022 m3/day (Papp, 1942). Geochemical analyses confirmed the similarity of all thermal waters discharging within the Gellért Hill area. They are characterised by elevated Ca2+, 2+ - 2- Mg , HCO3 , SO4 and TDS content. These higher values correspond to lower temperatures within a narrower range (35–47 °C), compared to the central discharge area. With the application of radionuclides (Erőss et al., 2012b) it was revealed that in this area the discharge of thermal waters is overwhelming, no mixing components could be identified. The established differentiated conceptual model for the Gellért Hill area is illustrated on Fig. 19. The hydraulic cross section on Fig 20 based on measured archive well data explains the flow conditions in the Gellért Hill area.

Fig. 19. Conceptual flow model for the Gellért Hill discharge area (modified after Erőss, 2010; in Mádl-Szőnyi et al., 2017a)

Fig. 20. Hydraulic cross section across the Gellért Hill (hydrostratigraphy adapted from Fodor, 2013) (Erhardt et al., 2017)

The north-western Cenozoic strike-slip fault (F#1) could be presumed as a hydraulically barrier fault, which induces significant difference in hydraulic head values on its two sides (Fig. 20.). This is supported by the westward thickening mediocre aquifer and aquitard units (in the cover of the carbonates) west of the fault, which are juxtaposed to the carbonate aquifer of the eastern side of the fault. This presumably barrier fault and the appearance and thickening of the moderate aquifers on its eastern side have a significant effect on the transversal fluid flows of

30 the area, because the recharging water in the mountain range—located in the west of the Gellért Hill—cannot or can just moderately discharge at the Gellért Hill discharge area. However, a positive hydraulic anomaly occurs around the junction of a Cretaceous fault and a Cenozoic strike-slip fault (Fig. 20., F#2 and F#3 respectively; 3,500–3,700 m section) at shallow depth. This anomaly denotes the fluid-flow conduit character of these fault zones. As a result, upwelling thermal fluids can discharge at the Gellért Hill discharge area preferably along these structural elements. It is also worth mentioning that although mediocre aquifers can be found between the two faults and in the east of the Cenozoic strike-slip fault—F#3; above z = −(200)m—which could slightly hamper fluid flow, these rocks may have a significant secondary porosity in the highly fractured junction of the two faults. As a consequence, there is no remarkable difference in hydraulic heads between the two sides of the faults.

Fig. 21. Hydraulic cross section across the Gellért Hill (hydrostratigraphy adapted from Fodor, 2013) (after Mádl-Szőnyi & Tóth, 2017)

Steady-state flow and heat transport simulations were also carried out in 2D along a hydrostratigraphic section of BTK through the Gellért Hill (Fig. 21.). The physical properties of the units and the applied boundary conditions of the flow and heat transport simulations are summarised in Fig. 21. The flow pattern of the western dominantly unconfined sub-basin is very similar to the Unit basin-type pattern. The conductive faults enhance recharge at 10–22 km and discharge at the Gellért Hill and 2 km west of the Gellért Hill. The eastern, confined part of the basin represents dominantly through-flow from east toward the Danube with downward flow

31 component in the lower part of the confining siliciclastic units. The asymmetric character of flow can be explained by the different rate of recharge across the unconfined and confined parts of the carbonate system. Under the confined sub-basin the vectors show converging directions toward the discharge area. Most intense discharge appears at the Danube. The temperature field shows cold water around 15 °C in the upper part of the system on West; while the water temperature is higher than 70 °C on top of the carbonates in the confined part of the system under the Pest Plain. Significant heat accumulation can be observed under the confined sub- basin of Pest (32-40 km). A heatplume can be observed toward the discharge area at the Danube (Mádl-Szőnyi & Tóth, 2017).

2.2. Field stop 5: Ősforrás (Spring of the Gellért Spa)

Only one natural spring outlet is known from the Gellért Hill discharge area, which is rather a spring group. The ancestor of the Gellért Spa, the so called „Muddy Spa” was established on this spring (Fig. 22.). The place was named after the accumulation of carbonate precipitates the so called mud. Later on the Gellért Spa (opened in 1918) also used this spring and a capture room was constructed above it, which is a ~ 100 m2 pool with trenches, where 17 spring outlet points were detected. The height of the room is 9.5 m, which is located 3.5 m below street level. The discharge rate of the spring depends on the level of the Danube, on the average it is 20 L/s. The spa used this springwater until 1978 directly, then it was substituted by wells drilled into the aquifer from a tunnel running parallel to the river in the interior of the hill (Papp, 1942; Schulhof, 1957).

Fig. 22. The so called “Muddy Bath”, the ancestor of the Gellért Spa in the 19th century (from the archive of I. Dobos)

The springs in the area are influenced by the transient effect of the Danube (discharge rate, chemical composition, temperature), however was according to an important observation the river water never mixes with the spring waters. It is already known from the earlier studies (e.g. Schafarzik, 1920) that the thermal waters discharge directly into the riverbed during normal stage or low-flow conditions (Fig. 23a). However, during flood conditions the discharge of the underwater (riverbed) springs is shifted to the riverbank zone owing to the elevated hydrostatic pressure of the river. At the Gellért Hill discharge zone and to the North this model describes the interaction of discharging thermal waters and the Danube (Fig. 23b) (Somogyi, 2009).

Fig. 23. The hydraulic connection between the Danube and the groundwater of the surroundings: a) during normal stage or low-flow conditions at the whole discharge zone along the Danube in Budapest; b) during flood at the Gellért Hill and to the North (Somogyi, 2009)

2.3. Field stop 6: Gellért Tunnel: Wells and in situ experimental study of the chemistry of discharging thermal water and related precipitates

The tunnel was built between 1969 and 1978, in order to capture the water of the riverbed springs in the Danube front of the hill. The tunnel with its 1100 m length connects the Gellért, Rudas and Rácz Spas. 18 wells were drilled on both sides, from which 4 are used for water supply of the Gellért and Rudas Spas. Besides chemical calcite precipitates so called biogeochemical precipitates can also be observed in the spring pools (Erőss, 2010; Erőss et al., 2012a; Dobosy et al., 2016). To understand better the evolution of calcite and biogeochemical precipitates in spring pools, in situ experimental studies were carried out in the canal of Gellért Tunnel (Kovács-Bodor et al. 2018 a, b). Thermal water was pumped at a constant rate (with ~ 1*10-4 m3/s volume discharge) during the experiments. Due to the location, the effect of changing weather conditions and vegetation could be neglected. In the first phase, the variations of the water’s physico-chemical parametres were studied along a 400-m-long flow path. Water parametres were measured (temperature, 2+ - pH, concentration of Ca , HCO3 ) and modelled (SIcalcite, pCO2) at certain distances from the source of thermal water. In addition also a reactive transport model (PHREEQC) was carried out.

The results show the drop of pCO2, resulted in CO2 degassing and indirectly in pH increase and 2+ - then in the decrease of SIcalcite, Ca and HCO3 concentration (Fig. 24.). In the second phase thermal water was pumped into the canal for 12 weeks. Evolved precipitates were analysed on glass slides, which were put into the canal at certain distances from the source of thermal water. By the end of the experiment, red, amorphous, network- structured biogeochemical precipitate formed between 0 and 15 m, built-up by curly filamentous bacteria (Fig. 25.). After 15 m rhombohedral and dendritic calcite crystals formed (Fig. 26.).

Fig. 24. Results of the experimental study in the Gellért Tunnel; dashed line: measured parametre changes, solid line: modelled parametre changes (Kovács-Bodor et al., 2018b)

Fig. 25. Red, amorphous biogeochemical precipitate in the canal and its SEM image (Kovács-Bodor et al., 2018b)

Fig. 26. Calcite in the canal and the SEM image of dendritic calcite crystals (Kovács-Bodor et al. 2018b)

The biogeochemical precipitates were found to highly adsorb trace elements (like Mn, Al, Si) and radionuclides (226Ra and progeny) from the water already after 12 weeks (Kovács-Bodor et al., 2019b), as it was previously found by Erőss (2010) and Dobosy et al. (2016) in case of unknown aged precipitates.

2.4. Field stop 7: Aragonite Cave

There is only limited documentation about the caves of Gellért Hill. The reason is probably, that their dimensions are much smaller, and therefore cannot compete with the famous caves of the Rózsadomb sector. Korpás et al. (2002) in their comprehensive review on the geology of the area provided a summary also about the caves. Cavities open to the surface and small caves can be found at several levels of the Gellért Hill (Papp, 1942) (Fig. 27.). Most of them developed in Triassic dolomite, however there are also caves documented in Eocene-Oligocene marl (e.g. Citadella Cave, Leél-Őssy et al., 2009, Leél-Őssy, 2013) or in Eocene cherty breccia (Szent Iván and Aragonite Caves, Oravecz, 1970).

Fig. 27. Cavities of the Gellért Hill front wall (Virág et al., 2011)

The Gellért Hill caves are characterised by isometric spherical cavities (up to 12 m size) and fracture-related caves are reported (with 5-6 m length, up to 2 m width and up to 3 m height) (Korpás et al., 2002). The Gellért Hill caves are mainly isometric spherical cavities (up to 12 m size) with fracture- related caves and some elongated, flat “water-table caves” also reported (with 5-6 m length, up to 2 m width and up to 3 m height) (Korpás et al., 2002). There are two larger and well known caves at lower levels, but still above the today’s spring discharge level: the Szent Iván (120 asl) and directly underneath, the Aragonite Cave (109 asl) (Kadić, 1914; Kessler, 1963; Szablyár, 2000) (Fig. 28). Both caves are developed in Eocene breccia and can be characterised as horizontally elongated “water-table” chambers (Aragonite cave: width (w): 8 m, length (l): 10 m, height (h): 1.5 m; Szent Iván cave: w: 11 m, l: 12 m, h: 8 to 13 m). However, owing to constructional works the original morphologies today are obscured. Kessler (1963, 1965) and Szablyár (2000) attributed these caves to the dissolution effect of thermal waters. Kessler (1963) reported aragonite and gypsum minerals from the Aragonite Cave. Oravecz (1970) was the first to document the Aragonite cave and its minerals systematically and described red limonitic clay at the bottom of the cavity above which a thick isopachous calcite- crust developed. The Aragonite cave situated at about 15 to 20 metres above the recent karst water-table, was discovered in 1963 and made accessible in 1969 at the time of the construction of the water- collecting tunnel within the body of Gellért Hill. The originally 8x10x5 m size elongated cavity was partially destroyed by the mining activity but still the major features of dissolution and precipitation are observable. Based on detailed study of the mineralisation covering the cave-

35 walls Virág et al. (2013) and Pásztor (2016) attempted to reconstruct the paleo-hydrological story behind the paragenetic order of the mechanical and chemical infills of the cave (Fig. 29.).

Fig. 28. Cross section in the Gellért Hill along the St. Ivan and Aragonite caves after Szablyár (2000)

Fig. 29. Profile of the Aragonite cave (Pásztor, 2016)

The first infill is a reddish goethitic- clay which accumulated at the bottom of the cave. As shown by SEM observations, it is partly of biogeochemical origin, however, partly it contains also detrital grains identified as dissolution residue from the country-rock. Desiccation cracks of the red clay demonstrate that after the accumulation of the red sediment, the water-table must have descended so that the formerly underwater environment became dry. The fact that the desiccation cracks are filled by coarse-crystalline isopachous calcite (cave clouds or mammillary calcite) covering almost all the walls of the cavity and obviously having been formed under phreatic conditions, suggests that the dried-up cave must have been flooded again. U/Th age dating of the isopachous calcite suggests that this flooding event took place at about 130 kyr BP. The last members of the mineral paragenesis record the final decrease of the water-table. They consist of calcite-II, aragonite-needles, epsomite, huntite and gypsum – all formed under vadose conditions - and the sulphate minerals show that in addition to CO2, sulphur compounds were also present and contributed both to dissolution and mineral precipitation. At present the discharge level of almost all springs of the Gellért Hill area are situated in small caves close to the actual base-level of erosion at about 105 m asl. – Similar to the ancient Aragonite Cave these cavities are only 5 to 6 m long, up to 2 m wide and their vertical extent is about 3 m. They are usually connected to faults and fractures. Based on hydrogeological and speleo-morphological considerations and recent microbiological investigations (Borsodi et al., 2012), microbially mediated sulphuric acid speleogenesis is suggested as the dominant cave

36 forming process for the area. This process is further supported by recently forming gypsum crusts on the cave walls above the water table.

3. References

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Notes:

Day 2 stops 1-3: Hydrothermal alteration zones in outcrops of the porphyry-epithermal system of Palaeogene age at Recsk2

Ferenc Molnár (Geological Survey of Finland)

1. Introduction – regional geology and types of mineralisation

The Recsk Ore Complex is hosted by an intrusive-volcanic complex of Paleogene age. The complex is located in northeast Hungary and is a part of the ~300 km long, NE-SW oriented Hungarian Paleogene Volcanic Belt (Fig. 30.). This belt is the continuation of the Alpean granite- tonalite batolite belt along the Periadriatic Line. The ore complex at Recsk consists of high sulphidation type epithermal Cu-Au-Ag deposits at the Lahóca Hill – Lejtakna area near the village of Recsk and occurrences of intermediate-sulphidation type epithermal Au-Ag-Pb-Zn ores exposed near the village of Parád (Fig. 31.). Porphyry Cu-(Mo-Au), Cu-Zn(-Fe) skarn and carbonate hosted vein type and metasomatic-replacement ores are known at depth (Fig. 32.).

Fig. 30. Location of the intrusive-volcanic complexes at Recsk and Velence Mts. in the Alp-Carpathian Palaeogene Volcanic Belt.

The shallow epithermal ore zones were discovered in the late 18th century and exploitation from underground mine workings on Lahóca Hill continued sporadically until 1979. Reported mine production was 3.1 million tonnes at an average grade of 0.61% Cu and 2.5 g/t Au (Baksa et al., 1980). The porphyry Cu(-Mo-Au), Cu-Zn(-Fe) skarn and metasomatic-replacement Pb-Zn ores were discovered during a drilling program commenced in the late 1950s. Following 156,000 m of surface diamond drilling, development of a deep mine started in the 1980s with the construction of two 1,200 m deep shafts and 9.3 km of drive development on the -700 m ASL and -900 m ASL levels. Despite this extensive work the exploitation of the deep mineralisation has yet to commence (Földessy & Szebényi, 2008). The gold ore potential of the shallow epithermal zones was recognised again during the 1990s and at the same time scientific re-evaluation of the whole complex was started using the latest models of porphyry-epithermal systems in subduction-

2 This chapter of the guidebook was compiled mostly on the basis of Gatter et al. (1999), Molnár et al. (2008) and Takács et al. (2017).

43 related settings (Gatter et al., 1999; Molnár et al., 2003; Molnár, 2007). The current known resources of the Recsk Ore Complex are summarised in Table 1.

Fig. 31. Geology of the Recsk volcanic complexes (Quaternary cover removed) and locations of the field stops (from Molnár et al., 2008)

Due to its special geodynamic setting, the Recsk Ore Complex is one of the best preserved porphyry-skarn-epithermal systems in the Alpine-Balkan-Carpathian-Dinaride metallogenic

44 province, and as such provides a unique opportunity to study the relationships between the shallow and deep ore-forming environments (Molnár, 2007).

Fig. 32. Section across the northern part of the Recsk Ore Complex.

Ore Types Resources [Mt]* Grade (cut-off) Estimation based on data from Reference Cu-porphyry 109.4 0.96% (@ 0.8) Cu underground drillings Fodor et al. 1998 0.65% (@ 0.4) Cu Copper ore (porphyry and skarn) 779.3 surface and underground drillings Kontsek et al. 2006 0.19 g/t Au Zn-skarn 11.5 4.98 % Zn underground drillings 3.15-3.56% Zn Fodor et al. 1998 Pb-Zn metasomatic ore 36.6 surface drillings 1.19-2.15% Pb 1.26% Pb (+Ag content) Pb-Zn deep ore ("Recsk deep") 48.9 3.02% Zn underground drillings Kontsek et al. 2006 0.44% Cu HS epithermal ore 36.5 1.47 g/t (@ 0.59) Au surface drillings Kontsek et al. 2006

* - not compliant with internationally accepted resource classes; most closely correspond to indicated resources Table 1. Ore resources in the Recsk porphyry-skarn-epithermal ore complex.

2. Geology and mineralisation of the Recsk Ore Complex

The mineralised intrusive and volcanic units seen in outcrop and intersected in drill holes at Recsk and Parád are parts of a stratovolcanic complex with related subvolcanic intrusions. The outcropping volcanic units are exposed within an approximately 30 km2 large area, which is bounded by younger fault zones on the eastern and western sides, while the northern and southern boundaries are unexposed and largely unexplored (Fig. 31.). The stratovolcanic units are partly covered and surrounded by Late Oligocene and younger sedimentary rocks and by the Neogene volcanic complex of the Mátra Mountains to the south. The stratovolcanic units are deposited mostly on pelitic carbonate sedimentary rocks of Triassic-Jurassic age, but locally Late Eocene to Early Oligocene sedimentary units are also underlain or intercalated with the

45 volcanic and volcaniclastic rocks (Földessy, 1975; Less et al., 2008). The stratovolcanic succession consists of three major units: a submarine lava and hyaloclastite strata, a lacustrine-terrestrial stratovolcanic strata, and a subaerial lava-flow unit (Földessy, 1975; Földessy et al., 2008; Fig. 31.). The early stage eruptions with peperitic pyroxene-amphibole andesite, agglomerate and lava flows were followed by accumulation of biotite-amphibole andesite lava, agglomerates, volcanogenic breccias and other pyroclastic deposits at Lahóca Hill. To the west and south of Lahoca Hill dacite (“quartz-andesite”) dome-flow complexes formed (Pantó, 1952; Földessy et al., 2008; Molnár et al., 2008; Fig. 31.). The youngest volcanic unit consists of a biotite-amphibole andesite lava flow and agglomerate and covers the northeastern side of the Lahóca Hill (Fig. 31.). Porphyritic diorite stocks intrude the Mesozoic basement and the lower part of the stratovolcanic succession. They host Cu(-Mo-Au) ores extending from ~500 m to at least 1,200 metres depth and they are clustered in a zone trending NNE-SSW (see inset map with contours of the intrusive bodies on Fig. 31. and also see Fig. 32.). The contacts between the diorite intrusions and the host sedimentary units are marked by Cu-Zn-Au(-Fe) skarn and metasomatic Pb-Zn (carbonate-replacement) ores, which are widespread but erratically distributed. The strongly altered dacitic domes, lava flows and tuffs of the stratovolcanic unit to the west and south of Lahóca Hill (Fehérkő, Veresvár, Hegyes Hill, Macskabérc and Veresagyagbérc; Fig. 31) host up to 2 m thick siliceous veins and hydrothermal breccia dikes of the intermediate sulphidation (IS) epithermal Au-Ag-Pb-Zn mineralisation (Molnár, 2007; Molnár et al., 2008, Sillitoe, 2010). The zones of this kind of epithermal mineralisation are located above the southern apex of the porphyry Cu diorite stock. Intermediate argillic alteration, characterised by the presence of illite, illite with smectite interlayering and sericite, is the most widespread alteration type in the volcanic rocks of this area. Locally K-feldspar alteration is noticeable in dacite, dacite tuff and hydrothermal breccia. Intense silicification occurs in zones with sporadic pyritisation, where rock forming minerals are mostly replaced by mosaic-textured quartz and narrow (<0.1 mm) stockwork veinlets are filled with quartz. K-Ar dates for adularia and illite confirm the Oligocene age (~28 Ma) of the mineralisation (Molnár et al., 2008). The andesitic crystal tuff and diatreme breccia complex at Lahóca Hill hosts high sulphidation (HS) epithermal Cu-Au-Ag ore mineralisation within massive-brecciated sulphide bodies, vuggy silica bodies, various types of hydrothermal and volcanic breccias, and ore-bearing quartz stockwork (Molnár et al., 2008). The host volcanic units are characterised by advanced argillic alteration with Kandite-group clay minerals (kaolinite-nacrite-dickite) which enclose irregular- lentoid and dissected-brecciated vuggy silica bodies. The argillic alteration is associated with silicification and pyritisation of variable intensity. The youngest volcanic unit on the northeastern slide of Lahóca Hill is an exception, as it is not affected by hydrothermal alteration. Intermediate argillic alteration with illite locally occurs at lower elevations of the northern slope of Lahóca Hill and in the deeper parts of some drill holes (Molnár et al., 2008). The HS type ore forms two distinct groups of ore bodies, the Lahóca cluster consisting of eleven ore bodies and the Lejtakna ore body. The former is located peripheral to the northernmost apex of the known diorite stock while the Lejtakna ore body is situated above the northern part of the intrusion (see inset map on Fig. 31. showing the location of the orebodies). These orebodies have irregular-lentoid shape and consists of enargite-luzonite-pyrite ore. Sulphide minerals occur as disseminations and vug fillings in the vuggy silica bodies, as vein-fillings in stockwork ore, and as fillings in the open spaces between clasts in the hydrothermal breccias. Funnel-shaped diatrema- breccia as host rock to mineralisation is known in the southern part of the Lahóca Hill: this

46 funnel-shaped body is probably the major conduit of the volcanic explosions resulted in formation of the diatreme breccia bodies and acted as conduit for hydrothermal fluid flow. Both high and intermediate sulphidation epithermal mineralisation are characterised by the presence of liquid and semi-solid hydrocarbons as films and cavity fillings where porosity is higher. The timing of hydrocarbon migration and entrapment is unclear and may have been either late during the hydrothermal activity or during a younger event unrelated to mineralisation (Gatter et al., 1999; Molnár et al., 2008).

Mineral Lahóca Hill Lejtakna Parád area Name Formula Stage I. Stage II. Stage III. Stage I. Stage II. Stage III. list*

Pyrite FeS2 ? ? Sphalerite ZnS Galena Pb(Se,S) Clausthalite PbSe

Chalcopyrite CuFeS2

Tennantite Cu12As4S13 ?

Tetrahedrite Cu12Sb4S13

Goldfieldite Cu12Te4S13

Enargite Cu3AsS4

Luzonite Cu3AsS4

Famatinite Cu3SbS4

Stibnite Sb2S3

Chalcostibite CuSbS2

Emplectite CuBiS2

Bournonite PbCuSbS3 Unidentified Bi-Cu-Ag-Se-S mineral Unidentified Bi-Ag-Se mineral Unidentified Pb-Cu-S mineral Unidentified Sb-Bi-Cu-S mineral

Aikinite-Bismuthinite PbCuBiS3-Bi2S3

Wittichenite Cu3BiS3 Native gold Au(Ag)

Calaverite AuTe2

Hessite Ag2Te ?

Krennerite (Au,Ag)Te2

Sylvanite (Au,Ag)2Te4

Petzite Ag3AuTe2 ? Empressite AgTe

Bohdanowiczite AgBiSe2

Kesterite Cu2(Zn,Fe)SnS4

Berndtite SnS2

Stannoidite Cu6Cu2(Fe,Zn)3Sn2S12

Vinciennite Cu7Cu3Fe2Fe2Sn(As,Sb)S16 Unidentified Ag-Cu-S-Te mineral

Kawazulite Bi2Te2Se Vulcanite/Weissite CuTe Coloradoite HgTe Tiemanite HgSe Native tellurium Te

Acanthite Ag2S

Pyrargyrite Ag3SbS3

Tetradymite Bi2Te2S Marcasite FeS2 Table 2. Ore mineral parageneses in the HS type (Lahóca, Lejtakna) and IS type (Parád) epithermal zones of the Recsk ore complex (after Takács et al., 2017).

The sulphidation state of hydrothermal processes in the shallow epithermal zones in the Recsk Ore Complex is reflected by the ore mineral assemblages (Takács et al., 2017). In the HS type ores of the Lahóca Hill – Lejtakna area the most common ore minerals are enargite, luzonite, pyrite, tetrahedrite-tennantite, whereas enargite-luzonite occur in traces only and the

47 major ore minerals are tetraedrite-tennantite, pyrite, sphalerite in the IS type system near Parád. There are further differences in the assemblages of Au-Ag-telluride, Bi-telleuride and other sulphide accessory mineral assemblages (Table 2.), as well as in the compositions of tetrahedrite-tennantite group minerals. In the HS system at the Lahóca Hill, this latter group of minerals tends to be more Ag and Te rich comparing to the IS system at Parád. Variation of sulphidation state within the HS type ore is also reflected by the differences in accessory mineral assemblages among stages of ore forming processes (Takács et al., 2017; Fig. 33.).

Fig. 33. Variation of the sulphidation state during the stages of hydrothermal processes in the HS type epithermal mineralisation of the Lahóca Hill (from Takács et al., 2017)

Fluid inclusion studies were applied to temperature-pressure-composition the characterisation of hydrothermal fluids (Monár et al., 2008). The abundance of coarse-grained enargite-luzonite in the HS ore and transparency of these minerals in infrared light allowed to gather fluid inclusion data directly from the major sulphide components of the ore. From the IS ore, fluid inclusion data were collected from hydrothermal quartz only. According to the results of fluid inclusion studies, the palaeotemperature of fluids has been varied between 300 and 150°C at the currently exposed levels of the HS and IS type epithermal systems of the Recsk Ore Complex (Fig. 34.). However, the ore and post-ore stages in the HS system are characterised by elevated (up to 10-15 NaCl equiv. wt %) salinities of fluids, whereas these high salinity fluids are absent in the fluids with similar temperatures in the IS system. In the IS system the slight increase of salinities to max. about 5 NaCl equiv. wt.% salinities in the low temperature fluids can be explained by boiling of hydrothermal fluids. The much higher salinities in the HS system may reflect upwelling of high salinity magmatic fluids from the deep porphyry copper system into the shallow volcanic zones during the evolution of the HS system. However, this latter model needs further constraints from studies of stable isotopes or sulphide mineral trace elements.

Fig. 34. Fluid inclusion data for the HS and IS type epithermal zones of the Recsk porphyry-epithermal system (modified after Molnár et al., 2008).

3. Field stops in the epithermal zones of the Recsk Ore Complex

3.1. Field stop 1: advanced argillic alteration zone in the HS type epithermal system on the Lahóca Hill, northern gulley

On the northern slope of the Lahóc Hill (Fig. 31.), the intensely altered dacite crystal tuff host rock of the HS type ore crops out. In the least altered zones, the andesitic crystal tuff consist of plagioclase, biotite and amphibole phenocrysts, as well as andesitic lithoclasts in fine grained, rock flour matrix. The grain size of phenocrysts is up to 1-1.5 cm. Large plagioclase crystals are mostly euhedral, however, smaller grains are crushed and there is a continuous range of size of grains from the euhedral large crystals to the microscopic (<0.1 mm) crystal fragments. Biotite and amphibole are usually euhedral and coarse grained (0.5-1.5 cm). A very few crushed quartz fragment (0.2-0.5 cm) also occur in the tuff.

The andesite crystal tuff is generally characterised by advanced argillic alteration with variable intensity of silicification and pyritisation in the highly altered zones around the orebodies. Plagioclase, biotite and amphibole phenocrysts and crystal fragments are usually replaced by clay minerals whereas the groundmass is impregnated by mosaic-textured microcrystalline quartz. By increasing of intensity of silicifation some of the phenocrysts are also replaced by mosaic-textured quartz in the vicinity of veinlets and breccias. Pyritisation is typical for mafic rock forming minerals, but pyrite is also present in the groundmass. Like in the ore bearing structures, pyrite disseminations in the crystal tuff consists of two types of pyrite: euhedral cubes and round-framboidal gel-pyrite with 0.1-1mm grain sizes. Pyrite content of altered rocks can be as high as 10% especially in the proximity of ore-bearing structures. In association with pyrite, a few grains of enargite, grey ore and barite was also found by SEM-EDS studies. The most common clay mineral replacing phenocrysts belong to the Kandite-group according to results of X-ray diffraction analyses on separated and oriented samples. Kandite minerals are indistinguishable by routine X-ray diffraction methods, therefore results were refined by IR- spectroscopic analyses. Results indicate that the most common clay mineral is dickite.

3.2. Field stop 2: Etelka pit and adit, northern slope of the Veresvár Hill

The IS type peithermal gold mineralisation is hosted by dacite dome-flow complexes in the southern part of the mineralised area (Fig. 31.). Dacitic lava rocks occur at several places, but mostly at the higher elevations (above 300 m) of the Veresvár, Veresagyagbérc and Hegyeshegy Hills (Fig. 31.). The most common rock forming minerals in the porphyritic rock are feldspars which form subhedral-euhedral up to 2 cm large prisms. Two types of feldspars can be distinguished in the less altered rocks: plagioclase with andesine-labradorite compositions and sanidine. Less common phenocrystst (5-10 mm) are biotite and amphibole (always strongly altered) and rock forming quartz. Quartz is euhedral-hexagonal or round-resorbed. The groundmass is always silicified and argillitised and only ghosts of plagioclase microcrysts can be recognised at some places. Typical accessory minerals are apatite and zircon. Dacitic tuff/pyroclastic deposits occur mostly at the lower elevations of the area. Their common feature is the presence of large amount of crystal fragments (with up to 10-15 mm sizes) consisting of mostly feldspars (up to 25-30%) and less abundant amphibole and biotite phenopcrysts (less than 5%). The appearance of the rock resembles to the crystal tuff of the Lahóca Hill at some places. The most important difference is the occurrence of hexagonal quartz crystals and fragments (less than 5%) and the higher abundance (up to 5-10%) of lithoclasts. There are two different types of lithoclasts (up to 5 cm size) in the tuff: one of them has dark grey matrix with plagioclase phenocrysts (andesitic lava rock) and the other has very fine grained texture without pheocrysts (always altered). The groundmass (up to 50-60%) is very fine grained and usually silicified in the form of mosaic-crystalline (<0.1 mm) quartz or altered to a siliceous- argillic material. In some areas (e.g. the western side of the Veresvár Hill) the tuff has well bedded-layered appearance – at these places abundances of phenocrysts decreases and the original rock was probably a well bedded glassy tuff. Bedding cannot be recognised in the scale of most of the outcrops. In the volcanic rocks of the IS type apithermal zones of the Parád the most widespread alteration is the intermediate argillic alteration with illite, sericite (K-mica), and illite-smectite mineralogy. Argillic alteration is more intense in the tuffaceous units. Kaolinitisation was observed in a very few samples, especially in zones where the oxidation of high pyrite content resulted in supergene acidic conditions (kaolinite is associated with alumn, gypsum and alunite).

The sericite, illite and illite-smectite alteration affected mostly rock forming feldspars in dacite, in some fragments of hydrothermal breccia and groundmass of the tuff. Sericite (K-mica with very sharp peaks with narrow half-width on oriented XRD pattern) and illite (sharp peaks, but somewhat broader half-width in comparison to sericite on oriented XPD pattern) are well crystallised and where illite contains smectite interlayering the smectite content does not exceed 10% according to calculations based on Eberl et al. (1987). This indicates that illite alteration took place on a temperature higher than 200-220°C (Parry, 2002). Rock forming plagioclase also show K-metasomatic alteration in some dacite and dacite tuff samples, as well as fragments of hydrothermal breccia, however, it is largely masked by the overprinting sericitic alteration. Carbonate alteration of feldspars was occasionally detected on the northern slope of the Veresvár. Mafic rock forming minerals (amphibole, biotite) are replaced by fine-grained mass of argillic material, pyrite and mosaic-textured quartz. Pyritisation is also typical for the groundmass of rocks; pyrite occurs as dissemination of euhedral (hexahedron) crystals (0.1-1 mm) and framboidal/gel-pyrite patches (0.1-1 mm). Pyrite content is usually less than 1-2%, however, by increase of intensity of silicification pyrite may be more abundant (up to 10-15 %). Silicification affected mostly the groundmass of the rocks in the form of fine (<0.1 mm) mosaic-textured aggregates. In more intensely silicified zones, rock forming minerals are also replaced by mosaic- textured quartz and thin occurrence of thin (<0.1 mm) quartz stockwork is also typical. The siliceous veins and irregular masses of strongly silicified rocks which are surrounded by the agillic alteration zones usually consist of fine-grained, sugar textured massive to banded and brecciated white quartz haloed by a stockwork of thin (1-2 cm thickness) veinlets of milky to comb textured quartz. Comb-textured coarse grained (up to 1-2 cm grain size) quartz texture also occurs in drusy and ruptured parts of the veins. On the Hegyeshegy (Fig. 31.), banded texture of fine grained quartz vein with pseudomorphs after bladed calcite also occurs. Boundaries of veins are usually gradational to the host rocks by decrease of intensity of silicification: sharp walled banded-symmetric veins were described in some parts of veins in old adits of the Etelka pit only. Small amounts (5-10%) of disseminated sulphide minerals are present in several parts of veins, but their distribution is generally erratic. The most common sulphide minerals are tetrahedrite, sphalerite, galena, pyrite and chalcopyrite. These minerals occur as euhedral to anhedral grains (0.01-1 cm) or patches consisting of aggregates of grains (max. 1- 2 cm) in the fine grained quartz matrix. Nagy (1985) described occurrences of Au-, Ag- and Bi- telluride minerals in the veins in the Etelka pit of the Veresvár Hill. The Etelka pit and the short adit at the bottom of the pit was opened and mined at the turn of the 18th and 19th centuries and produced silver bearing copper ores (e.g. tetraedrite rich ores were exploited). In the 1950`s, the pit and adit was re-opened during the exploration of the area, but the mineralisation was found to be sub-economic.

3.3. Field stop 3: hydrothermal breccia dike, Timsós pit, Ilona Valley

The Ilona Valley forms the western tectonic boundary of the IS type epithermal zones south of Parád (Fig. 31.). This fault was probably also active during the hydrothermal processes as parllell with it a hydrothermal dike can be followed for about 1 km in a N-S trending zone. Thickness of dike is up to 6-8 metres in some outcrops. This breccia dike at this field stop is characterised by highly variable texture from matrix- supported to clast-supported character and composition and roundness of fragments. In general, the matrix is fine grained (0.01 mm) mosaic-textured quartz with variable clay content

(sericite, illite, smectite). The matrix also contains euhedral-pseudorhombic adularia, calcite and epidote at some places. Presence of pyrite disseminations (euhedral-subhedral cubes and framboidal bands, patches) is ubiquitous and occasional occurrences of tetrahedrite and sphalerite grains (0.01-1 mm) are also common. Tetrahedrite is Ag-bearing (around 1 wt%), Zn- rich (6-7 wt%) and has variable As-content (up to 12 wt%). Sphalerite is Fe-poor (<1 wt%) and has elevated Mn-contents (2-4 wt%). The clasts of the dike consist of variably altered (silicic, sericitic, illitic, carbonatic, K- metasomatised) fragments of porphyritic andesite/diorite, andesite with oriented flow texture, dacite and dacite tuff. Sericitic alteration is also characteristic to siliceous shale fragments with well observable lamination-foliation. Fragments of rock forming minerals (feldspar, biotite, amphibole) also occur in fine grained breccias. There are also fragments in which complete silicification washed away the original texture. Porphyritic diorite/andesite and sericitic shale fragments are probably originated from the subvolcanic intrusion and basement units which are located at about 800 m depth in the southern part of the Ilona Valley area. The sufficient vertical transport of fragments is also expressed by their round and round-flattened shapes. Less rounded fragments consist of dacite and dacite tuff. Size of fragments varies from 1 mm to 5 cm. In the R-424 drillhole, which was completed during exploration for epithermal gold in the 1990`s and followed the dike along dip (Fig. 35.), vertical sorting of grain size of fragments (increase of grain size from about 1mm to about 2-4 cm) was locally also observed. There are also zones where sorting and orientation of clasts suggest to fluidisation processes. In other parts of the drillhole, the matrix-supported and clast-supported character of breccia is highly variable and sometimes the change is sharp. The complex textural variation suggests that the breccia vent was filled up by a silica rich muddy matrix material which was intruded by pulses of upward moving packages of fragments. Some packages achieved vertical sorting in clam periods of breccia flow and remained intact by later pulses, and some packages were mixed up with each other in the still semi-soft state of the matrix.

Fig. 35. Zoning of metal concentration and hydrothermal mineralogy as a function of depth in the hydrothermal breccia of the Ilona Valley.

Solidification of the matrix of the breccia was associated with precipitation of quartz and other silicates, sulphides and gold. This is well expressed by the vertical zoning of matrix mineralogy and variation in gold content in the R-424 drillhole (Fig. 35.). The observed vertical zoning of matrix mineralogy is typical for low to intermediate sulphidation type of epithermal systems. Concentration of gold is increased in the zone where adularia is present in the mineralogy of the matrix. This is a common feature in a low-sulphidation type system where zones of adularia precipitation marks boiling of fluids and boiling is one of the major factors driving precipitation of gold.

4. References

Baksa, Cs, Cseh-Német, J, Csillag, J, Földessy, J, Zelenka, T (1980): The Recsk porphyry and skarn copper ore deposit, Hungary. In : Jankovic, S & Sillitoe, R (eds.): European copper deposits. Belgrade University, 73–76 Eberl DD, Srodon, J, Lee, M, Nadeau, PH, Northrop, HR (1987): Sericite from the Silverton caldera, Colorado: Correlation among structure, composition, origin, and particle thickness. Am. Min., 72:914-934 Fodor, B, Forgács-Gombár, G, Káli, Z (1998): Mineral resources and reserves of Hungary. Ann. Rep. Hung. Geol. Surv. 1998: 261 p. (in Hungarian) Földessy, J (1975): The stratovolcanic andesite formation of Recsk. Bull. Hung. Geol. Soc., 105: 625–645 (in Hungarian) Földessy, J & Szebényi, G (2008): The mineralizations of the Recsk deeps and Lahóca—short geological overview. In: Földessy, J & Hartai, É (eds.): Recsk and Lahóca geology of the Paleogene ore complex. University of Miskolc, Series A, Mining, 73: 85–98 Földessy, J, Hartai, É, Kupi, L (2008): New data about the Lahóca high-sulphidation mineralization. In: Földessy, J & Hartai, É (eds.): Recsk and Lahóca geology of the Paleogene ore complex. University of Miskolc, Series A, Mining, 73: 129–143 Gatter, I, Molnár, F, Földessy, J, Zelenka, T, Kiss, J, Szebényi, G (1999): High- and low-sulfidation type epithermal mineralization of the Mátra Mountains, northeast Hungary. Soc. Econ. Geol., Guidebook Ser., 31: 155–179 Kontsek, T (ed.) (2006): Mineral resources and reserves of Hungary: Annual Report of the Hungarian Geological Survey. Budapest, Hung. Geol. Surv., 115–119 (in Hungarian) Less, Gy, Báldi-Beke, M, Pálfalvi, S, Földessy, J, Kertész, B. (2008): New data on the age of the Recsk volcanics and of the adjacent sedimentary rocks In: Földessy, J & Hartai, É (eds.): Recsk and Lahóca geology of the Paleogene ore complex. University of Miskolc, Series A, Mining, 73: 99–128 Molnar, F (2007): The Cu-Au-Ag-Zn-Pb ore complex at Recsk: A uniquely preserved and explored porphyry-skarn-epithermal system in the Paleogene magmatic belt of the Alp-Carpathian- Dinaride system. In: Andrew, CJ et al. (eds.): Mineral exploration and research: Digging deeper, Biennial SGA Meeting, 9th, Proc., 1: 153–157 Molnar, F, Gatter, I, Zelenka, T, Pécskay, Z, Bajnóczy, B (2003): Metallogeny of Paleogene and Neogene volcanic belts in Hungary. In: Eliopoulos, DG (ed.): Mineral exploration and sustainable development. Biennial SGA Meeting, 7th, Proc., 2: 1205–1208 Molnar, F, Jung, P, Kupi, L, Pogány, A, Vagó, E, Viktorik, O, Pécskay, Z, Hurai, V (2008): Epithermal zones of the porphyry-skarn-epithermal ore complex at Recsk. In: Földessy, J & Hartai, É (eds.): Recsk and Lahóca geology of the Paleogene ore complex. University of Miskolc, Series A, Mining, 73: 99–128

Nagy, B (1985): Gold, silver, and bismuth telluride in the paragenesis of the Parádfürdö mineralization. Ann. Rep. Hung. Geol. Inst. 1983, 321–357 (in Hungarian) Pantó, G (1952): Mining geology survey around Recsk and Parad. Ann. Rep. Hung. Geol. Inst. 1949, 67–75 (in Hungarian). Parry, WT, Jasumback, M, Wilson, PN (2002): Clay mineralogy of phyllic and intermediate argilic alteration at Bingham, Utah. Econ. Geol. 97: 221–239 Sillitoe, RH (2010): Porphyry copper systems. Econ. Geol., 105: 3–41 Takács, Á, Molnár, F, Mogessie, A, Menzies, JC (2017): Ore Mineralogy and Fluid Inclusion Constraints on the Temporal and Spatial Evolution of a High-Sulfidation Epithermal Cu-Au-Ag Deposit in the Recsk Ore Complex, Hungary. Econ. Geol., 112/6: 1461-1481

Notes:

Day 2 stop 4: Darnó Hill, Nagy-Rézoldal quarry: submarine hydrothermal alteration of Triassic pillow basalt3

Gabriella B. Kiss (Eötvös Loránd University, Department of Mineralogy)

1. Regional geology

The basement of the Pannonian Basin consists of a mosaic-like pattern of allochtonous terranes derived from different parts of the Tethyan realm. The major structure of the basement is the Mid-Hungarian (Zagreb-Zemplén) Lineament which divides the Tisza Megaunit and the Pelso Megaunit, a part of the ALCAPA (Csontos, 1995). The Darnó Unit is a part of the Bükk Structural Unit and is located in the Pelso Megaunit. The ca. 7 km2 area can be interpreted as a continuation of the structural elements of the Bükk Unit in a nappe system (Kovács et al., 2008). The lowest unit is the “Bükk Parautochton”, containing Palaeozoic – Upper Jurassic formations. Above it, the Mónosbél Unit can be found with mainly Jurassic deep sea sediments. The Szarvaskő Unit is the next nappe, containing a Jurassic, incomplete ophiolitic sequence together with deep sea sediments, while on the top the Darnó Unit can be found with Triassic and Jurassic submarine volcanites and related sediments (Fig. 36, 37.) (Csontos, 1999, Haas and Kovács, 2001, Kovács et al., 2008).

Kovács et al., 2008 Fig. 36. The nappe system of the Bükk Unit

Several studies have dealt with the submarine basaltic rocks of the Darnó Unit (Fig. 2.) (e.g. Papp, 1938; Mezősy and Grasselly, 1949; Kiss, 1958; Buda & Kiss, 1980; Balla et al., 1980; Kubovics, 1984; Balla, 1987; Kubovics et al., 1990; Dosztály & Józsa, 1992, Józsa, 1999, Kovács et al., 2008, among others) and they arrived to controversial conclusions about the origin of these rocks. However, many of the authors recognised their different characteristics from mid-oceanic ridge basalts and from the nearby occurring Jurassic basalts of the Szarvaskő Unit. Similar origin and possible relationship between these submarine basaltic rocks and occurrences in the Dinarides and Hellenides were found on the basis of stratigraphy, petrochemistry and tectonic evolution of the region (Downes et al., 1990; Harangi et al., 1996; Józsa 1999; Haas & Kovács 2001). Recognition of the occurrence of volcanic units with peperitic facies called again the attention to this area and the most recent studies have suggested that the formation of these basaltic rocks is related to the Triassic advanced rifting stage of the Neotethys (Kiss et al., 2008, 2012; Kovács et al., 2008, 2010; Haas et al., 2010).

3 This chapter of the guidebook is based mostly on Kiss et al. (2010, 2012 and 2016). 55

Darnó Hill

Recsk RM-135 Nagy- Rézoldal quarry

Eger

RM-131 1 km

Földessy, 1975

Triassic and Miocene sediment Oligocene Jurassic basalt and volcanics sediment Quaternary Triassic and Jurassic cover shale, radiolarite Borehole

Fig. 37. Geological map of the Darnó Hill. The location of field stop "Nagy-Rézoldal quarry" is also shown.

2. Proximal and distal volcanological facies in a submarine basaltic lava flow

The structure of submarine basaltic lava flows can be reconstructed as follows (Goto & McPhie, 1998, 2004, Palinkaš et al., 2008) (Fig. 38.).

Lava flow

2008 Fig. 38. Structure of a submarine lava flow with the occurring volcanological facies

In the centre of the lava flow, coherent pillow lava (1) is presented. This represents the main feeder zone of the volcanism, where cooling of the basalt caused polygonal columnar joints. In this facies, no inter-pillow hyaloclastite breccia occurs as no magma–water interaction

56 happened at the time of solidification of the basalt. This facies is surrounded by the closely packed pillow lava (2) facies, where among the pillows some hyaloclastite breccia can be found. The breccia forms when the lava cools down fast, so the densely packed lobes will be surrounded by chilled glass. The amount of the inter-pillow hyaloclastite breccia is increasing as a function of distance from the centre of the lava flow. Hydrothermal infillings, such as amygdales, cooling-related jig-saw veins and former feeding channels can occur in this facies. Where the lava arrived into water-soaked sediment, peperitic facies (3) formed. In this facies, the basalt mingles with the sediment, thus the unconsolidated carbonate mud infiltrates into the cooling cracks of the basalt, admixes to the hyaloclastite breccias or larger semiconsolidated sedimentary blocks are mixed with the basalt in the form of anhedral clasts. On the seafloor, around the closely packed pillow, where very rapid cooling took place, in situ hyaloclastite (4), then pillow fragmented hyaloclastite (5) and finally isolated pillow breccia (6) developed as a function of distance from the closely packed pillow facies. All of these facies contain mainly pieces of basalt lobes and glassy shards, cemented by hydrothermal minerals. The in situ hyaloclastite contains less cement and bigger fragments, while the isolated pillow breccia has the lowest amount of small (few centimetres large) basalt fragments.

3. Submarine hydrothermal processes in the Darnó Unit

Hydrothermal alteration minerals are abundant at the Triassic basalt localities. The occurrences can be categorised into three main groups, according to the model of Hart (1973); (1) primary hydrothermal alteration, (2) cooling related hydrothermal alteration and (3) low- temperature hydrothermal alteration. The series of mineral precipitation as well as the temperature ranges are shown in Fig. 39., based on the study of several different natural and artificial outcrops of the Darnó Hill.

Fig. 39. Series of hydrothermal mineral precipitation in the Triassic basalts of the Darnó Unit. A series of three sequential processes (1. primary hydrothermal alteration; 2. cooling related hydrothermal alteration and 3. low temperature hydrothermal alteration) can be recognised.

The important determining factors of this small scale hydrothermal system are the extremely rapid cooling of the hydrothermal fluid, the dominance of slightly modified seawater as a

57 hydrothermal fluid and the dependence on the local, effective water/rock ratio (i.e. distal/proximal position within the lava flow). The observed features are typical of the volcanic facies, thus may also help in identification of the spatial relationships within a submarine lava flow. The overall characteristics of these local, submarine basaltic volcanic centre related hydrothermal systems are markedly distinct from the typical mid-oceanic ridge or back-arc-basin opening related, volcanogenic massive sulphide deposit forming large scale hydrothermal systems (see e.g. Bodnar et al., 2014 and the references therein).

4. Field stop 4: the Nagy-Rézoldal quarry

The Nagy-Rézoldal quarry is situated on the western part of the Darnó Hill, about 3 km east from Recsk (N 47.9273; E 20.1353) (Fig. 40.). Fig. 5/A

Fig. 40. Panoramic picture of the Nagy-Rézoldal quarry (A-B), sketch of the observable volcanological facies and hydrothermal phenomenon (C) and field photos of a pillow with jig-saw veins (D) and small amygdales (E).

Mainly blocks of greenish and greenish grey basalt are exposed in the quarry, though smaller blocks of possibly Jurassic sedimentary rocks (radiolarite, schist) can be found in its northern edge as well. The basalt blocks are in tectonic contact with each other and sometimes even the directions of the younger pillow is different in the adjoining blocks. Though almost all blocks represent the closely packed pillow facies, small and large amygdale containing as well as jig- saw vein containing sub-types can clearly be distinguished. One block of pillow fragmented hyaloclastite breccia is found in the middle of the quarry. Thus, only distal facies of the submarine lava flow crop out. The 30-60 cm big pillows contain amygdales with variable sizes (from 0.2 to 8 mm) and jig- saw veins in the various sub-types of the closely packed pillow facies. Basalt texture varies between sphaerolitic, variolitic and intersertal types. Besides the microcrystalline-chloritised ground mass, phenocrysts of plagioclase (with an average composition of

Na0,85Ca0,07Al1,08Si2,95O8), as well as pyroxene laths (with an average composition of Ca0,9Mg0,58Fe0,35Ti0,15Cr0,01Mn0,01Si1,65Al0,38O6) and calcite-chlorite pseudomorphs after olivine also occur in the basalt. Based on chlorite thermometry calculations (Zang & Fyfe, 1995), chlorite occurring in the groundmass formed at an average temperature of 222°C. The 0.2-0.8 mm thick cooling-related cross-cutting veinlets and the 1-3 mm thick jig-saw veins are filled by calcite and small amount of chlorite, quartz, prehnite and pumpellyite. Chlorite thermometry calculations (Zang & Fyfe, 1995) yielded an average formation temperature of 216°C for early chlorite occurring along the walls of the veins. The smaller (0.2-2 mm) amygdales are filled up by chlorite, chalcedony and calcite, whereas the larger (2-8 mm) ones contain mainly calcite with some early quartz and chlorite along the walls. Chlorite thermometry calculations (Zang & Fyfe, 1995) yielded an average formation temperature of 182°C for smaller and 213°C for larger amygdale filling early chlorite. Fluid inclusion microthermometry of large amygdale filling calcite yielded a homogenisation temperature of Th(LV-L)=102-180°C (average: 136°C) and a salinity of 2.9- 4.49 NaCl equiv. wt% (average: 3.49 NaCl equiv. wt%) (Fig. 41.). These data suggest upheated seawater related origin in a rapidly cooling environment. Interpillow hyaloclastite occurs in small amount. Its matrix is mostly calcite, while the fragments are glassy basaltic shards.

Homogenisation temperatures measured in the Homogenisation temperature vs. salinity diagram amygdales at Nagy-Rézoldal Quarry, Darnó Unit (amygdale, Nagy-Rézoldal Q., Darnó Unit)

n ) V

t m 12 10

w L

r L 10 .

s 8

8 u

e V

6 e

l 6

f 5 μm 4 C

r 4 2 N

e (

0 y

i m 2

9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 9 0

1 2 3 4 5 6 7 8 9 0 1 2 3 4 5 6 7 8 9 0 1 2 3 4 5 6 7 8 0

0 ------

1 1 1 1 1 1 1 1 1 1 2 2 2 2 2 2 2 2 2 3

0 0 0 0 0 0 0 0 0 ------

S 0

1 2 3 4 5 6 7 8 9 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0 0

0 1 2 3 4 5 6 7 8 9 0 1 2 3 4 5 6 7 8 9 0 50 100 150 200 250 300

1 1 1 1 1 1 1 1 1 1 2 2 2 2 2 2 2 2 2 2 Homogenisation temperature ranges (oC) Th (oC)

Fig. 41. Results of fluid inclusion microthermometry study.

A 5 m large block of pillow fragmented hyaloclastite breccia occurs in the middle of the quarry. The up to 30 cm large basaltic fragments are mixed with strongly hematitised glassy- chloritised material. The matrix is mainly calcite but some chlorite is also presented. Whole rock geochemical study clearly shows the effects of submarine hydrothermal processes (e.g. the behaviour of mobile elements). However, differences among this Triassic and the nearby Jurassic basalts are obvious. Within-plate magmatism related origin is suggested based on whole rock geochemical as well as pyroxene compositional discrimination diagrams (Pearce & Cann, 1973, Meschede, 1986, Nisbet & Pearce, 1977).

5. References

Balla, Z, Baksa, Cs, Földessy, J, Havas, L Szabó, I (1980): The Tectonic setting of the Ophiolites in the Bükk Mountains (North Hungary). Geol. Carpath., 31/4: 465-493 Balla, Z (1987): A Bükk-hegység mezozoós tektonikája és kapcsolata a Nyugati-Kárpátokkal és a Dinaridákkal (Tectonics of the Bükkian (North Hungary) Mesozoic and relations to the West Carpathians and Dinarides). Ált. Földt. Sz., 22: 13-54 Bodnar, R. J., Lecumberri-Sanchez, P., Moncada, D., Steele-Macinnis M. (2014): Fluid Inclusions in Hydrothermal Ore Deposits. In: Holland, H.D. & Turekian, K.K. (eds.): Treatise on Geochemistry, Second Edition, Oxford: Elsevier, 13: 119-142 Buda, Gy & Kiss, J (1980): Comparison some chromite and titaniferous magnetite, ilmenite ore bearing ultrabasic-basic complexes. UNESCO Int. Symp. Athens, 1: 21-45

Csontos, L. (1995): Tertiary tectonic evolution of the Intra-Carpathian area: A rewiev, Acta Vulc., 7/2: 1-13 Csontos, L (1999): A Bükk-hegység szerkezetének főbb vonásai (Main structural features ot eh Bükk Mts.). Földt. Közl., 129/4: 611-651 Dosztály, L & Józsa, S (1992): Geochronological evaluation of Mesosoic formations of Darnó Hill at Recsk on the basis of radiolarians and K-Ar age data. Acta Geol. Hung., 35/4: 371-393 Downes, H, Pantó, Gy, Árkai, P, Thirlwall, MF (1990): Petrology and geochemistry of Mesozoic igneous rocks, Bükk Mountains, Hungary. Lithos, 24/3: 201-215 Goto, Y & McPhie, J (1998): Endogenous growth of a Miocene submarine dacite cryptodome, Rebun Island, Hokkaido, Japan. J. Volc. Geotherm. Res., 84/3–4:273–286 Goto, Y. & McPhie, J. (2004): Morphology and propagation styles of Miocene submarine basanite lavas at Stanley, northwestern Tasmania, Australia. J. Volc. Geotherm. Res., 130/3– 4:307–328 Haas, J & Kovács, S (2001): The Dinaric-Alpine connection – as seen from Hungary. Acta Geol. Hung., 44/2-3: 345-362 Haas, J, Kovács, S, Karamata, S, Sudar, M (2010): Displaced South Alpine and Dinaridic elements in the Mid-Hungarian Zone. Bulletin T. CXL Classe des sciences mathématiques et naturelles, Sciences naturelles, 46: 81-103 Harangi, Sz, Szabó, Cs, Józsa, S, Szoldán, Zs, Árva-Sós, E, Balla, M, Kubovics, I (1996): Mesozoic Igneous Suites in Hungary: Implications for Genesis and Tectonic Settings in the Northwestern Part of Tethys. Int. Geol. Rev., 38: 336-360 Hart, R. A. (1973): A model for chemical exchange in the basalt-seawater system of oceanic layer II. Can. J. Earth. Sci., 10: 799–816 Józsa, S (1999): A darnó-hegyi óceánaljzati magmás kőzetek petrológiai-geokémiai vizsgálata – PhD Thesis manuscript, Eötvös Loránd University (Petrological and geochemical analysis of the submarine igneous rocks of the Darnó Hill) Kiss, J (1958): Ércföldtani vizsgálatok a Darnó-hegyen (Ore geological investigations on the Darnó hHll). Földt. Közl., 88: 27-41 Kiss, G, Molnár, F and Palinkaš, LA (2008): Volcanic facies and hydrothermal processes in Triassic pillow basalts from the Darnó Unit, NE Hungary. Geol. Croat., 61/2-3: 385-394 Kiss, G; Molnár, F ; Kovács, S ; Palinkaš, L (2010): Field characteristics and petrography of the advanced rifting-related Triassic submarine basaltic blocks in the Jurassic mélange of the Darnó Unit. Centr. Eur. Geol., 53/2-3: 181-204 Kiss, G, Molnár F, Palinkaš LA, Kovács S, Hrvatović, H (2012): Correlation of Triassic advanced rifting-related Neotethyan submarine basaltic volcanism of the Darnó Unit (NE-Hungary) with some Dinaridic and Hellenidic occurrences on the basis of volcanological, fluid–rock interaction, and geochemical characteristics. Int. J. Earth Sci., 101/6: 1503-1521 Kiss GB, Molnár, F, Palinkaš LA (2016): Hydrothermal processes related to some Triassic and Jurassic submarine basaltic complexes in northeastern Hungary, the Dinarides and Hellenides. Geol. Croat. 69/1: 39-64 Kovács, S, Haas, J, Szebény,i G, Gulácsi, Z, Pelikán, P, B.-Árgyelán, G, Józsa, S, Görög, Á, Ozsvárt, P, Gecse, Zs, Szabó, I (2008): Permo-Mesozoic formations of the Recsk-Darnó Hill area: stratigraphy and structure of the pre-tertiary basement of the paleogene Recsk orefield. In: Földessy J., Hartai É. (eds.): Recsk and Lahóca Geology of the Paleogene Ore Complex, Geosciences, Publications of the University of Miskolc, Series A, Mining, 73: 33-56 Kovács, S, Haas, J, Ozsvárt, P, Palinkaš, LA, Kiss, G, Molnár, F, Józsa, S, Kövér, Sz (2010): Re- evaluation of the Mesozoic complexes of Darnó Hill (NE Hungary) and comparisons with

Neotethyan accretionary complexes of the Dinarides and Hellenides: preliminary data. Centr. Eur. Geol., 53/2-3: 205-231 Kubovics, I (1984): On the petrogenesis of the North Hungarian basic-ultrabasic magmatic rocks. Acta Geol. Hung., 27/1-2: 163-189 Kubovics, I, Szabó, Cs, Harangi, Sz, Józsa, S (1990): Petrology and petrochemistry of Mesozoic magmatic suites in Hungary and adjacent areas – an overview. Acta Geod. Geoph. Mont. Hung., 25/3-4: 345-371 Meschede, M. (1986): A method of discriminating between different types of mid-ocean ridge basalts and continental tholeiites with the Nb-Zr-Y diagram. Chem. Geol., 56: 207-218 Mezősi, J & Grasselly, Gy (1949): The occurrence of Native Copper in the Mátra Mountains at Bajpatak. Acta Miner. Petrogr., III: 44-47 Nisbet, EG & Pearce, JA (1977): Clinopyroxene Composition in Marie Lavas from Different Tectonic Settings. Contrib. Miner. Petrol., 63: 149-160 Palinkaš, AL, Bermanec, V, Borojević-Šoštarić, S, Kolar-Jurkovšek, T, Strmić-Palinkaš, S, Molnár, F, Kniewald, G (2008): Volcanic facies analysis of a subaqueous basalt lava-flow complex at Hruškovec, NW Croatia-evidence of advanced rifting in the Tethyan domain. J. Volc. Geotherm. Res., 178/6: 44–656 Papp, F (1938): Notes sur les minerais de Recsk. Földt. Közl., 68/7-9: 208-214 Pearce, JA & Cann, JR (1973): Tectonic setting of basic volcanic rocks determined using trace element analyses. Earth Planet. Sci. Lett., 19: 290-300 Zang, W & Fyfe, WS (1995): Chloritization of the hydrothermally altered bedrock at the Igarapé Bahia gold deposit, Carajás, Brazil. Mineralium Deposita, 30: 30-38

Notes:

Days 3-4: Field stops in the Tokaj Mts.4

Ferenc Molnár (GeologicalSurvey of Finland)

1. Geological setting of the Tokaj Mts. in the frame of the Neogene-Quaternary Carpathian Volcanic Arc

The volcanic rocks in the Tokaj Mts. are of Upper Miocene (Badenian-Sarmatian, 14-10 Ma) age. The volcanic units of the Tokaj Mts. are parts of the Neogene-Quaternary intermediate- acidic calc-alkaline volcanic range of the Carpathians, and more specifically they belong to the Western Carpathian segment of that range (Fig. 42.). The geodynamic and magmagenetic aspects of the Neogene to Quaternary volcanism of the Carpathians has been discussed by several authors in details (Kaličiak et al., 1989; Lexa et al., 1993; Kaličiak, 1994; Pécskay et al., 1995, 2006; Lexa & Konečný, 1998, Seghedi et al., 2004a, 2004b, 2005). The widely accepted picture involves that the Carpathian volcanic arc has been developed in relation to the southwestward subduction of the Penninic oceanic crust of the Carpathian flysch basins. The subduction was generated by the northeastward escape of two continental lithospheric blocks from the Alpean collision zone (e.g. the ALCAPA and TISZA-DACIA microplates that have been amalgamated during the Upper Cretaceous collision of Africa and Europe in the Western Tethyan realm). Soft collision occurred in the Lower Miocene in the Western Carpathians and mostly in the Upper Miocene in the Eastern Carpathians. Thus – sensu stricto – the formation of the intermediate-acidic volcanic units of the Carpathians can be considered as the result of a syn- to post-collisional volcanism. The combined effect of the transtensional-transpessional tectonism, roll-back of the subducting slab and possible slab break-off processes (from northwest to southeast) resulted in opening of back-arc-like basins (e.g. Pannonian Basin, Transsylvanian Basin and several smaller ones), as well as temporal shift of volcanism from the West to the East and from the Northwest to the Southeast along the Carpathian Volcanic Arc. (Săndulescu, 1988; Szabó et al., 1992; Csontos et al., 1992, 1995; Pécskay et al., 1995; 2006, Vass, 1998; Lexa, 1999). The formation of the intermediate-acidic volcanism lasted from the Lower Miocene until the end of Miocene (from 22 to 9 Ma) in the Western Carpathians, whereas development of the arc along its Eastern part started in the Middle Miocene (15 Ma) and ended in recent times (<0.05 Ma). The last eruptions in the southeastern segment of the Eastern Carpathian volcanic range occurred about 42-11 Ka ago in the Ciomadul Massif. In addition to the intermediate-acidic volcanism, two additional types of volcanism occurred in the Carpathian- Pannonian Region (including both the Carpathian thrust-and-fold arc and the back-arc type basins): 1. Areal type calc-alkaline acidic volcanism. It is represented by extended sheets of dacite- rhyolite tuffs and ignimbrites associated with extrusive domes in the Pannonian and other basins of the Carpathian realm, largely covered by younger sediments. The volcanic activity of this type ranges from Eggenburgian till Lower Sarmatian (21-11 Ma, Pécskay et al., 2006) and the emplacement of volcanic products progressed from the South-west to North-east. From the petrological viewpoint the rocks are of the crustal origin being formed by anatexis owing to overheating of the crust in extensional regime by the mantle diapirism and penetrating mafic magma of astenospheric source. 2. Alkaline to ultra-alkaline volcanism. This type indicates the continuing extension in the back-arc space. It is represented by diatremes, maars, scoria cones and lava flows.

4 This chapter of the guidebook is based mostly on Molnár et al. (1999) and Molnár et al. (2010). 62

Occurrences of these volcanic units are restricted to areas of limited extension as monogenetic volcanic fields such as the Balaton Highland and the Little Plain area in Western Hungary and the Nógrád Volcanic Field in Northern Hungary. Scattered occurrences are also known in the Western Pannonian Basin (Banat region) and in Eastern Carpathians (Persani Mts.), too. This type of volcanism was most widespread in the Carpathian-Pannonian Region between 8 and 0.5 Ma. Its origin is related to various mantle sources whose melting was triggered by small asthenospheric thermal plumes or by decompressional processes.

Fig. 42. Location of the Tokaj Mts. in the Neogene-Quaternary Volcanic Arc of the Carpathians.

From the point of view of hydrothermal processes, metallogenesis, and volcanic raw material deposits, the calc-alkaline intermediate-acidic volcanism has the primary importance in the Carpathians. The famous epithermal gold-silver deposits (e.g. Selmecbánya-Banská Stiavnica, Körmöcbánya-Kremnica, Verespatak-Rosia Montana among many others) were formed in these units and current mineral exploration still finds interesting and promising targets in the old mining fields of the Carpathian volcanic range.

2. General geology and mineral deposits of the Tokaj Mts.

2.1. Geology and volcanism

The Tokaj Mts. comprise the southern part of the Slanské-Tokaj Unit, a volcanic range which is about 150 km long and 15 to 20 km wide in Northeastern Hungary and eastern Slovakia (Fig. 42.). The Tokaj Mts. covers approximately 1200 km2 area and has a moderate topographic relief. The highest peaks are between 700 and 900 m above sea level. The centre of the Tokaj Mts. is located south of Regéc (Fig. 43.) at 21°24’E, 48°19’N. The Middle-Upper Miocene (Badenian-Sarmatian-Pannonian) volcanic-sedimentary sequence of the Tokaj Mts. fills up an approximately 2 km deep, N-S oriented graben-like volcano-tectonic structure and is bordered by the north-northeast trending Hernád Fault and

63 the northwest trending Szamos Fault (Pantó 1968, Gyarmati 1977). The southeast margin of the graben is bordered by the northeast-trending Bodrog Fault. These faults can be related to the major strike-slip, left-lateral Mid-Hungarian Line of the Pannonian Basin. Within the graben, the major faults have north-south orientations or are aligned nearly perpendicular to the Bodrog Fault (Figs. 43. and 44.).

Fig. 43. Geological sketch map of the Tokaj Mts. with the locations of the most significant hydrothermal zones.

The strike-slip movement along the Mid-Hungarian Line caused about 300 km northeastward displacement of crustal units during the Palaeogene – Early Miocene (Fig. 42.). Paleomagnetic data indicate that during the Miocene an approximately 30° counter-clockwise rotation occurred in some regions of the Tokaj Mts. (Balla, 1987; Csontos et al., 1991). The σ1 direction of the regional stress field in the Pannonian area changed from north to the east after the Badenian (Csontos et al., 1991; Csontos et al., 1992). This change in the stress field was caused by the migration of the Carpathian subduction front from west to east. This migration was also accompanied by the change of subduction from oblique in the west to perpendicular in the east. During this tectonic evolution local pull-apart features developed along the regional strike slip faults (Horváth, 1993). Thus the opening of the north-south oriented graben of the Tokaj Mts. may be related to pull-apart extension generated by the large-scale strike slip faulting and the pattern of major faults within the graben reflect the change of the regional stress field from north-south to east-west directions.

Fig. 44. Depth of the pre-Tertiary basement under the Tokaj Mts.

The basement within the eastern part of the graben consists of Precambrian gneiss and mica schist related to an Assyntian metamorphism (962 Ma), Paleozoic porphyroids (metamorphosed during the Caledonian orogenesis at about 450 Ma), sandstone, conglomerate and shale (metamorphosed during the Saalian Orogeny), and Triassic to Jurassic limestone and dolomite. Along the eastern border of the graben these rocks are exposed north of the Szamos Fault, in the Zemplénikum Unit (Fig. 43.); west of the fault, these rocks occur at 960 m depth in the Füzérkajata-2 (Fk-2) borehole (Ilkey-Perlaky & Pentelényi 1978; Fig. 45.). Basement rocks have not been penetrated by deep drilling (-1500 m below sea level) in the western and central part of the graben (Hi-1, Tb-2, Ta-15 and B-3 drillholes; Figs. 43. and 45), but presence of xenoliths in volcanic rocks indicate the existence of Paleozoic shale lithology in the basement (Gyarmati, 1977). The basement along the western part of the depression is at least at 1500-2000 m depth (Zentai, 1991). Volcanism started in the Middle Miocene (Badenian) with accummulation of thick rhyodacitic tuff. K-Ar ages for this unit are between 15.2±1.3 and 13.0 ± 0.6 Ma (Pécskay et al., 1986). Ignimbrite, ash flow, and crystal tuff deposits of this volcanic phase are exposed along the northeastern part of the Tokaj Mts. only (Figs. 43. and 46.). Along the western boundary of the graben, this tuff occurs beneath the approximately 1400m thick younger sedimentary and pyroclastic rocks. The eruptive centres of the Badenian tuff are related to the northwest trending Szamos Fault. The early eruptive phase was followed by subsidence of the graben`s basement and marine transgression from the northeast. The andesitic and dacitic volcanic rocks succeeding this stage were mostly emplaced in submarine environment, resulting in various types of peperitic and brecciated rocks intercalated with shallow marine clays, marls and fine sands. These submarine

65 volcanic accumulations are known from drilling in the central part of the Tokaj Mts. at depths below 800 m (B-3 drilling, Figs. 43. and 45.). Small and shallow subvolcanic andesitic-dacitic intrusions were also associated with the Badenian volcanic activity and these intrusions now crop out in the northeastern part of the Tokaj Mts. or are known from the deepest parts of the Telkibánya (Tb)-2, Füzérkajata (Fk)-2 and Tállya (Ta)-15 drillings (Figs. 43., 45. and 46.).

Fig. 45. Logs of deep drillings completed in the Tokaj Mts. Locations of drillholes are shown on Fig. 43.

At the end of the Badenian, volcanic activity temporarily ceased and uplift of various segments of the basement resulted in the regression of the Badenian sea. Deposition of shallow marine-brackish water clay, marl, sand and reworked volcanic material was restricted to small basins in some parts of the Tokaj Mts. The Sarmatian-Pannonian (Upper Miocene) volcanic phase consisted of successive deposition of acidic and intermediate rocks and basaltic volcanism in the final stage. According to Mátyás (1974), the evolution of the Sarmatian-Pannonian volcanism in different parts of the Tokaj Mts. was initiated by eruptions resulting in rhyolitic tuff sequences up to 600 m thick, accumulating under partly subaqueous, partly subaerial conditions. In the volcanic centres, these tuffs are associated with small rhyolite domes. The products of this volcanic stage crop out in the northern part of Tokaj Mts., as well as in the eastern and southern Tokaj Mts. in an area around Sárospatak, Erdőbénye, Mád and Szerencs (Figs. 43. and 46.). The Lower Sarmatian acidic volcanic stage was followed by or at some places it was contemporaneous with andesitic-dacitic eruptions producing aerially extensive lava flows and local tuff accumulations. Centres of this volcanic stage form caldera-like depressions and circular structures in the vicinity of Telkibánya, Regéc and Mád, as well as volcanic cones in the central parts of the Tokaj Mts (Fig. 46.). The andesitic products of these eruptive events can be divided into two units (Gyarmati, 1977). The thick-bedded andesite lava flows of the lower unit are often intercalated with andesite pyroclastic beds. Most typically, the lava flows of the lower andesite unit have weathered appearance, but propylitic alteration is also typical in areas proximal to

66 volcanic centres. The subvolcanic-intrusive bodies of the lower andesite are located in the caldera-like structures and are characterised by propylitic and adularia-sericite alteration (Telkibánya, Mád, Regéc, Fig. 43.). The caldera-like depressions are also associated with andesitic cones (Telkibánya, Regéc) and extrusive domes of rhyolite and dacite (all centres). The K-Ar ages of rocks from these volcanic centres cluster between 12.5 and 11.5 Ma (Pécskay et al., 1986, Pécskay & Molnár, 2006).

Fig. 46. Volcanological sketch map of the Tokaj Mts.

During the major stages of the Sarmatian andesitic volcanism, some areas of the Tokaj Mts. were still covered by a shallow sea. Therefore, tuffaceous accumulations are often well bedded and intercalated with siliciclastic-pelitic sedimentary rocks deposited in brackish water. Synchronously with or succeeding the accumulation of the lower andesite units, rhyolitic dome-flow centres formed independently between Telkibánya and Sárospatak, as well as in the vicinity of Erdőbénye and Mád (Figs. 43. and 46.). The most typical K-Ar ages of these rocks are between 10 and 12 Ma (Pécskay et al., 1986). The late stage of volcanic activity was characterised by either pyroxene andesite lavas or local dacite-rhyolite products. The andesitic lava flows occurring at higher elevations in the western and central part of the Tokaj Mts. form the so called upper andesite unit, and they are correlated with the emplacement of andesitic dikes (Telkibánya, Mád). Dacitic extrusions and intrusions of late stage volcanism occur sporadically in the area of the Tokaj Mts; the most important

67 occurrence is at town of Tokaj (Figs. 43. and 46.). K-Ar ages of rocks formed during the late, Pannonian stages of volcanic activity are mostly between 9.5 and 11 Ma (Pécskay et al., 1986). The olivine basalt of the final stage of volcanism in the Tokaj Mts. (9.4 ± 0.5 Ma; Pécskay et al., 1986) is not exposed; it is covered by Pannonian and younger sediments and is known only from drillings in the vicinity of Sárospatak along the eastern boundary of the Tokaj Mts. (Figs. 43. and 46.). The Sarmatian-Pannonian volcanic cycle shows a more differentiated petrochemical character compared with the Badenian stage (Gyarmati, 1977). Both volcanic cycles started with acidic products and evolved to an intermediate-basic composition. However, the full rhyolite- rhyodacite-dacite-acidic pyroxene/amphibole andesite-pyroxene andesite series developed only during the Sarmatian-Pannonian cycle. According to Gyarmati (1977), Póka (1988) and Szabó et al. (1992), the petrochemical characteristics of these volcanic rocks have calc-alkaline signature with transitional character between island-arc and active continental-margin magmatic series.

However, the rocks show much diversity, with high-K character (K2O>2.5 wt%) for rhyolite and some dacite and medium K-character (0.75-2.5 wt% K2O) for andesite and dacite (Fig. 47.).

Fig. 47. Geochemistry of volcanic rocks in the Tokaj Mts.

Mátyás (1974) suggested that the temporal variation in the composition of volcanic rocks originated from different levels of secondary magma chambers. Szabó et al. (1992) concluded that the bimodal character of volcanism indicates the existence of different magma chambers for different rock types, or a different degree of contamination in the same magma chamber. The initial 87Sr/86Sr (0.7060-0.7135), 134Nd/144Nd (0.51221-0.51255) and Pb-isotopic ratios presented by Salters et al. (1988) and Downes et al. (1995) indicate a strongly contaminated character of the volcanic rocks in the Tokaj Mts. The isotopic characteristics most probably resulted from the contamination of metasedimentary or acid meta-igneous upper crust in a mantle-derived melt which composition was already modified by the subducted slab.

2.2. Hydrothermal systems, fluid-rock interaction, and mineral deposits in the Tokaj Mts.

Both Badenian and Sarmatian-Pannonian volcanic cycles of the Tokaj Mts. were associated with intense hydrothermal activity. In general, both volcanic cycles generated intermediate to low-sulphidation type epithermal systems in which metallic and non-metallic mineral deposits were formed depending on the palaeodepth of hydrothermal processes, as well as proximal or distal setting in relation to the major conduits of hot fluids. Because of the different depths of erosion, different levels of epithermal environments are exposed in different parts of the region (Fig. 43.). In addition to the mineral deposits of hydrothermal origin, acidic volcanic rocks, especially perlitic rhyolite, and their non-hydrothermal alteration products (e.g. zeolite) have important economic values in the Tokaj Mts. Resources and reserves of non-metallic mineral deposits are listed in the Table 3. Type of industrial Resource Reserve

mineral (1000 tons) (1000 tons)

Kaolin 8200 3100

Illite 1700 800

Bentonite 7700 5400

Silica 10765 5100

Diatomite 5000 1700

Alunite 200 50

Potassium tuff 700 200

Zeolitic tuff 26900 17500

Perlite 30500 15100

Table 3. Non-metallic mineral resources in the Tokaj Mts.

The location of the most strongly mineralised zones is controlled by the major faults and volcanic centres; most commonly their orientation is elongated in a north-south direction. The hydrothermal alteration zones at Telkibánya, Regéc, Komlóska, north of Sárospatak and around Mád are characterised by strong potassium anomalies (Fig. 48.). Chemical analyses of altered rocks from these areas contain above 5-8 wt% K2O (Széky-Fux, 1970; Gyarmati, 1977). These hydrothermal zones with sulphide-poor quartz veins surrounded by adularia-sericite alteration zones (resulting in potassium anomalies) were the sites of the medieval gold and silver mining at Telkibánya and at Rudabányácska (Fig. 43.). In similar zones at Komlóska and around Regéc (Fig. 43.), exploratory pits associated with the medieval mining can also be found. The adularia- sericite alteration zones have been formed at about 200-500m palaeodepth in relation to the palaeogroundwater-table, most typically in the 200-250°C temperature zone of the hydrothermal convection cells driven by small andesitic and dacitic intrusions (Molnár, 1994; Molnár & Zelanka, 1995; Molnár et al., 1999). Shallower zones are characterised by intense argillic alteration: an unique example of this kind of hydrothermal clay deposits is located near Füzérradvány (north of the location of Fk-2 drillhole, along the Hungary-Slovakian state border on Fig. 43.) where a high quality illite deposit occur along the main conduits of the palaeohydrothermal fluid flow. In the last 100 years, the alunite-kaolinite alteration zones (the

69 very shallow steam-heated acid alteration part of epithermal systems) in the Mád area and north of Sárospatak, as well as the strongly argillised (bentonitic) host rocks to veins at Komlóska (Fig. 43.), have been the primary targets of clay exploration and exploitation. Recognition of shallow erosion depths of several hydrothermal fields in the Tokaj Mts. has also been generated exploration programs for epithermal gold deposits in the past 15 years. These programs were partly focussed on the known medieval gold-silver mining fields but also recognised new gold bearing zones in the area of Mád and Füzérradvány. However, further exploration is needed to determine mineable resources. The small basins and their lacustrine environments formed at the late stages of volcanism in several parts of the Tokaj Mts. host to diatomite, bentonite, kaolinite and silica deposits, especially in the areas around the village of Mád (Figs. 43. and 44.). These lacustrine environments were also fed by hot springs and accumulation of raw materials can be interpreted as a result of combined distal hydrothermal and local sedimentary processes.

Fig. 48. Aero-gamma map of the Tokaj Mts. recalculated to the potassium contents of rocks. The high-potassium zones mark intense adularia-sericite alteration.

A generalised model for the hydrothermal systems of the Tokaj Mts. with the occurrences of various mineral deposits is shown on the Figure 49. This model was established on the basis of detailed mineralogical, petrographycal, geochemical, fluid inclusion and K-Ar studies (Molnár et al., 1999). In various hydrothermal fields of the Tokaj Mts., the erosion level is different and thus they can be placed into different palaeodepth position in the generalised model. Detailed K-Ar studies on hydrothermal minerals proved that the age of hydrothermal mineralisation in the more deeply eroded zones (e.g. adularia-sericite alteration zones with Au-Ag accumulations) in the northern part of the Tokaj Mts. are slightly older (12-13 Ma) than the less eroded hydrothermal fields (e.g. steam-heated alteration zones with alunite and kaolinite) in the southern part of the Tokaj Mts. (10-11 Ma; Pécskay & Molnár, 2002; Fig. 50.).

Fig. 49. Model for the shallow zones of epithermal systems in the Tokaj Mts. The model is based on the observed vertical and horizontal zoning of alteration, vein mineralogy and fluid inclusion data in the hydrothermal centres with different depths of erosion in the hydrothermal centres with different depths of erosion in the Tokaj Mts.

3. Field stops in the Tokaj Mts.

3.1. Field stop 1: Fresh ignimbrite with gas-emanation channels: quarry at Vizsoly village

The quarry is located along the western side of the Tokaj Mts., at the southern boundary of the village of Vizsoly (Fig. 43.). The quarry is a protected site of geological interest as it is the type locality of the Vizsoly Rhyolite Tuff Formation. The pumiceous ignimbrite, which is exposed up to 10 metres thickness along the walls of the quarry, has been accumulated during the last phase of the volcanism in the Tokaj Mts. The radiometric age data indicate Sarmatian-Pannonian (11 Ma) age. The most probable source of the pyroclastic flow is a volcanic eruption centre near the village of Boldogkőváralja, approx. 2 km distance to the southeast. The fresh, weakly welded ignimbrite contains pumice fragments with up to 20 cm diametre and variable sized, but usually not larger than 2-3 cm angular lithoclasts, as well as some crystals or fragments of crystals of biotite and feldspar in a fine grained volcanic ash matrix. The volume of matrix dominates over the fragments. Lithoclasts mostly consist of rhyolite and intermediate volcanic rock fragments and minor amounts of laminated-layered fragments of sedimentary rocks. Layering or internal sorting of pumice fragments and lithoclasts cannot be observed.

Fig. 50. K-Ar ages for volcanic rocks and hydrothermal minerals in hydrothermal centres of the Tokaj Mts.

A peculiar characteristics of the ignimbrite at Vizsoly is the occurrence of vertical gas- emanation channels. Formation of these channels can be related to the deposition of ignimbrite in a wet, most probably lacustrine/swamp environment. The fast accumulation of the still hot

72 pyroclastic material on the top of wet, unconcolidted sediments caused evaporation of water. Steam formed by this way and volcanic gases enclosed in the ignimbrite emanated along numerous channels. The fast emanation of volatiles blowed out the volcanic ash matrix along those channels but it was unable to move larger fragments. Due to the episodic eruption of volatiles, significant alteration of the tuff have not taken place along those channels. The vertical position of the gas-emanation channels suggest that the pyroclastic flow deposit is in its original position. The weakly welded pumiceous ignibrite is easy to cut and grind. Due to the relativey high porosity its density if low and heat-insulator capacity is high. Therefore it was a preferred building material in the past centuries used by the local habitant. Nowadays, this kind of rock is not mined at Vizsoly. However, there are huge deposits of ignimbrite in the southern and southeastern part of the Tokaj Mts. where excavation and processing produces various valuable materials for the modern building industry.

3.2. Field stop 2: Steam-heated alteration zone in ignimbrite: Király Hill north of the village of Mád

The exploitation of kaolinite, bentonite, zeolite and pure silica deposits started before World War II in the area around Mád, in the southern part of the Tokaj Mts (Figs. 43. and 51.). State owned mining was most active between the 1950s and 1970s; nowadays private companies exploit these raw materials. The formation of clay deposits is related to steam-heated alteration zones of epithermal systems near Király, Bomboly, and Diós Hills (Fig. 51.). Diagenetic alteration of glassy rhyolitic tuffs resulted in formation of clinoptilolite and mordenite deposits. Bentonite together with kaolinite and pure silica which were deposited in local fresh-water basins fed by palaeo-hot springs (distal environments of hydrothermal systems) also form important deposits 1.5 km west of Mád (Fig. 43.).

Fig. 51. Geological sketch map of the hydrothermal zone northeast of Mád.

The products of the Badenian volcanic cycle of the Tokaj Mts. are not exposed in the area around Mád. The 1200 m deep Tállya-15 (Ta-15) drillhole (see its location on Fig. 43. and log on Fig. 45.) reached the Badenian rocks (14.2 Ma K-Ar age; Pécskay & Molnár, 2002) at 900 m depth below the present surface. The Mád-23 drillhole in the centre of the mineralised area northeast of Mád (Fig. 51.) found only Sarmatian-Pannonian volcanic and sedimentary rocks (11.5-12.2 Ma K-Ar age, Pécskay et al., 1986) to a depth of 712 m. According to Zelenka (1964) and Mátyás (1974), the major part of the Sarmatian-Pannonian volcanic sequence is composed of five rhyolitic tuff units around Mád (Fig. 52.). Each unit reflects the variation of local, subaerial or subaqueous conditions of accumulation and they are locally intercalated with shallow marine clay and marl beds. The pyroclastic units are predominantly pumiceous glass tuff and lapilli bearing pumiceous tuff with subordinate crystal tuff as well as ignimbrite. The explosive tuff accumulations are associated with extrusive domes and pumiceous lava flows.

Fig. 52. Lithostratigraphy and major volcanic phases in the southern part of the Tokaj Mts.

In the mineralised area on the Király Hill and its surroundings, northeast of Mád, the rhyolitic tuff and associated domes of the fourth explosive phase can be found on the surface (Fig. 51.). This phase was followed by andesitic volcanic activity producing lava, tuff and agglomerate beds deposited in shallow water and terrestrial environments. Deposition of the pumiceous rhyolite glass tuff (ignimbrite) of the fifth volcanic phase together with subsequent rhyolite domes and dacitic eruptions (9.8-10.8 Ma K-Ar age; Pécskay et al., 1986) represents the last stages of

74 volcanism in this area. Rhyolite extrusive domes now form the Király Hill, Bomboly Hill and Diós Hill (Fig 51.). The major faults of the area have NE-SW, N-S and E-W directions. The N-S and NE-SW trending faults were active during the hydrothermal processes, indicated by the predominant strike of siliceous veins (Fig. 51.). The E-W trending faults were also active after the volcanic-hydrothermal stages, resulting in block tectonism of the volcanic and sedimentary sequences. The mineralised zones and quartz veins of this area occur in various host rocks including rhyolite and rhyolitic tuff as well as andesitic rocks. K-Ar ages of alunite from the rhyolite of Mogyorós Hill (Fig. 51.) are 11.7 Ma, whereas alunite from the ignimbrite of Király Hill has an age of 10.9 Ma (Pécskay & Molnár, 2000). These data suggest protracted hydrothermal activity in this area and are in the range of K-Ar ages for volcanic rocks. A horizontal zonation of the steam-heated alteration at shallow depths (along and above the palaeogroundwater-table; see Fig. 49.) of an epithermal system is exposed in the quarry of the Király Hill (Fig. 51.). Here the most intensively silicified zones in the the fourth and fifth rhyolitic tuff (ignimbrite) units formed along a tectonic contact with the fifth stage rhyolite (Fig. 53). The silicification is most intensively developed in the porous rhyolitic tuff; rhyolite is altered only just along the fault. The silicified zone of the fifth tuff unit exposed in the quarry laterally changes into a kaolinite zone with decrease in the SiO2 content (from about 95-98 wt% to about 80 wt%) and increase of the Al2O3 content (from 2-3 wt% to 10-11 wt%). The average sulfate content increases to about 3 wt%, but locally can be as high as 15-20 wt% (chemical data from

Mátyás, 1966 and Varjú, 1974). Sulfate-rich zones also have 3-5 wt% K2O, suggesting that the enrichment of sulfur is associated with the presence of alunite. Nemecz (1973) reported dickite in addition to kaolinite in the kaolinite-alunite zone. The kaolinite bearing zone is fringed by illite- montmorillonite alteration. According to Nemecz (1973), the illite-montmorillonite has an interlayered structure with ordered pattern of illite and montmorillonite layers (this type of clay mineral was described as ‘allevardite’ and ‘rectorite’ in the older literature). The outermost part of this alteration zone contains montmorillonite only. In the strongly silicified zone, the original texture of ignimbrite is totally masked by silicification. As the intensity of silicification decreases towards the kaolinite-alunite zone, the rock becomes highly porous due to hydrothermal leaching. The resulting cavities are mostly found in places of former pumice fragments and formed the preferred sites for precipitation of alunite. Alunite is mostly pure K-alunite (minor Na-content may occur in some growth zones; Bajnóczi et al., 2002) and forms aggregates of fine-grained individual crystals (0.1-1 mm) with rhombohedral and platy habit. The 10.8 Ma K/Ar age reported in Pécskay & Molnár (2000) was determined on cavity-filling alunite from this locality. Minor quartz, chalcedony and opal is associated with alunite. Very fine alunite is also present in the kaolinitic-siliceous groundmass of the rock. In the illite-montmorillonite zone, the color of the tuff changes from white to grey/greyish-green. The cavities formed by acidic leaching are less abundant and filled up by illite-montmorillonite or montmorillonite. No sulphide minerals were found in the alteration zones. The major Fe-bearing mineral is the hematite and other Fe-oxyhydroxides. The presence of hematitic patches and disseminations are typical in the marginal parts of the kaolinite-rich alteration zone in other similar shallow epithermal environments of the Tokaj Mts. (Mátyás, 1973). Steam-heated alteration with alunite forms in a near-surface portion of a low-sulphidation type epithermal environment, where H2S exsolved through deep boiling of hydrothermal fluids is absorbed into steam-heated water and undergoes oxidation to form H2SO4 (Schoen et al.,

1974; Henley & Ellis, 1983). At the paleogroundwater table, massive silicification occurs in a porous unit, probably as a result of mixing of three different fluids: (1) acidic fluids descending and mobilizing silica from the acid-leached zone; (2) ascending near-neutral, also silica rich hydrothermal fluids and (3) cooler aquifer fluids (Schoen et al., 1974; Sillitoe, 1993). The horizontal zonation with less acidic character of alteration moving apart from the silicified and kaolinite-alunite bearing zones suggests lateral flow of fluids along the palaeowater table at Király Hill.

Fig. 53. Mineralogical zoning in the steam-heated alteration zone exposed on the Király Hill, near the village of Mád.

Several other steam-heated alteration zones also occur in the vicinity of the Király Hill (Fig. 51.). The widespread occurrences of these zones suggest shallow erosion of a large scale epithermal system and thus the area has potential for epithermal gold mineralisation at depth. The vertical zoning of hydrothermal alteration is well represented in the Mád-23 drillhole completed near to the Király Hill (Figs. 51. and 54.). The upper 50 m of the sequence is characterised by kaolinite-alunite alteration with silicification in andesite and andesite tuff. This is underlain by a zone of K-feldspar-chlorite-carbonate (-smectite-illite) alteration in the subvolcanic andesite. The deeper zones (116-240 m) of this intrusive body are characterised by epidote-carbonate-pyrite-chlorite alteration. Below 240 m depth K-feldspar is present again in the alteration assemblage. The vertical zonation of the hydrothermal alteration within and above the shallow subvolcanic intrusion corresponds to a typical low-sulphidation type epithermal alteration pattern with propylitic zones at depth, adularia-sericite(-chlorite- carbonate) alteration at intermediate depths and steam-heated (kaolinite-alunite) alteration in the shallow portion of the system (Silberman & Berger, 1985).

Fig. 54. Lithology and alteration mineralogy in the Mád-23 drillhole. Location of this drillhole is shown on Fig. 51.

3.3. Field stop 3: Silica and clay deposits in a hot-springs fed lacustrine basin west of Mád

The sedimentary sequences of the lacustrine basin west of Mád crop out over an approximately 8 km2 area (Figs. 43. and 55.). The original size of the basin is not known because it is bordered by normal faults to the west and south. The basin history was reconstructed by Mátyás (1966) from observations of more than 460 shallow drillholes and exposures in open pits.

Fig. 55. Geological sketch map, lithostratigraphy and mineralogy, as well as a section across the siliceous beds deposited in the lacustrine basin west of the village of Mád.

The limnic sequence of the basin is underlain by a pyroxene andesite lava flow related to the intermediate volcanic activity between the fourth and fifth rhyolitic tuff unit known of the area (Fig. 52.). The upper parts of the lava beds contain hyaloclastite breccia indicating subaqueous accumulation condition. In the lacustrine sedimentary sequence, three major consolidated siliceous layers occur (Fig. 55.). The mineral composition of massive silica beds is quartz according to the XRPD studies. There also are unconsolidated beds (‘microsil’; Fig. 55.) which are composed of uncemented quartz plates of 5-20 µm size and very little argillic material. The ‘microsil’ contains

93 wt% SiO2 while the silica content of the massive siliceous layers is above 98 wt%. The siliceous layers show bedding in which the individual layers have very variable thickness from a few centimetres up to 1-2 metres. Some of these layers show fine internal lamination, however, chaotic-brecciated structures are also common probably due to the slumping of the unconsolidated and semi-consolidated silica mud at the time of deposition. The thick siliceous layers also contain interbedded argillic material and rhyolitic tuff whithout silicification. This field evidence suggests that the silica was mostly precipitated as very fine mud from the hydrothermal solutions and was not the product of superimposing silicification of sedimentary and pyroclastic units.

The color of individual beds varies from white to grey, green, brown, red and black. The variation of the color from layer to layer is related to the presence of minor mineral particles and organic material. Red and brown layers contain fine Fe-oxides, while black layers contain framboids of pyrite and marcasite. These sulphides occasionally correlate with the distribution of organic material. The siliceous layers often contain petrified wood and fossilised plants (branches and leaves) indicative of a shallow water or swamp environment. Typical plant fossils comprise Gliptostrobus, Osmunda, Taxodium, Populus and Salix species. Prints of Quercus mediterranea, Quercus kubinyii and Acer decipiens transported into the basin from the surrounding areas. The occurrence of relatively low temperature conditions during the accumulation of the silicieous beds is reflected by the presence of tortoise and fish fossils. The siliceous beds are generally characterised by high concentrations of Sb (up to 400- 600 ppm) and lesser As and Hg, as well as some Ag (Vető, 1969; Szakáll, 1990; 1991). Stibnite, stibiconite and realgar sporadically occur in the siliceous beds. The siliceous layers appear in a cyclic sedimentary series in which the periodicity of sedimentation resulted from the relative intensity of accumulation of terrigenous, pyroclastic and chemically precipitated materials (Fig. 55.). The occurrences of tuffite, tuff sand and local conglomerate units indicate the periods of the predominantly pyroclastic and terrigenous source of basin filling. These sediments predominate at the base of sedimentary cycles. Angular pumice fragments are present in each sequence suggesting that basin filling correlates with the fifth eruptive phase in the area. The grain size of sediments gradually decreases, not only horizontally (to the west) but also vertically in each cycle. Thus, the upper parts of sedimentary cycles are characterised by pelitic sediments which also show intensive kaolinitic and bentonitic alteration near to the hydrothermal vents. The localities of feeder zones within the basin (Fig. 55.) are marked by this horizontal zonation of clay minerals (kaolinite close to the feeder zones fringed by a montmorillonite halo) as well as by the thickness of silica layers. Close to the centres, the thickness of siliceous beds is up to 10-12 metres, while on the margins they totally pinch out. In addition to the feeder zones within the basin, Mátyás (1979) also outlined subaerial hydrothermal centres around the depression. These centres now form small hills with strongly silicified and brecciated rocks. The formation of thick silica mud accumulations in the lacustrine basin near Mád is interpreted to be distal manifestation of hydrothermal systems of the region. Fluids discharging from the paleo-hot springs transported silica into the local fresh water basin and mixing with cooler and acidic fluids of the organic material rich shallow water/swamp enviroment caused precipitation of siliceous material. In the period of ceased hydrothermal activity mostly terrigeneous sediments and pyroclastic material accumulated in the basin and subaqueous alteration of glassy tuffaceous material led to formation of bentonite and kaolinite beds among the silica layers.

3.4. Field stop 4: Perlite pit, Gyöngykö Hill at the village of Pálháza

In the Tokaj Mts., rhyolite dome-flow complexes are widespread in two areas: in the Mád- Szerencs area and in the northern part of the mountains, between Telkibánya and Pálháza (Fig. 43.) The quarry in the Gyöngykő Hill at Pálháza village is located in the northern rhyolite field of the Tokaj Mts. (“gyöngykő” means “pearl-rock” in Hungarian). The quarry on the northern slope of the Gyöngykő Hill exposes a perlitic part of a rhyolite intrusive and extrusive dome-flow complex which penetrated the lower Sarmatian shallow marine clay and marl units (Fig. 56.). Perlite occurs as marginal facies of lava flows on the slopes

79 of a cryptodome with dip angle of bedding varying from subhorizontal to fairly steep angles. The size, the geometry and the contact relationship between the host sediments and the coherent igneous bodies indicate a multiple injection and pulsatory magma supply that formed the cryptodome and dome complex exposed here.

Fig. 56. Section across the Gyöngykő Hill with location of the rhyolite-perlite quarry at the village of Pálháza.

In general, petrochemistry of the well differentiated Sarmatian-Pannonian volcanic series show medium to high calk-alkaline character with up to 5 wt % potassium content for rhyolites in the Tokaj Mts. (Fig. 47.). SiO2 content of rhyolite from the Tokaj Mts. ranges between 70 wt% and 77.5 wt%. No significant difference can be detected in the major element and also the trace element concentrations of perlite and fresh rhyolite in the Tokaj Mts., except that the H2O and the Cl contents of perlite are remarkably higher (H2O from 3.9 to 4.7 wt%; Cl = from 676 to

740 μg/g) then those of the fresh rhyolite (H2O> 1.4 wt% and Cl> 460 μg/g; geochemical date are from Németh et al., 2008). Boron-contents of rhyolite and perlite are between 20 and 68 μg/g; this range of boron concentrations is typical for subduction related felsic volcanic rocks. K-Ar ages are in accordance with biostratigraphic data which shows that perlite at Pálháza is slightly older than perlite from the adjacent areas in the Carpathians (13.6 Ma at Pálháza, 12.5- 11.6 Ma in other areas e.g. Oas-Gutaii Mts. in Romania, Transcarpathia in Ukraine and Slansky Vrchy in Slovakia according to Németh et al., 2008). The quarry reveals a 3D view of an at least 200 m thick succession of complex rhyolitic cryptodomes and endogeneous lava domes, partially intruded into wet, unconsolidated marine sediments (Fig. 57.). The individual coherent magmatic bodies are slightly concave upward and they are up to 50 m wide and 25 m thick. The columnar jointing is well developed in the inner core of the individual lensoid coherent lava facies. Many of the margins of the coherent rhyolite are glassy and strongly perlitic in texture. Columnar jointing forms a slightly radial joint pattern and platy fracture zones perpendicular to the main jointing. The more intact rhyolitic coherent units are commonly flow banded with alternating cm-wide black and grey zones. The coherent bodies have only occasional vesicular texture, and the vesicles exclusively located in the outer margin of the individual coherent igneous bodies. The coherent units are very glassy in general, and porphyritic, aphanitic in texture. The interior of the coherent silicic magmatic bodies are commonly false vitriclastic in texture with small dm-size relicts of original perlitic, black glassy zones. Along the dome margins, thin

80 darker colored, variably devitrified obsidian perlite can be identified that are commonly flow banded and shows also false vitriclastic texture.

Fig. 57. Development of the rhyolite-perlite dome-flow complex at Pálháza. 1. Cryptodomes intruding unconsolidated wet sediments; 2. Stacking of rhyolite domes under shallow marine conditions.

The contact relationship with the underlying marine sediments records shallow marine (from few tens to few hundred metres depth) depositional environment and suggest subaqueous emplacement of the igneous masses. Fragmented volcanic rocks form a complex architecture enclosing most of the coherent rhyolitic rock units in the quarry. The matrix supported, angular, glassy commonly flow banded rhyolitic clast rich, massive volcaniclastic breccias are interpreted to be hyaloclastite associated with the emplacement of the silicic igneous units at Pálháza. The general increase of jig-saw fit hyaloclastitic breccia near the rhyolitic, rhyodacitic coherent zones indicates that the

81 hyaloclastite is in situ and no significant remobilisation of clasts took place through gravity flows. Besides fragmented volcanoclastic units the coherent rhyolite domes are also surrounded by few metres thick breccia zones along the contacts with shallow marine sediments which were still unconsolidated and water-rich at the time of volcanism. Proportion of the marine sediments increases dramatically, especially on the base of the exposed volcanic successions. These reaction rims of the rhyolitic lava and the soft in-situ marine sediments are considered as a peperitic facies of the volcanic structure. The identification of peperitic margin along the Pálháza rhyolite complex attest its predominantly shallow subsurface intrusive origin, where rhyolitic magma initially intruded unconsolidated and wet fine grained sediments. Unlike it is common in the case of mafic lavas peperites here are not much fluidal (globular) but display a rather blocky nature which is interpreted to be the result of the progressive cooling of the magma during intrusion, and also from the breakdown of fluidisation when the limited supply of fine glassy matrix in the host hyaloclastite was exhausted similarly to the processes well described in mafic intrusions into coarse grained sediments. Increase of magma viscosity during cooling (and overall lower intrusion temperature compared to mafic magmas), combined with multiple amplacement of cryptodomes were inferred to be important in the production of the blocky peperite. Muddy sedimentary dikes of various thicknesses also penetrate into the jig-saw breccia zones of the coherent rhyolitic bodies. What makes perlite an important raw material is its water content. Compared to other volcanic felsic glassy materal (pumice, obsidian), it contains a relatively large amount of water even up to 2-7 wt%. Globally, most of the known perlite occurrences are in young volcanic areas because of easy devitrification of the material. Average composition of perlite used for industrial purposes is:

SiO2: 73%, Al2O3: 13%, Alcalines: 7%, Water: 2-7% FeO: 1-2%, MgO: 1% CaO: 1.5% TiO2: <0.3% P2O5: <0.03% Perlite after fine-crushing is heated in furnaces. During this process water content of the rock evaporates while the rock expands resulting in a very porous material with tube-like microscopic pore spaces, resistant to moisture, heat, pests and weathering. Volume increase during the processing can be as high as 1:20. Depending on the properties and quality of the material, expanded perlite can be widely used in different applications. Thermal and noise insulations, filters, absorbents are the main fields of use of high quality perlite. Large amount is used as admixture to concrete and other constructional material. Agricultural use is also significant. Although perlite also occurs in other areas of the Tokaj Mts., as well as in Neogene clardera stucture of the Mátra mountains, exploitation takes place at Pálháza only. Mining of perlite near Pálháza began in 1957 and continued without disruption up until now. The quarry and a newly opened pit to the south provides so large production that Hungary is listed among the 5 largest perlite producers of the World. Today perlite from Pálháza adds 6% to the world’s and 30% of Europe’s perlite consumption.

3.5. Field stop 5: Adularia-sericite alteration in pumiceous rhyolite/crystal tuff: historic gold mining site at the village of Rudabányácska

The history of mining of gold in the Tokaj Mts. goes back to the Medieval Ages. The first documentation of gold mining at Telkibánya (Fig. 43.) is known from 1341: an order by the royal court described the boundaries of the mining field. However, gold mining has been gradually deceased after the 17th century due to the collapse of the old mine caused by an earthquake. The last report on the delivery of a small amount of gold exploited from the old mines at

Telkibánya to the mint at Körmöcbánya (nowadays Kremnica, located in Slovakia) is from the second part of the 19th century. Mining of gold at Rudabanyácska was probably started at the same time than at Telkibánya: in 1347 the mines at Rudabányácska were affixed to the mining district of Telkibánya. However, the production of gold was probably much smaller as in addition to the numerous pits, only one major adit was constructed at Rudabányácska. Mining at this locality was already abandoned during the 15-16th centuries. During the past decades, renewed interest in epithermal gold deposits has generated some exploration programs in the Tokaj Mts., but currently no one of the potentially gold bearing areas is under exploration. The mineralised area near the village of Rudabányácska is hosted by felsic pyroclastic deposits (ignimbrite, welded ignimbrite, crystal tuff) of Badenian age (Fig. 58.). The cumulative thickness of these units are only a few hundred metres and they cover Triassic-Jurassic carbonate basement rocks that are underlain by metamorphic rocks of Precambrian to Palaozoic age (Gyarmati & Pentelényi, 1973; Figs. 43. and 44.). K-Ar ages for biotite and sanidine from the pyroclastic deposits are between 15.2 and 13.1 Ma (Pécskay et al., 1986; Pécskay & Molnár, 2002). The eruption centres for ignimbrite are probably along the Szamos Line (Gyarmati, 1977); however, there also are small rhyolite domes in the vicinity of Rudabányácska that have similar ages (13.0 Ma). The pyroclastic rocks were intruded by dacitic bodies (Száva Hill; Fig. 58.) which suffered hydrothermal alteration. Subsequent andesitic dacitic-andesitic lava flows covering ignimbrites are fresh. K-Ar data for whole-rock and amphibole samples suggest a Sarmatian age for intermediate lava flows of this area (12.4 - 12.2 Ma, Pécskay et al., 1987; Pécskay & Molnár, 2002).

Fig. 58. Geological sketch map of the medieval gold mining area near the village of Rudabányácska.

The hydrothermally altered rocks of the area occur in a NW-SE oriented zone nearly parallel to the Szamos Line (Figs. 43. and 58.). The most intense adularia-sericite-pyrite-hematite alteration affected the felsic pyroclastic units of the Bányi Hill and the dacite intrusion of Száva

Hill (Fig. 58.). The zone with adularia-bearing alteration assemblage is surrounded by regional propylitised rocks (Varga-Máthé, 1961). Rocks with adularia-sericite alteration host a siliceous- pyritic disseminated-stockwork gold deposit that was exploited during medieval time. According to fluid inclusion data, the deeper zones of gold mineralisation (Bányi Hill) were formed between 230-300°C during boiling and dilution of fluids by meteoric water (Fig. 59.). Fluid inclusion data suggest that the adularia-sericite alteration and deposition of gold took place at about 300 m minimum depth below the paleogroundwater table (Molnár, 1994).

Fig. 59. Interpretation of hydrothermal alteration and fluid inclusion data from the point of view of deposition of gold in the hydrothermal system at Rudabányácska.

K-Ar data for intrusive dacite of Száva Hill that shows intense K-metasomatic (K-feldspar- sericite) alteration is 13.2 Ma (Fig. 50.). K-Ar age for adularia from the stockwork deposit of Bányi Hill is 13.0 Ma (Pécskay & Molnár, 2002). These data suggest that hydrothermal mineralisation of the area near Rudabányácska took place at the end of the Badenian rhyolitic-dacitic volcanic stage and may be related to emplacement of shallow intrusions.

3.6. Field stop 6: Strongly silicified steam-heated alteration zone on the Botkö Hill, northwest of the town of Sárospatak

The palaeohydrothermal field at Sárospatak is located along the eastern boundary of the Tokaj Mts. and it covers 9 km2 area. The centres of hydrothermal activity are now exposed as low hills (150 to 300 m elevation) above the surrounding area (130-140 m elevation). The most intensively altered and mineralised zones at Király Hill and Megyer Hill form a semicircular range around a local basin opening to the south (Fig. 60.). The area has long history of raw material exploitation. The hydrothermally altered rocks provided excellent building materials for the surrounding settlements. The intensively silicified pyroclastic rocks were used as millstones and the first documentation of this industry at Sárospatak is known from the 16th century. The kaolin deposits has been explored and exploited during this century; however, none of these mineral occurrences have recently been worked. The basement of the Tokaj Mtns. around Sárospatak is elevated relative to that of the central and western zones (Fig. 44.). Geophysical data indicate less than 500 m depth to basement (Zentai, 1991) The Sárospatak-5 drillhole found Triassic limestone and dolomite at 225 m depth in the central part of the hydrothermal field (Figs. 44. and 60.).

The basement is unconformably covered by the acidic pyroclastic rocks of the first, Badenian explosive phase of volcanic activity (Fig. 60.). The tuff conglomerate beds in the oldest rhyodacite tuff unit contain clasts of the carbonate and metamorphic rocks, indicating that some areas of the basement was still exposed to erosion during this time (Gyarmati & Pentelényi, 1973). In the upper horizon of this unit, the tuffaceous material is predominant and the Lower Badenian age for the subaqueous accumulation is indicated by the macrofauna.

Fig. 60. Geology and hydrothermal alteration zones, as well as an E-W oriented profile across the hydrothermal zone near Sárospatak.

Towards the end of the Badenian volcanism was evolved to a rhyolitic composition, and the related tuff is characterised by alteration (clinoptilolite, heulandite, cristobalite, chlorite, montmorillonite), and occurrence of redeposited tuff, tuffite and argillic beds reflecting subaqueous accumulation. Some parts of the area east of the major north trending fault (Fig. 54.), was uplifted. In these subaerial zones, terrestrial tuff accumulations can be found. The Mollusca fauna of the subaqueous units indicates an Upper Badenian age (Hoffer, 1925; Frits, 1960; Csepreghy-Meznerics, 1966). K/Ar ages of the terrestrial unaltrered tuff are between 14.6- 12.9 Ma (Pécskay et al., 1986). At the end of Badenian and Early Sarmatian a transgression occurred and rhyolite tuff and tuffite accumulation and reworking as well as shallow-water clay and sand sedimentation took place in the southern foreland of Király Hill and Megyer Hill (Fig. 60.). According to Mátyás (1979) the recent morphology and distribution of tuff and sedimentary units reflects the Early Sarmatian paleogeography of the area. The products of the Middle-Upper Sarmatian intermediate-acidic volcanism mostly occuring west and north of the major fault bordering this area (Fig. 54.). The Sarmatian andesitic lava flows are not faulted in its zone; therefore Gyarmati and Pentelényi (1973) concluded that the

85 subsidence of the basement west of the fault occurred during Late Badenian. The fracture system parallel to the major fault is characterised by an elevated thermal gradient (Frits, 1964). In the southern continuation of the hydrothermal zone between the Megyer Hill-Botkő (Fig. 54.), Ca-Mg-sulfate-carbonate thermal water at 48°C was found in drillholes reaching fractured Triassic carbonates in the basement. The hydrothermal alteration of this area occurs in two, north-south zones parallel to the major fault to the west (Fig. 60.). The host rocks are both the Upper Badenian rhyolitic tuff and Lower Sarmatian reworked rhyolitic tuff units, suggesting that hydrothermal activity took place later than the Lower Sarmatian. The epithermal alteration has two major types in this area: - siliceous horizons with surrounding and underlying alunite and kaolinite alteration zones (Botkö, Király Hill, Megyer Hill and Cinegés) - adularia-sericite alteration under the siliceous and kaolinite-alunite alteration zones (Király Hill) In the hydrothermal zones, the local upflows are marked by strongly silicified cap-like rock bodies on the top of Király Hill, Megyer Hill, Cinegés and Botkő Hill (Fig. 60.) The concentration of silica in these centres is > 95 wt% SiO2 and the silicification totally masks the original rock texture. The siliceous material is mostly quartz at Megyer Hill, Király Hill and Botkő; however, it is predominantly opal at Cinegés (Ilkey-Perlaki, 1989). The silicified centres are also characterised by intense hydrothermal brecciation at Király Hill and Botkő. The size of the angular breccia clasts is highly variable between 0.1 and 10 cm. In the open spaces of these breccias, as well in the groundmass of the silicified rock, barite and hydrothermal quartz is ubiquitous. Cinnabar is also present as dust-like encrustations in the vugs or dendritic dissemination in the silicified rock. Rare pyrite, and abundant hematite also associates with these minerals (Kulcsár & Barta, 1969; Molnár, 1993). The strongly silicified rocks are characterised by geochemical anomalies of Sb (up to 370 ppm), Hg (up to 41 700 ppm), As (up to 210 ppm), Tl (up to 18 ppm), as well as Ba and Sr (Kulcsár & Barta, 1969; Gyarmati & Pentelényi, 1973; Szakáll, 1991). The central silicified zones are surrounded and underlain by kaolinite-alunite, kaolinite, illite and montmorillonite alteration (Fig. 61.). In the kaolinite-alunite and kaolinite zones the degree of silicification decreases apart from the brecciated siliceous bodies and opal/cristobalite is present instead of the quartz. The presence of hydrohematite is also typical for the kaolinite- bearing alteration zones. Alunite occur as fine-grained (nx10-nx100m) replacements of the groundmass and pumice fragments of the tuff, as well as later fracture-filling aggregates with slightly coarser crystals (Mátyás et al; 1971; Ilkey-Perlaky, 1989). At the lower elevation of the Király Hill (Figs. 60. and 61.) adularia-sericite alteration is also present (Kulcsár & Barta, 1969; Molnár, 1993). In these zones the orginal sanidine shows sericite (illite) alteration. Fresh adularia is present in the mosaic-textured quartz of the silicified groundmass. In the adularia-bearing levels of Király Hill, the common occurrence of liquid-rich and gas-rich fluid inclusions plus solid adularia inclusions in quartz indicate boiling. Vapor-rich inclusions are also common in barite cementing the hydrothermal breccia at higher elevations (Molnár, 1988 and 1993). Homogenisation temperatures are between 200 and 250C from the adularia-bearing zones, and are mostly between 150 and 200C from the breccia of strongly silicified zones (Fig. 61.). The salinities for these inclusions are between 0.7 and 4.2 NaCl equiv.wt%. Crush-leach and ICP-decrepitation analyses indicate Na/K ratios from 0.95 to 1.16; the Ca and Mg content of

86 fluids is low (Molnár, 1993). Therefore sea-water or connate origin of hydrothermal fluids can be excluded despite considerations of the geology of the area (Fig. 60.).

Fig. 61. N-S oriented section with zoning of alteration and fluid inclusion data in the hydrothermal zone at Sárospatak.

The temperature differences between the different types of alteration are related to their different levels within the paleohydrothermal system (Fig. 61.). Temperature data correspond to a minimum 150 m paleodepth for the upper, silicified horizon, and a minimum 350 m paleodepth for the adularia-sericite alteration of the Király Hill (Fig. 61.). The location of samples from different levels along the same boiling curve indicates a genetic link between the deeper adularia-sericite alteration and the silicification and brecciation at the shallow levels. The silicified horizon on the top of the Király Hill may correspond to a preserved remnant of a thick strata-bound zone of massive silicification formed beneath an already eroded steam-heated argillic alteration zone according to a model suggested by Sillitoe (1993). The sealing of upflow zones by massive silicification between the boiling level and the uppermost steam-heated alteration zone may lead to a development of a high-pressure gas-rich fluid regime beneath the level of silicification resulting hydrothermal brecciation (Hedenquist & Henley, 1985b). The appearance of alunite-kaolinite alteration beneath the silicified zone of the Király Hill can be considered as an overprinting steam-heated alteration caused by the drop of the paleogroundwater table. However, the presence of at least two adularia-bearing and silicified horizons in the mass of the Király Hill indicates that steam-heated overprint may also be related to repeated hydrothermal activity associated with different levels of paleowater tables. At Botkő, which is located at 140m elevation, the most common homogenisation temperatures of secondary fluid inclusions from rock-forming quartz are about 140-150C and salinities are between 0.4 and 2.5 NaCl equiv. wt% (Molnár, 1988). The mode of homogenisation temperatures at about 140°C is considered as paleotemperature due to the presumed low pressure trapping conditions. This paleotemperature is similar to the temperature of steam- heated fluids in active geothermal systems (Hedenquist et al., 1990). However, salinities are slightly higher which may indicate the mixing of the recharging very dilute steam-heated fluids

(max. 0.8-1 NaCl equiv wt%, mostly resulted from the CO2 content of fluids; Hedenquist & Henley, 1985a, 1985b) with the upflowing hydrothermal fluids. This mixing may be one of the factors leading to the intensive silicification and the presence of steam-heated alteration with alunite beneath the silicified horizon (Figs. 60. and 61.) may also indicate the downward moving of the paleowater table during the hydrothermal activity (Sillitoe, 1993). Structural and volcanological data indicate that faulting with uplift of units east of the major fault of the area took place before the accumulation of the unaltered lower andesite lava flows of Sarmatian age. On the other hand, the Lower Sarmatian reworked rhyolitic tuff of Botkő is hydrothermally altered. Therefore the change in the level of paleowater table resulting steam-heated overprint on deeper alteration zones can be attributed to the tectonic processes synchronously affecting the area during the hydrothermal activity.

4. References

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