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Geothermics 50 (2014) 180– 188

Contents lists available at ScienceDirect

Geothermics

journa l h om epa ge: www.elsev ier .com/ locate /geothermics

ap rock efficiency and fluid circulation of natural hydrothermalystems by means of XRD on clay minerals (Sutri, Northern Latium,taly)

veva Corradoa,∗, Luca Aldegab, Antonio Stefano Celanoc, Arnaldo Angelo De Benedetti a,uido Giordanoa

Dipartimento di Scienze, Sezione di Scienze Geologiche, Università degli Studi Roma Tre, Largo San Leonardo Murialdo 1, 00146 Roma, ItalyDipartimento di Scienze della Terra, Università degli Studi Sapienza, Piazzale Aldo Moro 5, 00185 Roma, ItalyVia Vitaliano Brancati 65, 00144 Roma, Italy

r t i c l e i n f o

rticle history:eceived 28 December 2012ccepted 18 September 2013

a b s t r a c t

We performed XRD investigations on the sedimentary cap rock of the geothermal system developedin the area of Vico volcano (Northern Latium) to assess its effectiveness and degree of interactionwith fluids. The system consists of a positive thermal anomaly, a permeable carbonate reservoir atshallow depths and a low permeability siliciclastic cap rock. Unfractured cap rock shows maximum

eywords:ap-rockhermal alteration-ray diffractionixed layers illite-smectiteorthern Latium

taly

paleo-temperatures <50–60 ◦C, interpreted as the thermal signature of the original sedimentary basin.Fractured cap rock is characterized by kaolinite, calcite, short-range ordered mixed layers illite-smectitewith paleo-temperatures between 85 and 110 ◦C indicating strong interaction with hot fluids from acarbonate reservoir.

© 2013 Elsevier Ltd. All rights reserved.

. Introduction

Geothermal systems dominated by water convection consist ofour main elements: the heat source, the reservoir, the cap rock andhe recharge area. Most of the literature on geothermal explorations focussed on the characterization and modeling of the reservoire.g. Giordano et al., 2012; Ganguly and Kumar, 2012, for a review)nd on the evaluation of the P-T conditions (Erkan et al., 2008; Deenedetti et al., 2010; Rossetti et al., 2011; De Filippis et al., 2012).urthermore investigations with geothermal purposes start withydrogeological studies devoted to the regional characterization ofacro-circulation and thermo-baric state of fluids (Serpen, 2004;

a Vigna et al., 2012; Mazza et al., 2013).Less attention has been generally addressed to the characteri-

ation of the cap rock as a necessary component of the system thatuarantees the maintenance in the subsurface of the adequate flu-ds pressure and temperature conditions. Traditionally the cap rock

s described as a sufficiently, but not totally impermeable, thick and

idespread rock unit but its integrity and composition are oftennferred mainly through surface data with many uncertainties.

∗ Corresponding author. Tel.: +39 06 57338002.E-mail addresses: sveva.corrado@uniroma3.it, sveva.corrado@gmail.com

S. Corrado).

375-6505/$ – see front matter © 2013 Elsevier Ltd. All rights reserved.ttp://dx.doi.org/10.1016/j.geothermics.2013.09.011

Concerning seal integrity, surface monitoring of CO2 flux fromthe ground may provide discrete distribution of gas leakage fromactive systems (Chiodini et al., 2005; Annunziatellis et al., 2008).This method has been proven to be particularly useful in locat-ing successful production wells (Chiodini et al., 2007). MoreoverCO2 flux from the ground has been rarely compared with cap rockstructural integrity (Todesco and Giordano, 2010).

Cap rocks in volcanic contexts may be represented by largeignimbrites or fine ash-dominated deposits (e.g. base surges, falldeposits). In other cases cap rocks may be made up of altered pyro-clastic deposits that acquired cap efficiency by strong alterationdue to fluids interactions. Cap rocks geometry (e.g., thickness) isnot strictly related to the original depositional mechanisms but toweathering and alteration either due to temperature increase (Yanget al., 2001) or to endogenous fluids circulation, driven by perme-ability anisotropy (e.g., Todaka and Akasaka, 2004). On the contrary,in systems developed in sedimentary environments the cap rockthickness always depends on the original factors (e.g., sedimentssupply and composition) controlling deposition and compactionof rocks. Thus post-diagenetic thermal alteration in sedimentaryrocks due to hot fluids upward migration provides precious evi-

dence of two factors that are essential in evaluating the potential ofa geothermal system: (1) hydrological conditions prevailing withinan active geothermal system (e.g., temperatures and nature offluids; Battaglia et al., 2007) and (2) effectiveness of the cap rock.

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n contexts characterized by strong alteration processes mineraloning in the cap rock is controlled by temperature distribution,ock and fluid compositions, and rock textures and permeability.hanges in clay minerals structures and compositions are progradeith depth and temperature increases (Muffler and White, 1969;oagland and Elders, 1978; Battaglia et al., 2007). Nevertheless itay be difficult to detect burial and thermal history in low perme-

ble pelitic rocks (Wolde Gabriel et al., 2001). Thus the geochemicaltudy of veins cross cutting cap rocks have been lately adoptedGasparrini et al., 2012).

Concerning the effectiveness of the cap rock, the influence ofap rock thickness on fluids P-T conditions and system stabilityave been recently investigated by means of numerical modelingTimlin, 2009). Performed models allowed to define a critical thick-ess of the seal above which the system becomes more stable andurface thermal anomalies are not detectable (e.g., blind geother-al systems). A further aspect concerning cap rock effectiveness is

hat, at comparable boundary conditions, fluid-rock interaction onolcanic assemblages may be higher than in siliciclastic rocks (Coxnd Browne, 1998).

Aim of this paper is to present an example of thermal alter-tion of the out-cropping sedimentary cap-rock of a geothermalystem located in the Central Italy by means of semi-quantitative-ray diffraction analysis. This investigation was performed onoth whole-rock composition and clay grain-size fraction of sedi-ents integrated with structural analysis at the outcrop scale. This

pproach is widely very useful in geothermal exploration for:

characterizing alteration facies in hydrothermal systems in orderto detect the nature and degree of fluid-rock interaction (Cox andBrowne, 1998);deciphering clay minerals structural and compositional changes(e.g., stacking order and illite content in mixed layer illite-smectite) to determine the thermal maturity of sediments, fromdiagenesis to very low-grade metamorphism. Thus to constrainthe maximum temperatures experienced by rocks (Merriman andFrey, 1999).

The study area is located in one of the most geothermally inter-sting regions of the Italian peninsula extended between Tolfa Mts.nd Vico and Bracciano Lakes (Fig. 1). In this area there are sev-ral surface manifestations such as mineral springs, hydrothermaleposits and diffuse degassing, linked to the Pliocene and Qua-ernary volcanism along the peri-Tyrrhenian margin. These piecesf evidence suggest the presence of magmatic heat sources and aermeable carbonate substratum at shallow depths (Barberi et al.,994; Cinti et al., 2011). Within the study area, the outcroppingertiary siliciclastic deposits (e.g., sandstones and pelites of theuscan succession) potentially represent the cap rocks of a shalloweso-Cenozoic carbonate reservoir (Aldega et al., 2010).

. Geological setting – main geothermal features

Pliocene-Quaternary magmatism and crustal extension linkedo the Tyrrhenian Sea evolution are the most distinctive geologicaleatures recorded along the peri-Tyrrhenian margin in Northernatium from the Vulsini Mts toward the South to the Roman areaMattei et al., 2010 and references therein). This evidence makeshe area one of the most perspective for geothermal explorationn Italy (Cataldi et al., 1995; Doveri et al., 2010). The Quaternaryolcanic complexes in Northern Latium are characterized by diff-

sed surface geothermal manifestations and a series of geothermaleservoirs at shallow depths (between 1 and 2 km) drilled betweenhe 1970s and the early 1990s for high-medium enthalpy targetse.g., Torre Alfina, Latera, Cesano geothermal fields: Funiciello and

cs 50 (2014) 180– 188 181

Parotto, 1978; Buonasorte et al., 1987, 1995; Barberi et al., 1994;Cataldi et al., 1995). The heat flow along the peri-Tyrrhenian mar-gin in Northern Latium is higher than 100 mW/m2 with localizedmaxima exceeding 250 mW/m2 (Buonasorte et al., 1995) and lateralgradients varying from sharp (50 mW over a linear distance of lessthan 1 km) to diffuse (50 mW over a linear distance of more than7 km) (Buonasorte et al., 1987; Cataldi et al., 1995). These maximaare generally outlined by CO2 pressure values higher than 0.18 barrecorded in waters at shallow depths (Doveri et al., 2010). Heatflow rapidly decreases to values lower than 50 mW/m2 moving tothe East towards the Appennine chain.

The stratigraphy of the typical hydrothermal system of NorthernLatium is characterized from top to bottom by (Buonasorte et al.,1987; Barberi et al., 1994; Capelli and Mazza, 2005):

- Pleistocene volcanics and alluvial deposits with medium perme-ability and thicknesses varying from 0 to 500 m.

- Upper Miocene-Pliocene clayey-arenaceous neo-autochtonousdeposits with low permeability and thicknesses ranging from 0to 1000 m, generally acting as cap rocks (CAP1 in Fig. 2).

- Highly tectonically disrupted calcareous-marly-siliciclastic suc-cessions belonging to Cretaceous-Oligocene pre- and syn-orogenic sedimentary cycles (e.g., Ligurian and sub-Ligurian unitsand foredeep deposits) with low permeability and variable thick-nesses that range between a few hundred meters to 2 km, actingas cap rock (CAP 2 in Fig. 2).

- Jurassic-Eocene calcareous-marly-siliceous sedimentary multi-layer that acts as reservoir (RES 1 in Fig. 2) with frequent clayeyinterlayers and variable permeability as a function of clay contentand fracturing of carbonates. Thicknesses range between 400 and1000 m.

The regional structural evolution was characterized by eastwardmigration of Apennines compression since Oligocene-Miocene(Meneghini et al., 2012) followed from Upper Miocene to Qua-ternary by extension-transtension linked to the Tyrrhenian Seadynamics (Mattei et al., 2010; Fig. 2).

3. Methods and materials

Clay minerals in shales and sandstones undergo diagenetic andvery low-grade metamorphic reactions in response to heating.Reactions in clay minerals are irreversible under normal diageneticto epizonal conditions, so that sedimentary sequences generallyretain indices and fabrics indicative of their maximum thermalalteration. In particular, mixed-layer clay minerals (I-S) and theirtransformation sequence, di-smectite-random ordered mixed lay-ers (R0)-ordered mixed layers (R1 and R3)-illite-di-octahedralK-mica (muscovite), can be used as indicators of the thermal evo-lution of sedimentary sequences (Corrado et al., 1998, 2009; Bottiet al., 2004; Aldega et al., 2007, 2011) and as a tool for exploration(Pollastro, 1990).

Investigated lithotypes are mainly pelites and sandstones. Theyare attributed to the siliciclastic unit stratigraphically overlyingthe reservoir of the Tuscan-Umbrian successions and to the vari-coloured clays belonging to the Ligurian domain which thrust overthe siliciclastic rocks (Fig. 3).

Samples were analyzed by XRD using a Scintag X1 system (CuK� radiation) at 40 kV and 45 mA. Randomly oriented whole-rockpowders were run in the 2–70◦ 2� interval with a step size of0.05◦ 2� and a counting time of 3 s per step. The <2 �m grain-

size fractions were separated by centrifuging without crushingthe samples and oriented slides were prepared by the pipette-on-slide method, keeping the specimen thickness as constant aspossible, within the range of 1–3 mg of clay per cm2 of glass slide

182 S. Corrado et al. / Geothermics 50 (2014) 180– 188

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ig. 1. Hillshade from Digital Elevation Model of Northern Latium (Central Italy) wafter Buonasorte et al., 1995). Study area is located to the South of the Vico Lake in

Giampaolo and Lo Mastro, 2000). Oriented air-dried and ethylene-lycol–solvated samples were scanned from 1◦ to 48◦ 2� and from◦ to 30◦ 2�, respectively, with a step size of 0.05◦ 2� and a countime of 4 s per step. The illite content in mixed-layer I-S wasetermined according to Moore and Reynolds (1997) using theelta 2� method after decomposing the composite peaks between◦–10◦ 2� and 16◦–17◦ 2�. Peaks in relatively close position (in�) and at angles >10◦ 2� were selected for clay mineral semi-uantitative analysis of the <2 �m (equivalent spherical diameter)raction in order to minimize the angle-dependent intensity effect.he relative proportions of clay minerals were calculated by using

ineral intensity factors as calibration constant (see Moore and

eynolds, 1997 for details). Integrated intensity were measured byecomposing composite peaks [e.g., kaolinite(0 0 1)-chlorite(0 0 2)nd I(0 0 2)-S(0 0 3)] through Pearson VII functions and the DMSNT

elevant geological units of the regional geothermal system and heat flow isolinesuared area between the towns of Sutri and Ronciglione.

Scintag associated program and then normalized to 100%. Non-clayminerals (e.g., quartz) identified in the <2 �m grain-size fraction,were not included in the quantitative analysis of the oriented aggre-gates, thus the data refer to the phyllosilicates group only.

4. Results

4.1. Structural and stratigraphic data

The study area extends for about 7 km2 along a NW-SE elongatedridge to the SE of the Vico Lake in the surroundings of the towns

of Sutri and Ronciglione (Fig. 1). The ridge, with elevations notexceeding 350 m a.s.l., represents a structural high of the sedimen-tary Meso-Cenozoic marine substratum unconformably coveredby Pleistocene volcanoclastic products of the Vico and Sabatini

S. Corrado et al. / Geothermics 50 (2014) 180– 188 183

Fig. 2. Simplified cross section showing the tectonic and stratigraphic setting of the Northern Latium geothermal area. 1: Pleistocene volcanics and continental clastics; 2:Upper Miocene-Pliocene clayey-arenaceous neo-autochtonous complex (a: CAP1) and Cretaceous-Oligocene pre and syn-orogenic sedimentary cycles (b: CAP2); 3: Jurassic-E ower

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ource: Redrawn and modified after Capelli and Mazza (2005).

olcanic complexes (Figs. 1 and 3a). Volcanic deposits are gener-lly preserved along the flanks and river incisions surrounding thetructural relief and consists of two main lithotypes:

a few tens of meter-thick red brown ignimbrite with highly vesic-ulated black scoriae, at the bottom (Sutri Fm., Bear et al., 2009);gray massive ignimbrites rich in lithics and leucite crystals at thetop.

The ridge backbone is mainly made up of decimetric to centimet-ic alternating medium-fine grained beds of turbiditic sandstonesnd pelites, dipping to the SW by about 35◦–45◦ (Fig. 3d). Thisnit has been interpreted as the upper portion of the Tuscan-mbrian succession (Buonasorte et al., 1987). The turbiditic unit isverthrusted along a low angle thrust surface by well bedded mud-tones locally alternated with reddish to greenish scaly shales ofhe Ligurian domain, moderately dipping to SW and S. The tectonicontact crops out along the south-western flank of Mt. Croce ands marked by varicoloured shales. Furthermore local detachments

ithin the calcareous-marly unit enhanced the development ofesoscale asymmetric SW-verging folds in the hangingwall. These

ompressive structures developed in Late Oligocene-Early Mioceneimes and were dissected by NW-SE and SW-NE extensionalaults in Upper Miocene-Lower Pliocene. These faults uplifted theresent-day structural high and finally were sutured by Pleistoceneolcanic products. To the SE of Mt. Croce these buried faults arearked at surface by a series of CO2 emissions aligned along a

E-NW trend (Capelli and Mazza, 2005, Fig. 3b).Structural analysis was concentrated along the SW flank of

t. Croce (Fig. 3a1 and a2). Here the thrust footwall (siliciclas-ics belonging to the Tuscan-Umbrian domain) is affected byub-vertical not-mineralised tensile fractures, developed in Lateligocene-Early Miocene times, striking between N020◦ and N040◦,ainly orthogonal to bedding that moderately dips to the SW

Fig. 3a2). The fractures H/S (Height/Spacing) ratio is of about.4, typical of compressive regimes (Tavani et al., 2010). Theangingwall presents two sets of non-mineralised fractures eithererpendicular (between N030 and N060◦) (Fig. 3a1) or parallel toedding. They may have formed analogously to the fracture sets

reserved in the footwall. They may be interpreted either as enu-leated by the compressive regime and then slightly rotated duringectonic transport along the Mt. Croce thrust or developed latern in the extensional regime. Meso-scale asymmetrical secondary

Jurassic); 4: Triassic succession (RES2); 5: thermo-metamorphic halo; 6: projected

folds parallel to the regional thrust also developed in the hanging-wall block.

4.2. X-ray diffraction data

Samples for XRD investigation derive from footwall and hang-ingwall units either on unfractured rocks or along footwall fracturesand thrust plane separating Ligurian from Tuscan units (Fig. 3c).

Analysis of randomly oriented whole-rock powder patterns andof oriented mounts of the <2 �m grain-size fraction (summarizedin Tables 1 and 2, Figs. 4 and 5) allowed to group XRD results in fivedistinct classes that correspond to specific structural or lithologicalfeatures.

(1) Footwall unfractured rocks: The whole-rock composition ofunfractured fine-grained sandstones (SU1, SU2) and pelites(SU3, SU5) from the footwall is made up of phyllosilicates(mica and clay minerals), quartz, albite, minor amounts ofK-feldspar and traces of goethite with a relative higher abun-dance of quartz (21–23%) and albite (37–38%) in sandstoneswhen compared to pelites (respectively 7–8% and 13–14%).Oriented mounts of the <2 �m grain-size fraction display illite-rich assemblages, which constitute at least 48% of the overallcomposition, mixed-layered minerals with amounts of about30%, and subordinate amounts of kaolinite (up to 18%). Chloritelocally occurs in pelites with percentages that do not exceed 1%.Non-clay minerals such as quartz, albite, goethite and gismond-ina have been also observed in this fraction. Mixed-layered clayminerals generally consist of random ordered I-S with an illitecontent of about 30–33% (Table 2).

(2) Footwall damage zone: Sample SU11 has a whole-rock composi-tion similar to that observed in footwall unfractured sandstonesexcept for the presence of calcite (12%) and lower amounts ofalbite (29%). On the contrary the mineralogical assemblage ofthe <2 �m grain-size fraction differs from those of group 1 withsmectite as major component (96%) and subordinate amountsof illite and kaolinite (2% each). Non-clay minerals such asquartz, calcite, K-feldspar and albite have been identified aswell.

(3) Fractures fill in the footwall: Highly plastic green clays filling

fractures from the footwall (SU4 and SU6) are made up of morethan 90% of phyllosilicates with low amounts of quartz andalbite and traces of K-feldspar and goethite. The <2 �m grain-size fraction contains kaolinite (mean value 50%), illite (25%)

184 S. Corrado et al. / Geothermics 50 (2014) 180– 188

Fig. 3. (a) Geological map of the study area between Fonte Vivola locality and Mt. Croce, with lower hemisphere projections (Schmidt diagram) by Daisy3 software (byFrancesco Salvini, version 2010) of hangingwall (a1) and of footwall fractures (a2). (b) Simplified structural section across the southern flank of the structural high whereCO2 shows and normal faults downthrowing the structure are located; 1: Pleistocene volcanics and continental clastics; 2: Upper Miocene-Pliocene clayey-arenaceous neo-autochtonous complex; 3: Cretaceous-Oligocene pre- and syn-orogenic sedimentary cycles; 4: inferred normal fault; 5: well; 6: water table; 7: gas manifestation; redrawnafter Capelli and Mazza, 2005; (c) Inset of map a for location of samples and of structural analysis sites. (d) Simplified geological sections perpendicular to the regional foldaxis (traces in map a).

Table 1Mineralogy and relative abundance of the mineralogical phases of the whole-rock samples. Qtz = quartz; Cal = calcite; Kfs = K-feldspar; Ab = albite; Phy = phyllosilicates;Hem = hematite; Go = Goethite; tr = traces.

Sample Unit Rock type Whole-rock composition

Qtz Cal Kfs Ab Phy Hem Go

SU1 Siliciclastics Sandstone 21 – 1 37 41 – trSU2 Siliciclastics Sandstone 23 – 1 37 39 – –SU3 Siliciclastics Pelite 8 – – 14 78 – trSU4 Siliciclastics Alteration 6 – tr 3 91 – trSU5 Siliciclastics Pelite 7 – 1 13 79 – –SU6 Siliciclastics Alteration 4 – tr 2 94 – –SU11 Siliciclastics Sandstone 22 12 1 29 36 – –SU8 Ligurian Pelite 6 3 – 1 90 – –SU10 Ligurian Pelite 5 2 – 1 90 2 –SU9 Ligurian Alteration 2 66 – – 32 – –

S. Corrado et al. / Geothermics 50 (2014) 180– 188 185

Table 2X-ray semi-quantitative analysis of the <2 �m grain-size fraction. Sm = smectite; I = illite; I-S = mixed layer illite-smectite; K = kaolinite; Chl = chlorite; Qtz = quartz; Ab = albite;Go = goethite; Gis = gismondine; Cal = calcite; Kfs = K-feldpar; R = stacking order (Jagodzinski, 1949); %I in I-S = illite content in mixed layer illite-smectite.

Sample Unit X-ray quantitative analysis of the <2 �m grain-size fraction (%wt.) R %I in I-S

Sm I I-S K Chl Other

SU1 Siliciclastics – 48 35 17 – Qtz, Ab, Go 0 30SU2 Siliciclastics – 49 33 18 – Qtz, Ab 0 30SU3 Siliciclastics – 60 28 11 1 Qtz, Ab, Go 0 32SU4 Siliciclastics – 24 22 54 – Qtz, Go 0 54SU5 Siliciclastics – 67 29 3 1 Qtz, Ab, Gis 0 33SU6 Siliciclastics – 26 29 45 – Qtz 0/1 40/70SU11 Siliciclastics 96 2 – 2 – Qtz, Ab, Cal, Kfs – –

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SU8 Ligurian – 10 46

SU10 Ligurian 7 13

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and mixed layers I-S (25%) as major minerals. Two populationsof mixed-layer I-S are observed. They are random ordered R0I-S with an illite content of 40% and low expandable R1 I-S withan illite content of 70%.

4) Hangingwall unfractured rocks: Varicoloured clays (sample SU8)show a mineralogical assemblage mainly composed of phyl-losilicates, with minor amounts of quartz, calcite and albite(Table 1). The <2 �m grain-size fraction contains mainly kaolin-ite (40%) and mixed layer I-S (46%) and subordinate amountsof illite (10%) and chlorite (4%). Mixed-layered clay mineralsconsist of random ordered I-S with an illite content of 30%.

5) Hangingwall damage zone: Two samples have been collected (SU

9 and SU10) from fractures walls. Sample SU9 shows a predom-inance of calcite and phyllosilicates, whereas sample SU10 hasa mineralogical assemblage more similar to that observed ingroup 4. Stronger affinities exist in the <2 �m grain-size fraction

ig. 4. Schematic field section at the footwall of Mt. Croce thrust with location of collectnd <2 �m fraction. Qtz = quartz; Kfs = K-feldspar; Ab = albite; Phy = phyillosilicates; Sm =

0 4 Qtz 0 305 5 Qtz 0/1 34/658 5 Qtz, Cal 0/1 33/64

of this group. Clay minerals are mainly composed of kaolinite(higher than 70%) and subordinate amounts of mixed layer I-S (about 10%), illite (7–8%) and chlorite (5%). Mixed-layer clayminerals are R0 and R1 I-S with illite contents of 30 and 65%respectively.

5. Discussion

In order to better constrain the seal effectiveness of the NorthernLatium geothermal system, XRD and structural results acquired inthe study area are compared:

(1) with similar outcrops in the central Apennines developed insectors with lower heat flux;

(2) to worldwide case histories on thermal alteration of clays ingeothermal systems.

ed samples and related graphs showing percentages of minerals of the whole rocksmectite; I = illite; I-S = illite-smectite mixed layers; K = kaolinite; Chl = clorite.

186 S. Corrado et al. / Geothermics 50 (2014) 180– 188

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ig. 5. Schematic field section across Mt. Croce thrust with location of collected saraction. Qtz = quartz; cal = calcite; Kfs = K-feldspar; Ab = albite; Phy = phyillosilicateshl = clorite.

In the footwall rocks, XRD analyses on the clayey-arenaceousnit show kaolinite crystallization within vertical fractures and twoopulations of mixed layers I-S (R0 and R1) along the wall rocks.n the other hand, the unfractured portion of the same unit is char-cterized by highly expandable R0 illite-smectite. The hangingwallonsists of undeformed rocks characterized by mixed layers illite-mectite with low amounts of illite layers (30%) and kaolinite and

more deformed portion where two populations of mixed layersllite-smectite (R0 and R1) occur. Moreover calcite dominates inracture fills.

These data show that unfractured rocks from both the hanging-all and footwall record early diagenetic conditions characterized

y R0 I-S with illite contents of 30% that are indicative of temper-tures lower than 50–60 ◦C (Hoffman and Hower, 1979; Merrimannd Frey, 1999). These data are consistent with those from coldhermal regimes, typical of deep marine environments (Allen andllen, 2005). These environments developed in post-rift and fore-eep setting as in the case of varicoloured clays of the Ligurianomain and the Tuscan siliciclastics, respectively, with sedimenta-ion rates from very low to high. These kinds of basins are generallyharacterized by low heat flow values (e.g. 1HFU, cfr Allen and Allen,005). Furthermore similar siliciclastic deposits were analyzed aew tens of km farther NNE in the surroundings of the Trasimenoake (e.g., Aquerino unit in ISPRA National geological map nr. 310cale 1:50,000). These deposits show similar results in terms ofhermal maturity derived from vitrinite reflectance data (imma-ure stage of hydrocarbon generation) which were interpreted as

result of shallow sedimentary/tectonic burial at paleo-gradient

onditions of <25 ◦C/km (Caricchi et al., 2011). As the unfracturedamples from the footwall and hangingwall indicate maximumemperatures lower than 60–70 ◦C, the paleogeothermal gradient,f assumed similar to that of the Aquerino unit, defines a maximum

and related graphs showing percentages of minerals of the whole rock and <2 �m = hematite; Sm = smectite; I = illite; I-S = illite-smectite mixed layers; K = kaolinite;

burial lower than 2 km and the influence of hot fluids circulationseems negligible.

Mineral assemblages such as smectite and mixed layer illite-smectite ± calcite or kaolinite detected along fracture walls and invein filling both in the footwall and in the hangingwall rocks areinstead typical of sediment-hosted geothermal systems (McDowelland Elders, 1980; Schiffman and Evarts, 1984; Cho et al., 1988) thatnormally form in shallow zones characterized by: (1) diagenetictemperatures, (2) interaction with hot and CO2 rich fluids, and (3)an increase of illite layers in mixed layers illite-smectite as functionof temperature. The mineralogical composition of the clay matrixchanges with increasing distance from the fractures. Kaolinite con-stitutes the prevailing clay mineral in the fracture filling of footwallrocks whereas smectite is more abundant farther in the altered wallrocks and in the thrust damage zone. These results are consistentwith the intermediate clayey alteration described by Wohletz andHeiken (1992) characterized by the kaolinite and smectite associa-tion, with kaolinite forming nearer veins and smectite farther fromveins. The formation of these two minerals, in fact, is strongly influ-enced by chemical factors such as pH and Na+/H+ and Ca2+/H+ ratio(Velde, 1985). In zones of low fluid/rock ratio, the interaction offluids with the host rock may increase the pH value or Na+/H+ andCa2+/H+ ratio and favor smectite crystallization. On the other hand,in current fluid circulation zones or in zone with high water/rockratio, the crystallization of kaolinite is favored.

Comparison with similar studies worldwide and with temper-ature models for I-S geothermometry may provide more preciseconstraints about the temperature of the illitization process in

this hydrothermal system. Muffler and White (1969) reportedthe appearance of R1 ordering in the Salton Sea (California,USA) hydrothermal system at 98–135 ◦C in reasonable agreementwith R1 occurrence at 85–110 ◦C in the Wairakei (New Zealand)

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ydrothermal system (Srodon and Eberl, 1984; Harvey and Browne,991) or at about 100–110 ◦C in Milos island (Greece; Christidis,995). Nevertheless the appearance of I-S mixed layers with morehan 95% illite has been reported at 200 ◦C in the Shinzan areaJapan; Inoue and Utada, 1983), at 203–217 ◦C in the Salton Seaystem (Muffler and White, 1969), and at 230–240 ◦C in WairakeiSrodon and Eberl, 1984). Furthermore Harvey and Browne (1991)eported illite formation at 200 ◦C and formation of I-S with 80–90%f illite layers at 155–190 ◦C in the Wairakei hydrothermal field andhristidis (1995) in Milos island correlated I-S with 87% of illite lay-rs generically at temperatures lower than 200 ◦C. Regarding theemperature models for I-S geothermometry, applied to short-livedeating events as in areas of relatively recent thermal activity andlevated geothermal gradients and/or hydrothermal environments,n I-S mixed layers the transition between R0 and R1 I-S occurs in aemperature range of 120◦–140◦ (Jennings and Thompson, 1986).

Following previous interpretations our host rocks close toirculation conduits are characterized by the coexistence of R0nd R1 mixed layer I-S that can be interpreted as indicative ofulses of hot fluids whose temperature ranged between 85 and40 ◦C. Further comparison with comparable case histories ana-

yzed with methodologies analogous to those adopted in this paperould in future contribute to better constrain this temperatureange.

. Conclusions

We performed XRD investigations of the whole-rock compo-ition and <2 �m fraction of the sedimentary cap rock of theeothermal system developed in the area of Vico volcano closeo Ronciglione and Sutri towns. This investigation allowed toecognize differences in clay minerals composition between thenfractured and fractured rock assemblages of the cap rock.nfractured cap rock shows low thermal maturity and maximum

emperatures lower than 50–60 ◦C which are interpreted as thehermal signature of the original cold sedimentary basins. Fracturedap rock is characterized by kaolinite, calcite and the coexistencef R0 and R1 illite-smectite. These evidences indicate interactionith hot fluids deriving from the buried carbonate reservoir whose

emperatures are between 85 and 140 ◦C.From our data we infer that the cap rock in the study area is effec-

ive as indicated by the preservation of the original sedimentaryhermal signature (e.g., cold thermal regime). The fluid circulations instead structurally-controlled and recorded only by limited clayecrystallization along the walls of discrete tensile fractures. Thusractures and thrust faults pre-dating the still active hydrothermalctivity acted as passive anisotropies that focussed localized alter-tion.

In conclusion the study of thermal maturity of siliclastic rockss a function of fracturing distribution of cap rocks has shown howompositional and structural changes in mixed layers I-S are func-ion of hydrothermal fluid temperature and fractures distribution.n addition, the distribution of neo-formed minerals such as calcitelong tectonic discontinuities helped to trace the path of hot-fluidirculation at shallow depths. Hence this approach may turn out toe useful in areas of geothermal interest where either subsurfaceata are scarce or surface thermal anomalies are lacking (e.g., blindeothermal systems).

cknowledgments

We are greatly indebted to S. Lo Mastro for providing facil-ties for X-ray diffraction analysis in Roma Tre University, F.alvini for the use of the structural software Daisy3, F. Balsamo ishanked for his help in the field. Two anonymous referees and the

cs 50 (2014) 180– 188 187

Associate Editor David Bruhn are kindly acknowledged for theirprecious comments and suggestions that contributed to improvethe original version of the manuscript. Research financial supportwas provided by the Department of Geological Sciences, “Roma Tre”University Fundings to G. Giordano and S. Corrado.

References

Aldega, L., Botti, F., Corrado, S., 2007. Clay mineral assemblages and vitrinitereflectance in the Laga Basin (Central Apennines, Italy): what do they record?Clays and Clay Minerals 55 (5), 504–518.

Aldega, L., Corrado, S., De Benedetti, A.A., Giordano, G., 2010. Valutazione emappatura dell’efficacia delle rocce sigillo nei sistemi geotermici del Laziosettentrionale: una nuova metodologia integrata mineralogica e petrograficaorganica. In: GeothermEXPO 2010, Abstract volume, Ferrara 21–23.09.2010.

Aldega, L., Corrado, S., Di Paolo, L., Somma, R., Maniscalco, R., Balestrieri, M.L., 2011.Shallow burial and exhumation of the Peloritani Mts. (NE Sicily, Italy): insightfrom paleo-thermal and structural indicators. Geological Society of AmericaBulletin 123, 132–149.

Allen, P.A., Allen, J.R., 2005. Basin Analysis, Principles and Application, second ed.Blackwell Scientific Publications, Oxford.

Annunziatellis, A., Beaubien, S.E., Bigi, S., Ciotoli, G., Coltella, M., Lombardi, S., 2008.Gas migration along fault systems and through the vadose zone in the Lateracaldera (central Italy): implications for CO2 geological storage. InternationalJournal of Greenhouse Gas Control 2, 353–372.

Barberi, F., Buonasorte, G., Cioni, R., Fiordelisi, A., Foresi, L., Iaccarino, S., Laurenzi,M.A., Sbrana, A., Vernia, L., Villa, I.M., 1994. Plio-pleistocene geological evolutionof the geothermal area of Tuscany and Latium. Memorie Descrittive della CartaGeologica d’Italia 49, 77–134.

Battaglia, S., Gherardi, F., Gianelli, G., Leoni, L., Origlia, F., 2007. Clay mineral reactionsin an active geothermal area (Mt. Amiata, southern Tuscany, Italy). Clay Minerals42, 353–372.

Bear, A.N., Cas, R.A.F., Giordano, G., 2009. Variations in eruptive style anddepositional processes associated with explosive, phonolitic composition,caldera-forming eruptions: The 151 ka Sutri eruption, Vico Caldera, central Italy.Journal of Volcanology and Geothermal Research 184 (3–4), 1–24.

Botti, F., Aldega, L., Corrado, S., 2004. Sedimentary and tectonic burial evolution ofthe Northern Apennines in the Modena-Bologna area: constraints from com-bined stratigraphic, structural, organic matter and clay mineral data of Neogenethrust-top basins. Geodinamica Acta 17 (3), 185–203.

Buonasorte, G., Fiordilesi, A., Pandeli, E., Rossi, U., Sollevanti, F., 1987. Stratigraphiccorrelation and structural setting of the pre-neoautochthonous sequences ofNorthern Latium. Periodico di Mineralogia 56, 111–112.

Buonasorte, G., Cameli, G.M., Fiordilesi, A., Parotto, M., Perticone, I., 1995. Results ofgeothermal exploration in Central Italy (Latium-Campania). Proceedings of theWorld Geothermal Congress. Florence, International Geothermal Association 2,1293–1298.

Capelli, G., Mazza, R., 2005. Inquadramento geologico del dominio vulcanico laziale.In: Capelli, G., Mazza, R., Gazzetti, C. (Eds.), Strumenti e strategie per la tutela el’uso compatibile della risorsa idrica nel Lazio, Quaderni di tecniche di protezioneambientale, vol. 78. Pitagora editrice, Bologna, pp. 15–39.

Caricchi, C., Aldega, L., Botti, F., Cifelli, F., Corrado, S., Mattei, M., Speranza, F., 2011.Preliminary results on the reconstruction of the Cenozoic thermal and kinematicevolution of the external portion of the Tuscan nappe (northern Apennines).Epitome 4, 147.

Cataldi, R., Monelli, F., Squarci, P., Taffi, L., Zito, G., Calore, G., 1995. Geothermalranking of the Italian territory. Geothermics 24, 115–129.

Chiodini, G., Granieri, D., Avino, R., Caliro, S., Costa, A., Werner, C., 2005. Car-bon dioxide diffuse degassing and estimation of heat release from volcanicand hydrothermal systems. Journal of Geophysical Research – Solid Earth 110,B082041.

Chiodini, G., Baldini, A., Barberi, F., Carapezza, M.L., Cardellini, C., Frondini, F.,Granieri, D., Ranaldi, M., 2007. Carbon dioxide degassing at Latera caldera (Italy):evidence of geothermal reservoir and evaluation of its potential energy. Journalof Geophysical Research – Solid Earth 112, B12204.

Cho, M., Liou, J.G., Bird, D.K., 1988. Prograde phase relations in the State 2-14 wellmetasandstones, Salton Sea geothermal field, California. Journal of GeophysicalResearch 93, 13081–13103.

Christidis, E.G., 1995. Mechanisms of illitization of bentonites in the geother-mal field of Milos Island Greece: evidence based on mineralogy, chemistry,particle thickness and morphology. Clays and Clay Minerals 43 (5),569–585.

Cinti, D., Procesi, M., Tassi, F., Montegrossi, G., Sciarra, A., Vaselli, O., Quattrocchi, F.,2011. Geochemistry and geothermometry in the western sector of the SabatiniVolcanic District and the Tolfa Mountains (Central Italy). Chemical Geology 284,160–181.

Corrado, S., Di Bucci, D., Naso, G., Giampaolo, C., Adatte, T., 1998. Applicationof organic matter and clay mineral studies to the tectonic history of theAbruzzo-Molise-Sannio area, Central Apennines, Italy. Tectonophysics 285,

167–181.

Corrado, S., Aldega, L., Balestrieri, M.L., Maniscalco, R., Grasso, M., 2009. Structuralevolution of the sedimentary accretionary wedge of the alpine system in East-ern Sicily: thermal and thermochronological constraints. Geological Society ofAmerica Bulletin 121, 1475–1490.

1 hermi

C

D

D

D

E

F

G

G

G

G

H

H

H

I

J

J

L

M

88 S. Corrado et al. / Geot

ox, M.E., Browne, P., 1998. Hydrothermal alteration mineralogy as an indicator ofhydrology at the Ngawha geothermal field, New Zealand. Geothermics 27 (3),259–270.

e Benedetti, A.A., Caprilli, E., Rossetti, F., Giordano, G., 2010. Metamorphic, metaso-matic and intrusive xenoliths of the Colli Albani volcano and their significancefor the reconstruction of the volcano plumbing system. In: Funiciello, R., Gior-dano, G. (Eds.), The Colli Albani Volcano, Special Publication of IAVCEI, vol. 3.The Geological Society, London, pp. 153–176.

e Filippis, L., Faccenna, C., Billi, A., Anzalone, E., Brilli, M., Özkul, M., Soligo, M., Tuc-cimei, P., Villa, I.M., 2012. Growth of fissure ridge travertines from geothermalsprings of Denizli Basin, western Turkey. Geological Society of America Bulletin124 (9–10), 1629–1645.

overi, M., Lelli, M., Marini, L., Raco, B., 2010. Revision, calibration, and application ofthe volume method to evaluate the geothermal potential of some recent volcanicareas of Latium, Italy. Geothermics 39 (3), 260–269.

rkan, K., Holdmann, G., Benoit, W., Blackwell, D., 2008. Understanding the ChenaHot Springs, Alaska, geothermal system using temperature and pressure datafrom exploration boreholes. Geothermics 37 (6), 565–585.

uniciello, R., Parotto, M., 1978. Il substrato sedimentario nell′area dei ColliAlbani: considerazioni geodinamiche e paleogeografiche sul margine tirrenicodell′Appennino centrale. Geologica Romana 17, 233–287.

anguly, S., Kumar, M.S.M., 2012. Geothermal reservoirs – a brief review. Journal ofthe Geological Society of India 79 (6), 589–602.

asparrini, M., Ruggieri, G., Chiodini, G., 2012. Burial vs. hydrothermal calciteveins and the record of ancient geothermal systems from the Monte Ami-ata region (Italy). In: Geofluids VII Conference-Paris 6-8.06.2012, Extendedabstracts. Available on http://www.geofluids2012.com/presentation-files.html(accessed December 2012).

iampaolo, C., Lo Mastro, S., 2000. Analisi (semi)quantitativa delle argille mediantediffrazione araggi X. In: Fiore, S. (Ed.), Incontri Scientifici, vol. II. Istituto di Ricercasulle Argille, Potenza, pp. 109–146.

iordano, G., Pinton, A., Cianfarra, P., Baez, W., Chiodi, A., Viramonte, J., Norini, G.,Groppelli, G., 2012. Structural control on geothermal circulation in the CerroTuzgle–Tocomar geothermal volcanic area (Puna plateau, Argentina). Journal ofVolcanology and Geothermal Research 249, 77–94.

arvey, C.C., Browne, P.R.L., 1991. Mixed-layer clay geothermometry in the Wairakeigeothermal field, New Zealand. Clays and Clay Minerals 39, 614–621.

oagland, J.R., Elders, W.A., 1978. Hydrothermal mineralogy and isotopic geochem-istry in the Cerro Prieto geothermal Field, Mexico. I. Hydrothermal mineralzonation. Geothermal Resources Council 2, 283–286.

offman, J., Hower, J., 1979. Clay mineral assemblages as low grade metamorphicgeothermometers – application to the thrust faulted disturbed belt of Mon-tana, USA. In: Scholle, P.A., Schluger, P.S. (Eds.), Aspects of Diagenesis. Society ofEconomic Paleontologists and Mineralogists, Special Publication 2, pp. 55–79.

noue, A., Utada, M., 1983. Further investigations of a conversion series of diocta-hedral mica/smectites in the Shinzan hydrothermal alteration area, northeastJapan. Clays and Clay Minerals 31, 401–412.

agodzinski, H., 1949. Eindimensionale Fehlordnung in Kristallen und ihr Einflussauf die Röntgen Interferenzen. Acta Crystallographica 2, 201–207.

ennings, S., Thompson, G.R., 1986. Diagenesis of Plio-Pleistocene sediments of theColorado River delta, southern California. Journal of Sedimentary Petrology 56,89–98.

a Vigna, F., Mazza, R., Capelli, G., 2012. Detecting the flow relationships betweendeep and shallow aquifers in an exploited groundwater system, using long-termmonitoring data and quantitative hydrogeology: the Acque Albule basin case(Rome, Italy). Hydrological Processes, Published online in Wiley Online Library.Available on http://onlinelibrary.wiley.com/doi/10.1002/hyp.9494/abstract.

http://dx.doi.org/10.1002/hyp.9494

attei, M., Conticelli, S., Giordano, G., 2010. The Tyrrhenian margin geological set-ting: from the Apennine orogeny to the K-rich volcanism. In: Funiciello, R.,Giordano, G. (Eds.), The Colli Albani Volcano, Special Publication of IAVCEI, vol.3. The Geological Society, London, pp. 7–27.

cs 50 (2014) 180– 188

Mazza, R., Taviani, S., Capelli, G., De Benedetti, A.A., Giordano, G., 2013. Quantitativehydrogeology of volcanic lakes with management, volcanological and geother-mal implications. In: Tassi, F., Christenson, B., Vandemeulebrouck, J., Rouwet,D. (Eds.), Volcanic Lakes, Advances in Volcanology IAVCEI Book Series. Springer,Heidelberg (in press).

McDowell, S.D., Elders, W.A., 1980. Authigenic layer silicate minerals in boreholeElmore I, Salton Sea Geothermal field, California, USA. Contributions to Mineraland Petrology 74, 293–310.

Meneghini, F., Botti, F., Aldega, L., Boschi, C., Corrado, S., Marroni, M., Pandolfi, L.,2012. Hot fluid pumping along shallow-level collisional thrusts: the MonteRentella Shear Zone, Umbria Apennine, Italy. Journal of Structural Geology 37,36–52.

Merriman, R.J., Frey, M., 1999. Patterns of very low-grade metamorphism inmetapelitic rocks. In: Frey, M., Robinson, D. (Eds.), Low-grade Metamorphism.Blackwell Science, Oxford, pp. 61–107.

Moore, D.M., Reynolds Jr., R.C., 1997. X-Ray Diffraction and the Identification andAnalysis of Clay Minerals. Oxford University Press, Oxford, 378 p.

Muffler, L.J., White, D.E., 1969. Active metamorphism of upper Cenozoic sediments inthe Salton Sea geothermal field and the Salton Trough, Southeastern California.Geological Society of American Bulletin 80, 157–180.

Pollastro, R.M., 1990. The illite/smectite geothermometer – concepts, methodologyand application to basin history and hydrocarbon generation. In: Nuccio, V.N.,Barker, C.E., Sally, J., Dyson, S.J. (Eds.), Application of Thermal Maturity Studiesto Energy Exploration. Society of Economic Paleontologists and Mineralogists,pp. 1–18, Rocky Mountains section.

Rossetti, F., Aldega, L., Tecce, F., Balsamo, F., Billi, A., Brilli, M., 2011. Fluid flow withinthe damage zone of the Boccheggiano extensional fault (Larderello–Travalegeothermal field, central Italy): structures, alteration and implications for hydro-thermal mineralization in extensional settings. Geological Magazine 148 (4),558–579.

Serpen, U., 2004. Hydrogeological investigations on Balcova geothermal system inTurkey. Geothermics 33 (3), 309–335.

Schiffman, P.W., Evarts, R.C., 1984. Oxygen isotope evidence for submarine hydro-thermal alteration of the Del Puerto ophiolite, California. Earth and PlanetaryScience Letters 70, 207–220.

Srodon, J., Eberl, D.D., 1984. Illite. In: Bailey, S.W. (Ed.), Micas. Reviews in Mineralogy,vol. 13. Mineralogical Society of America, Blacksburg, VA, pp. 495–554.

Tavani, S., Storti, F., Munoz, J.A., 2010. Scaling relationships between strataboundpressure solution cleavage spacing and layer thickness in a folded carbonatemultilayer of the Northern Apennines (Italy). Journal of Structural Geology 32,278–287.

Timlin, M., 2009. Effects of stratigraphy on geothermal reservoir performance. In:AAPG Annual Convention and Exhibition, Denver, CO, USA, 7–10.06.2009.

Todaka, N., Akasaka, C., 2004. Reactive geothermal transport simulations to studythe formation mechanism of an impermeable barrier between acidic and neu-tral fluid zones in the Onikobe Geothermal Field, Japan. Journal of GeophysicalResearch – Solid Earth 109, B05209.

Todesco, M., Giordano, G., 2010. Modelling of CO2 circulation in the ColliAlbani area. In: Funiciello, R., Giordano, G. (Eds.), The Colli Albani Vol-cano, Special Publication of IAVCEI, vol. 3. The Geological Society, London,pp. 311–330.

Velde, B., 1985. Clay minerals, a physico-chemical explanation of their occurrence.Developments in Sedimentology 40, 427.

Wohletz, K.H., Heiken, G., 1992. Volcanology and Geothermal Energy. University ofCalifornia Press, Berkeley, CA, pp. 432.

Wolde Gabriel, G., Goff, F., Aronson, J., 2001. Mineralogy and K-Ar geochronologyof mixed layered illite/smectite from The Geyser coring project, California, USA.

Geothermics 30, 193–210.

Yang, K., Browneb, P.R.L., Huntingtona, J.F., Walshea, J.L., 2001. Characterising thehydrothermal alteration of the Broadlands-Ohaaki geothermal system, NewZealand, using short wave infrared spectroscopy. Journal of Volcanology andGeothermal Research 106, 53–65.