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Phreatomagmatic activity and associated hydrothermal processes in the lamproitic volcano of Cancarix (Southeast Spain).

1. Introduction

The term lamproite, which includes lamproites and kimberlites in the category of ultrapotassic rocks, was initially used by Niggli (1923) to identify high K and Mg rocks, and its use is recommended by the IUGS (Mitchell and Bergman, 1991). Lamproite magmas came from partially melted mantle at depths from 80 to 170 km (e.g. Prelevic et al., 2012), although the depth proposed for the origin of the Spanish lamproite melts is considered slightly shallower, around 55-70 km (Contini et al., 1993; Ferriz et al., 1994; Solovova et al., 1994). They are characterised chemically by high MgO and [K.sub.2]O content and enrichment in incompatible elements. Typical minerals found in lamproite rocks include forsterite, leucite, phlogopite, and iron-rich sanidine.

The Volcano of Cancarix is an exceptional example of volcanic deposits characterised by phreatomagmatic activity followed by lamproitic extrusion (Seghedi et al., 2007). One variety of lamproite, rich in enstatite, sanidine and phlogopite, is referred to as Cancalite or Cancarixite due to its occurrence at the Volcano of Cancarix (Parga-Pondal, 1935; Fuster et al., 1967), though this term is not generally accepted.

The Volcano of Cancarix exemplifies part of a group of volcanic episodes that took place in southeastern Iberia during the Late Miocene (from 9.3 to 6.9 Ma, Duggen et al., 2005; Perez-Valera et al., 2013), sometimes involving volcano edifices and/or dykes. The volcanoes of the lamproitic province of Albacete-Murcia-Almeria--with main examples in Cancarix, Vera, Fortuna, and Jumilla--were initially studied in 1889 and have since become an international reference of lamproitic rocks (Fuster and Gastesi, 1965; Fuster et al., 1967; Pellicer, 1973; Hall, 1987; Venturelli et al., 1988; Mitchell and Bergman, 1991; Contini et al., 1993; Linthout and Lustenhouwer, 1993; Salvioli-Mariani and Venturelli, 1996; Seghedi et al., 2007; Prelevic and Foley, 2007; Prelevic et al., 2008; Reolid et al., 2009, 2013; Fritschle et al., 2013). In the southern Iberian Peninsula, lamproitic bodies are associated with faults (Perez-Valera et al., 2010) and Neogene basins (Elizaga-Munoz, 1994; Kuiper et al., 2006). Spanish lamproites can be divided in two groups according to their content in SiO2 and MgO (Duggen et al., 2005). The group known as phlogopite lamproites shows relatively low SiO2 and high MgO contents; it is developed on the Iberian crust and the Mesozoic sedimentary cover of Prebetic and Subbetic affinity. The other group, called lamproite derivates because they present higher Si[O.sub.2] and lower MgO contents than the typical lamproites, have been found only on the onshore Alboran domain. The Volcano of Cancarix is representative for the phlogopite lamproites and largest in size and preservation.

This article describes in detail the volcanic lamproitic edifice and its relationships with the sedimentary host rocks, whose interaction produced spectacular phreatomagmatic deposits and hydrothermal alteration in the host rocks. Such interactions have been poorly studied in the lamproitic outcrops of southern Spain to date (e.g. Seghedi et al., 2007), in contrast with the geochemical and petrological aspects of lamproite (e.g. Fuster and Gastesi, 1965; Lopez-Ruiz and Rodriguez-Badiola, 1980; Venturelli et al., 1988; Linthout and Lustenhouwer, 1993; Salvioli-Mariani and Venturelli, 1996). The analysis of the phreatomagmatic deposits enabled us to reconstruct the genesis and evolution of the eruption. In this work we present additional informations and interpretations based on the detailed analysis of five sections from the phreatomagmatic deposits as well as textural, mineralogical and geochemical analyses of the different components of these phreatomagmatic deposits. In addition, changes in the host rock related to hydrothermal activity are also analysed.

2. Methods

Several field campaigns were carried out in the area, resulting in the elaboration of a detailed geological map (1:25000). This meant exploring and evaluating the different sides of the volcanic edifice very precisely. Five cross-sections were selected for sampling, to characterise the contact between the volcanic materials and the sedimentary host rocks. Altogether, forty-two samples--of both types--were collected. The cross-sections represent five margins of the volcanic edifice (West Quarry, Cancarix southwest, Cancarix northeast, Cancarix East and East Quarry).

Mineral composition was determined by X-ray diffractometry (XRD) using Cu-K[alpha] radiation at 40 kV and 30 mA, in a Siemens D-5000 diffractometer (Universidad de Jaen). Unoriented powders were prepared using a holder filled from the side with halite as the internal standard. Oriented aggregates were prepared by sedimentation on glass slides. Ethylene glycol (EG) and dimethyl sulfoxide (DMSO) treatments were carried out on the aggregates to corroborate the identification of smectites and kaolinite.

Following XRD results and the optical microscopy study, ten samples were prepared as carbon-coated polished thin sections for Scanning Electron Microscopy (SEM) using back-scattered electron (BSE) imaging and energy-dispersive X-ray (EDX) analysis to obtain textural and chemical information. These observations were performed using a Zeiss DSM 950 SEM (Centro de Instrumentacion Cientifica of the Universidad de Granada) equipped with an X-ray Link Analytical QX-20 energy-dispersive system (EDX). An accelerating voltage of 20 kV, with a beam current of 1-2 nA and counting time of 100 s, was used to analyse the silicates by SEM. The following compounds were used as calibration standards: albite (Na), periclase (Mg), wollastonite (Si and Ca), orthoclase (K), and synthetic [Al.sub.2][O.sub.3] (Al), [Fe.sub.2][O.sub.3] (Fe) and MnTi[O.sub.3] (Ti and Mn).

The transmission and analytical electron microscopy (TEM-AEM) studies were performed with two different microscopes: a Philips CM20 scanning transmission electron microscope (STEM) operating at 200 kV and with a La[B.sub.6] filament (Centro de Instrumentacion Cientifica of the Universidad de Granada) and a Jeol JEM 2100 STEM operating at 200 kV and with a point-to-point resolution of 2.5 [Angstrom] in the TEM mode (Universidad Complutense of Madrid). Powdered portions were prepared using C-coated formvar Cu grids, and Cu rings were attached to representative selected areas of the thin-sections. Chemical analyses were made in the STEM mode with an EDX microanalyses system. Albite, biotite, spessartine, muscovite, olivine, titanite, MnS and CaS were used as standards to derive K-factors for the transformation of intensity ratios to concentration ratios following the procedures of Cliff and Lorimer (1975).

Whole-rock analyses of five samples were carried out in the X-Ray Assai Laboratories of Lancaster (Ontario, Canada): a lamproite, the matrix of a phreatomagmatic breccia, and three marl-limestone rhythmites located at different distances from the breccia contact. X-ray fluorescence (XRF) was used for the major elements and inductively coupled plasma-mass spectrometer (ICP-MS) for the trace elements.

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3. Geological setting and materials

The Volcano of Cancarix is located in the Sierra de las Cabras, South of Albacete province (southeast Iberian Peninsula, Fig. 1A), and just 2 km West of the village of Cancarix. All this area is included in the External Prebetic (Fig. 1B), the northernmost domain of the External Zones of the Betic Cordillera. The rocks constituting the main features of the Sierra de las Cabras are Jurassic carbonates corresponding to three formations (Figs. 1C and 2): (1) Massive limestones and dolostones of the Middle Jurassic, with 60 m of well-bedded thick layers (>1 m). When dolomitisation is absent, the limestone textures are wackestones and oolitic grainstones; numerous banks show parallel lamination and the karstification is more intense. (2) Marl-limestone rhythmites of the Oxfordian-Lower Kimmeridgian, characterised by a 140 m thick alternance of marl and limestone beds (wackestone-packstones of peloids and quartz) disposed in an upward-thickening sequence. (3) Oncolitic limestones of the Middle Kimmeridgian, which comprise a well-stratified upward-thickening sequence (50 m thick), with a maximum bed thickness around 2 m. The oncolitic limestones are characterised by grainstones of oncoids (19 mm in diameter) with small colonies of corals. These are coarse grained limestones.

The Sierra de las Cabras is formed by a broad open anticline of axis approximately N080[degrees]E, and features characteristics of a local anticlinorium (Fig. 1C). The folding affects the Mesozoic sedimentary sequence reaching the upper Miocene. South and East from the volcano the folds are cut by high-angle normal faults, N70-90[degrees]E, dipping to the South. East of the volcano, the normal faults constitute a set of discrete planes that descend by the North limb, each single fault having a limited slip of a few tens of metres. West of the volcano, two faults with a cartographic strike slip of ~100 m limit the western sector of the anticlinorium, which remains free of faulting (Fig. 1C). This feature suggests that these strike-slip faults acted as lateral ramps or transfer of the aforementioned normal fault system. By contrast, normal faults South of the volcanic edifice coalesce in a broad fault damage zone that omits the Jurassic and sinks the Cretaceous sequence. This southern fault system limits Plio-Pleistocene conglomerate outcrops but cannot be followed by the scree Quaternary sediments that cover the area. The minimum throw of this fault zone, considering the Upper Jurassic sediment thickness (see below), is around 200 m. The age of the main faulting stage is Miocene and it is related with the Socovos Fault activity and the lamproites emplacement (Perez-Valera et al., 2013).

The main volcanic edifice is a cylindrical hill, 900 m in diameter and 250 m high above the surrounding valleys, that crops out in the axial zone of the anticlinorium (Figs. 1C and 3A). The most evident volcanic rocks constitute a body of massive lavas with a characteristic columnar jointing that is evident along a perimetral cliff, up to 100 m height (Fig. 3A and B). Vertical joints are pervasive, with a predominant metric spacing, but a quarry escarpment shows that joints are curved at the base (Fig. 3B). In the eastern wedge of the volcanic edifice, the columnar jointing is superimposed on a banded structure resembling horizontal stratification (Fig. 3C). In the inner parts of the volcanic edifice, spheroidal weathering creates rounded boulders of lamproite (40 cm to 4 m in diameter) (Fig. 3D). The main massive lamproite edifice is almost completely surrounded by a phreatomagmatic tuff ring of variable thickness (Reolid et al., 2013), usually < 22 m, yet greater when the host rock is the mechanically less competent marl-limestone rhythmite (Fig. 1C). The phreatomagmatic ring consistently dips 20-30[degrees] towards the centre of the volcanic edifice. The higher values of dipping are at the southwest edge. The tuff ring is better developed South and East of the edifice; elsewhere the thickness decreases, particularly to the North and southwest, where the main lava body is almost in direct contact with the Jurassic carbonates (Fig. 1C). Contact of the lamproitic body with the pyroclastic beds of the phreatomagmatic-effusive complex is planar to irregular, depending on the autobrecciation of the lava at the boundary.

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A smaller outcrop of volcanic rocks is separated from the main edifice. It is situated in the continuation of the southern normal fault system (Fig. 1C). It has a maximum length of 100 m and is made of lava flow and volcanic ash. Quaternary sediments shelter the relationship between these lavas and the host rock. This body can be interpreted as a lateral conduit of the main volcanic edifice (Reolid et al., 2013).

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4. Phreatomagmatic-effusive complex

The studied sections distributed on five margins of the volcanic edifice (West Quarry, Cancarix southwest, Cancarix northeast, Cancarix East and East Quarry; Fig. 1C) present a thickness of the phreatomagmatic-effusive complex ranging from 8 to 24 m. Three different lithologies can be differentiated in the phreatomagmatic-effusive complex (phreatomagmatic complex in Reolid et al., 2013). According to a general stratigraphic order they are:

a.--Contact breccia, consisting only of carbonate clasts derived from the immediately surrounding host rocks (mainly massive limestones) (Fig. 4). The clast size is usually from 1 to 30 cm, but exceptionally ranges from 7 m maximum diameter down to microscopic size. The clasts are mostly angular to subangular, and locally exhibitjigsaw-fit textures (Lorenz and Kurszlaukis, 2007). The smallest clasts constitute the cemented carbonate matrix. The contact breccia is associated with the West edge of the main volcanic edifice.

b.--Phreatomagmatic breccia organised in beds, generally without grading, having a metric-scale thickness and high lateral variability (Fig. 5). The breccia is composed by a grey-to-white matrix (<2 mm in size) constituted by small grains of host sedimentary rocks and volcanic rocks (from lava fragments and lapilli-size pyroclasts to coarse ash) (Fig. 6A and E). The proportions of pyroclasts and fragments of sedimentary rocks are variable. The breccia is darker when volcanic material predominates. White carbonate clasts (5-100 cm) are calcinated fragments of sedimentary rocks (Figs. 5D and 6) and present spectacular white to pink alteration rings (Figs. 6B and 6C). The blocks composed by large fragments broken from previously consolidated lava (Fig. 5) are usually angular and show different textures, from fluidal with phenocrysts of forsteritic olivine, to highly vesicular scoriaceous. Subspherical vesicular clast up to 45 cm in diameter are resembling ballistical bombs. Some large pyroclasts with fluidal texture include white carbonate clasts (Fig. 6D).

c.--Lava interlayers of metric-scale beds (<3 m) pinch out laterally a few metres (Figs. 5 and 7). Two main types of lava may be differentiated: a) massive lava (sometimes clastogenic, produced by autobrecciation) with bedding (Fig. 7A); and b) banded vesicular lavas having a fluidal laminated texture and different vesicular density, with forsterite phenocrysts (Figs. 7B-F).

Figure 8 shows the distribution of the three components of the phreatomagmatic-effusive complex along the five studied sections. The contact breccia is very well exposed only in the West Quarry section, with a maximum thickness of 17.1 m that progressively decreases toward the eastern parts of the phreatomagmatic ring. In the same sense, the thickness of the phreatomagmatic breccia and the lava interlayers increases. Following a stratigraphic sequence from the carbonate host rock at the base to the lamproite edifice, the contact breccia, if present, is at the lower part of the sections. After the contact breccia, the phreatomagmatic breccia with lava interlayers is recorded. The white carbonate clasts are dominant at the base of the phreatomagmatic breccia and decrease progressively upwards; meanwhile, pyroclasts and lava interlayers increase upwards. As a general trend, the diameter of pyroclasts increases upwards. The fluidal vesicular lava interlayers are dominant with respect to the clastogenic massive lava inter-layers, which are more common at the upper part of the East Quarry section.

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5. Mineralogical and geochemical characterisation

This section is focused on the mineralogical, geochemical and textural features of the phreatomagmatic deposits and the altered host rocks (marl-limestone rhythmite and oncolitic limestones) surrounding the lamproitic edifice as a consequence of the volcanic activity. Previously, there is a brief description of the lamproites.

5.1. Lamproites

The studied volcanic outcrops are constituted by lamproites composed by forsteritic olivine, Ti phlogopite, diopside, enstatite, richterite with Ti and K, leucite, and Fe-rich sanidine. According to Fuster et al. (1967), the mineralogical and textural features depend on the crystallisation degree. Olivine almost disappears in the holocrystalline inner parts of the lamproitic main body, but is locally common in the glassy lavas where there is a porphyric texture with glassy to microcrystalline groundmass. Phlogopite and diopside are mainly present in holocrystalline rocks as subhedral and euhedral crystals (Fig. 9), while amphibole and sanidine appear as interstitials sometimes developing poikilitic crystals. The geochemistry of these rocks is extremely unusual due to the high contents in MgO, K2O, P2O5 Ni, Cr, Rb, Ba, Th, and Zr (Table 1).

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5.2. Matrix of the phreatomagmatic breccia

The phreatomagmatic breccia XRD patterns show the presence of clay minerals (smectites and serpentine minerals) in addition to enstatite, quartz and carbonates (calcite and dolomite) (Fig. 10A). The XRD diffractogram corresponding to the matrix of the phreatomagmatic breccia (Fig. 10B) shows a clear enrichment in clay minerals and a minor presence of carbonates. Green crystals with xenomorphic and hypidiomorphic habit (Fig. 6A) were observed and analysed by SEM; they correspond to enstatite ([Si.sub.1.89] [Al.sub.0.20])[O.sub.4]([Al.sub.0.19] [Fe.sub.0.11] [Mg.sub.1.48] [Ti.sub.0.01]) [Ca.sub.0.15]. Enstatite crystals are surrounded by calcite crystals (Fig. 11A) and a very fine-grained matrix composed mainly by clays (Fig. 11B). Figure 11C shows phenocrysts of forsterite ([Fe.sub.0.16] [Mg.sub.1.82]) SiO4. The smectites of the matrix are mainly trioctahedral smectites (Mg-rich saponite-like composition, Table 2). At the lattice scale, smectites show a texture with slightly curved and discontinued packets of 10 [Angstrom] layers (Fig. 11D).

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5.3. Lava interlayers from phreatomagmatic-effussive complex

The lava interlayers present almost exclusively forsterite phenocrysts under petrographic microscope (Figs. 7E, F). However, the XRD analysis of one of these lava layers, partially altered and intercalated in the phreatomagmatic breccia, presents a mineral association of phlogopite, sanidine, leucite, amphibols, augitic piroxenes, smectite and vermiculite (Fig. IOC). TEM-AEM microanalyses of smectites reveal an Mg-rich trioctahedral saponite-like composition. The chemical composition of amphibols is very rich in K, with values around 1.5-2 atoms per formula unit (a.f.u.).

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5.4. Host rocks

The emplacement of the volcanic edifice affected the marllimestone rhythmite and oncolitic limestones, promoting in situ mineralogical changes as well as the genesis of white carbonate clasts incorporated into the phreatomagmatic breccia and xenoliths in the lamproites.

Marl-limestone rhythmite

The marl-limestone rhythmite was sampled along the same bed at different distances from the volcanic edifice in order to check potential effects on the mineralogy and geochemistry related to the volcanic emplacement in the host rock. In the outcrop, the most evident effect on the marls and limestones closer to the phreatomagmatic-effusive complex is the intense red colour (Fig. 12). Yet three metres away from the contact, the marl-limestone rhythmite has the characteristic yellow to white colour.

The XRD patterns of the marls closest to the phreatomagmatic breccia show a dominant carbonate composition, mainly dolomite and secondarily calcite. Other components are quartz and phyllosilicates, mainly mica (muscovite and biotite). BSE images (Fig. 13) show crystals of dolomite, silicates with detrital appearance (K-feldspars with rutile inclusions, quartz, tri--and dioctahedral micas, chamositic chlorites and zircons) and very fine-grained kaolinite (Table 3).

The distance to the volcanic rocks does not imply significant mineralogical differences in the closest 25 m of the marl-limestone rhythmite. In samples farther from the volcanic edifice, however, calcite is the dominant carbonate instead of dolomite. Regarding the silicates, quartz and a very minor proportion of micas are present.

The geochemical composition of the marl, according to a proximal-distal transect with respect to the lamproite edifice, is consistent with the mineralogy recorded (Table 1). The Mg content is higher in the analysed marly bed closer to the phreatomagmatic breccia, given the abundance of dolomite and the influence of the lamproite edifice as a source of Mg. The Mg content tends to decrease from the lamproite-phreatomagmatic breccia to the reddish marls-yellow marls (see Table 1). This trend is paralleled by other elements, such as Ni, Rb, Cs, Ba, U, Th, Be, Cr, Zr and Co.

Oncolitic limestones

The oncolitic limestones show textural and mineralogical changes, which vary according to the proximal-distal gradient to the volcanic edifice. In a distance >20 m the oncolitic limestones do not show evidence of alteration, but relatively closer samples present a progressive increase in small rhombohedral crystals in the original sparitic cement of the oncolitic grainstone (Fig. 14A). These new crystals, white in hand sample, are yellow cream to dark brown in thin section and grow preserving the ooids and peloids located in the sparitic cement. Their altered aspect is characterised by darker edges. EDX-SEM microanalyses of the rhombohedral crystals indicate a solid solution between calcite and dolomite with a proportion of Mg always lower than Ca. The darker parts of these crystals are enriched in Si. The opacity of the grains increases with the content of Si and suggests the record of amorphous silica. Locally, very fine-grained Mg smectites were identified.

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At a distance of 5 m from the lamproite contact the new crystals form a continuous mass among the oncoids (Fig. 14B). The smaller grains--ooids, peloids and bioclasts--are consumed by the growth of the rhombohedra, now larger and darker due to the increase in amorphous silica content. The dissolution/ crystallisation advances following the sparitic cement, while the oncoids are preserved as relicts resisting the growth of the rhombohedra due to low porosity and the presence of dense micritic laminae in concentric coats acting as a barrier. At any rate, isolated idiomorphic crystals grow inside the oncoids. The new crystals locally appear with dissolved margins.

At the lamproite-oncolitic limestone contact, all the sparitic cement and the oncoid edges were re-emplaced by a new generation of carbonate crystals (Fig. 14C). The inner parts of the oncoids show a recrystallisation by xenomorphic sparitic calcite, and only occasionally an oncoid preserving remains of the original oncolitic lamination is recorded (micritized oncoids).

White carbonate clasts of the phreatomagmatic deposits

These are calcareous clasts from the host rock located inside the phreatomagmatic breccia and locally in the lava beds (Figs. 5 and 6). These calcareous fragments present different altered features depending on the temperature conditions undergone and the nature of the original rock (massive limestones and dolostones of the Middle Jurassic, marl-limestone rhythmite of the Oxfordian-Lower Kimmeridgian or oncolitic limestones of the Middle Kimmeridgian). The porosity, water content, presence of phyllosilicates (especially in marls and marly limestones) and the size of the clasts are factors controlling the final texture. In any case, clasts with a very high content in carbonates are recrystallised, having lost the original sedimentary fabric. When the recrystallisation degree is low it is possible to identify traces of the sedimentary fabric, including microfossils. The size of the white clasts ranges from a few centimetres to exceptionally 1 m.

Many clasts from the marl-limestone rhythmite are characterised by low density and concentric bands. The most evident textural changes are alteration rings or haloes (Fig. 6B, C and 15 A, B). The microscopic analyses evidence a diffuse texture with calcite crystal balls (Fig. 15A). XRD patterns of marly limestones clasts show a composition of mainly calcite and clay minerals such as trioctahedral smectites and lizardite, a mineral of the serpentine group (Fig. 15C). SEM reveals that the crystals of calcite present Si and a Ca content lower than 1 a.f.u. (0.7-0.8 a.f.u.). It is possible to infer contamination by amorphous silica and to consider the presence of CaO in the external haloes located in the margins of the white carbonate clasts as a product of calcination. Other white clasts with concentric haloes show the same mineralogy, but the calcite crystals are not observable (Fig. 15B). White clasts coming from the oncolitic limestones of the Middle Kimmeridgian evidence the same features as at the contact between the lamproites and the oncolitic limestones on the North side of the volcanic edifice (Fig. 14C).

Xenoliths

Two main types of xenoliths, located in the West side of the volcanic edifice, can be recognised in the lamproites, according to the original composition of the host rock: massive limestones and dolostones of the Middle Jurassic (Fig. 16A, B) and marly limestones of the Lower Kimmeridgian (Fig. 16C, D). In both cases these sedimentary-derived xenoliths, ranging from 3 cm to 1m, form discrete patches of rather clear rock within the lamproite. Their clear outlines suggest that they have not been assimilated by the magma.

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The xenoliths composed originally by massive limestones or dolostones have a strongly recrystallised and brecciated appearance, featuring jigsaw-fit textures (Fig. 16A). From the compositional point of view they present enrichment in amorphous silica. The marly limestones xenoliths, some of them included in vesicular lava layers just below the lamproitic edifice, are composed of carbonates, amphibols, quartz, K-feldspar and clay minerals (smectite and serpentine minerals).

6. Interpretation

6.1. Origin and evolution of the volcano

The Volcano of Cancarix is a small edifice that implies a relatively limited emission of magma, which required the aid of a free pathway along the crust, such as a fault or fault zone. There is no chief structure with a clear expression on the surface as occurs in nearby lamproitic outcrops (Perez-Valera et al., 2013), though a volcanic-diapiric alignment has been proposed (Jerez-Mir, 1973; Rodriguez Estrella, 1979). The extrusion of the magma appears to be associated with the normal and strike-slip local fault systems that shaped the Sierra de las Cabras (Reolid et al., 2009; Perez-Valera et al., 2013), probably during the Late Miocene.

The folding is previous to the faulting stage, thus most likely unrelated with the magmatism extrusion. On the other hand, although the volcanic outcrops are not faulted, there is a close relationship between outcrop distribution and shape, and fault distribution (Fig. 1C), suggesting that the volcanic eruptions took place during or immediately after the faulting. In this sense, the [sup.40]Ar/[sup.39]Ar age of the volcanic rocks at 7.0 Ma (Duggen et al., 2005) is congruent with the supposed age of faulting during the Upper Miocene. Moreover, lamproitic dykes, with a similar composition intruded along the neighbouring Socovos Fault, where they have been interpreted contemporaneous with the fault movement (Perez-Valera et al., 2010, 2013). Therefore, fault activity could be related with magma ascent in depth determining the final point of emplacement at the surface.

The effusion of the igneous rocks took place in two clearly different phases: a) an initial highly explosive phase, dominated by phreatomagmatic eruption as a result of the interaction between the magma and the phreatic water contained in the karstic system and in the fractures of the host Jurassic calcareous rocks; and b) a second phase dominated by the emission of lava, producing the filling of the crater where the lamproitic materials are more crystalline. Consequently, massive lava were deposited over the phreatomagmatic-effusive complex (Fig. 17).

According to this phreatomagmatic eruptive model, the initial explosions increased the diameter of the central vent of the volcano, producing a progressively wider crater (Fig. 18). The subsequent explosions led to underground fragmentation of the host rock, resulting in contact or explosive breccias (Grady and Kipp, 1987; Lorenz et al., 2002; Lorenz and Kurszlaukis, 2007). The phreatomagmatic-effusive complex shows alternance between lava flows and phreatomagmatic breccia beds, which suggests an open system in the magma/ phreatic water interaction. Lava flows correspond to effusive episodes between phreatomagmatic explosive activities. Aranda-Gomez and Luhr (1996) and Risso et al. (2008) propose this type of open system with different phases and a variable water input for examples from Mexico and Argentina. The intervals dominated by massive lava (sometimes clastogenic, produced by autobrecciation) in the phreatomagmatic-effusive complex are located atop the phreatomagmatic sections and suggest a decrease in the interaction processes between magma and phreatic water, as indicated by Seghedi et al. (2007). The intervals dominated by phreatomagmatic breccia indicate new water inputs producing phreatomagmatic explosions and the incorporation of abundant sedimentary rock fragments resulting later in white carbonate clasts.

The moderate vesicularity of pyroclasts, mainly lapilli size, indicates that vesicular texture in the magma developed early during the phreatomagmatic fragmentation by magmaphreatic water interaction (Nemeth et al., 2007). The high content of white carbonate clasts in the phreatomagmatic breccia suggests a transport of pyroclasts through a relatively narrow conduit in the first phases of eruption, when the conduit was unstable and collapse episodes of the explosion chamber were more likely.

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The phreatic water inputs to magma were controlled by jointing sets of the host rocks and the karstic system. Massive limestones and dolostones were the rocks with the highest groundwater flow, due to the penetratingjointing sets and the karstic system developed in these fractures. These rocks allowed the input of groundwater to the phreatomagmatic system. Lorenz and Kurszlaukis (2007) and Kurszlaukis and Lorenz (2008) propose a general model of phreatomagmatic eruptions with some aspects compatible with the Volcano of Cancarix features. According to this model, phreatomagmatic thermohydraulic explosions initially happen close to the surface and become progressively deeper with the advance of explosive activity (Fig. 18). This process produces an increase in the depth and diameter of the crater. Successive explosions produce host rock fragmentation, resulting in contact or explosive breccia (Grady and Kipp, 1987; Lorenz et al., 2002; Lorenz and Kurszlaukis, 2007). The effective underground fragmentation of once normallyjointed and un-weathered massive limestones requires a mechanical process entailing a high stress rate. In other words, it would have required an explosion (Grady and Kipp, 1987; Lorenz and Kurszlaukis, 2007), during which the host rocks were affected first by a shock wave, then by a rarefaction wave (Fig. 18). These contact breccias may collapse downward (total or partially) on the chamber formed by the previous explosion. After each explosion, the explosion chamber (s. Lorenz and Kurszlaukis, 1997) is temporarily evacuated by the eruption, giving a cavity filled in by the wall collapse with volcanic rock fragments as well as carbonate host rock fragments (the future white carbonate clasts, Fig. 18B). New magma inputs reach the explosion chamber and press on contact breccia, which may have phreatic water arriving from the host rock fractures by means of the piezometric gradient. This process produces new evaporation of water close to the magma and the increase of pressure in the explosion chamber, favouring another phreatomagmatic explosion (Fig. 18C). Because the water vapour has high pressure and temperature during the thermohydraulic phreatomagmatic explosion, it rapidly drives the expansion of gas to the surface, moving out great amounts of clasts from the contact breccia (Lorenz et al., 2002).

[FIGURE 15 OMITTED]

When the input of outside water from the host rock decreased, the eruption became drier and evolved, according to Seghedi et al. (2007) towards Strombolian fallout episodes with intercalated clastogenetic lava flows (Hawaiian-style). The water availability decreased as eruption continued, as a consequence of the exhausting of the outer water sources from the host rock and continuous magma supply. Subsequently, the highly viscose degasified crystalline magma advanced over the phreatomagmatic-effusive complex filling the crater. This second stage of the eruption is discussed by previous authors. According to Mitchell and Bergman (1991) this second phase resulted on a lava lake, nevertheless Seghedi et al. (2007) proposed an extrusive dome formed by viscose magma. From our observations, more crystallized magma suggests high viscosity and the curvature of the columnar jointing to the base on the East edge point to extrusive lava dome as documented in other European examples (Hetenyi et al., 2012).

The phreatomagmatic ring consistently dips 20-30[degrees] towards the centre of the vent and indicates the partial collapse of the crater walls after the phreatomagmatic explosive phase. This allows us to interpret the geometry of the walls during the Strombolian eruption phase. The dipping of the walls was lower toward the East edge of the volcano, at the contact with the marl-limestone rhythmite, and coincides with the thicker phreatomagmatic-effusive complex (Fig. 8).

The volcanic vent that fed the system is situated close to the West edge. This is deduced from: a) the tectonic structure of the Sierra de las Cabras, b) the highest values of dipping of the phreatomagmatic breccia at the southwest edge, c) the highest content of contact breccia on the West edge (Fig. 8), d) the record of host rock xenoliths West of the lamproite edifice (Fig. 16), e) the absence of phreatomagmatic-effUsive complex in the northwest edge (Fig. 1C), f) the curvature of the columnarjointing to the base on the East edge, indicating the lateral palaeo-boundary of the lava flow, and the lesser thickness of the lamproitic edifice in this sector (Fig. 3B), and g) the banded structure resembling horizontal stratification observed in the East wedge, likewise indicating the lava flow (Fig. 3C). Lamproites cover the vent located close to the northwestern sector, whereas the lava flows advanced mainly to the southeast, favoured by the presence of host rocks with less strength. For instance, the marl-limestone rhythmite may create a depressed palaeo-topography compared with massive limestones and dolostones of the Middle Jurassic and oncolitic limestones of the Upper Jurassic (see the geological map of Fig. 1C). This palaeo-topography controlled the eruptive process and the thickness and dipping of the phreatomagmatic-effusive complex (thicker and having less dip over marl-limestone rhythmite).

[FIGURE 16 OMITTED]

6.2. Hydrothermal alteration processes

The first phase of the effusion of the igneous rocks--which promoted initial dehydration, cracking and hydrofracturing in the host materials--was very explosive and dominated by phreatomagmatic eruption. Afterwards, further explosions led to the advance of fragmentation, generating contact or explosive breccias. As a consequence of these processes, the surrounding carbonates and marls underwent mineralogical and textural changes. According to several authors (Dennis et al., 1982; Buhmann, 1992; Bishop and Abbot, 1995), the size of the igneous body is one factor conditioning effects on sedimentary host rocks. Other important variables are the temperature of the melt, the fracture network and texture of the country rock, the fluids and the fluid fluxes. Central volcanoes, which are fixed in space, permit much longer intervals of high-temperature hydrothermal systems than those at mid-ocean ridges, with new crust continually moving out. Another factor to take into account is the depth of emplacement. When it is close to the terrestrial surface or on the surface, cooling is faster than when the emplacement is deep, resulting in short heating times and smaller intervals for changes. In the case of the Volcano of Cancarix, as previously mentioned, the extrusion of the magma appears to be related to the normal and strike-slip local fault systems that shape the Sierra de las Cabras, which suggests rapid cooling. In any case, the chemical, textural and mineralogical data of the host rocks clearly reveal a process of recrystallisation in the carbonates, above all in the metres closest to the contact with the igneous materials, due to the increasing temperature and the circulation of hydrothermal fluids in the surroundings of the volcanic activity.

[FIGURE 17 OMITTED]

The geochemical data along the transect lamproite--phreatomagmatic breccia--marls of the host rock (marl-limestone rhythmite) point to a decrease in the contents of the most characteristic components of lamproite (MgO, [K.sub.2]O, [P.sub.2][O.sub.5], Ni, Cr, Rb, Ba, Th and Zr). The marls close to the phreatomagmatic breccia present higher values in these elements, which progressively decrease farther from the phreatomagmatic ring. The chemistry of lamproite is key to explain the high content of dolomite in the carbonate rocks surrounding the volcano, instead of calcite as it happens in the farthest host rocks. A good example of this feature is found in the oncolitic limestone, where the dolomite crystals are described as small euhedral grains (rombohedra) dispersed in the original sparitic cement, together with amorphous silica and Mgsmectites, 30 m from the volcano. Nevertheless, the proportion of these minerals increases according to the proximity to the lamproite; by the contact of the oncolitic limestone with the lamproite (North face, Fig. 1) the dolomite constitutes a continuous mass of crystals among the oncoids, which are recrystallised by xenomorphic calcite (Fig. 14). These textural and mineralogical data point to a hydrothermal alteration that favoured the development of dissolution-recrystallisation processes. These processes were more accentuated at the contact with the lamproitic body, where the temperature would have reached the highest values for the longest duration, as deduced from its crystallinity compared to the lava layers of the phreatomagmatic-effusive complex. In subaerial hydrothermal systems like this, fluids can recirculate and boiling can occur, causing precipitation of hydrothermal minerals and even the partitioning of gases such as C[O.sub.2] in the vapour phase; it may condense later near the surface, producing carbonate-rich surface groundwaters. Whereas the recent emplacement of the lamproites affected the sparitic cement of the oncolitic limestone (due to the porosity and the coarse grained texture), which shows the advance of hydrothermal fluids among grains, the micritic coating of the oncoids would be more resistant to alteration (due to the scarce porosity and fine grained texture). Finally, by the contact with the lamproites the oncoids would have recrystallised during a later superimposed hydrothermal event under cooler conditions, after the total replacement of the sparitic cement by dolomite and amorphous silica. Silicification processes are common in the later stages of the hydrothermal cycle, when temperatures fall and silica is introduced into the system. The texture of limestone beds within the rhythmite (mudstone to wackestone) was not affected by these alterations because of the abundance of micrite. Nevertheless, the texture of the white clasts of the phreatomagmatic breccia, completely recrystallised and featuring alteration haloes, evidences calcination. In the case of the xenoliths, especially the marly limestone xenoliths, immersed in the lava layers or even in the lamproitic body, mineralogy provides clear evidence of reaction processes favoured for higher temperatures, incorporating not only carbonates but also amphibols and K-feldspars.

[FIGURE 18 OMITTED]

Smectites are typical products of hydrothermal alteration (e.g. Clayton and Pearce, 2000; Abad et al., 2003; Henry et al., 2007; Jimenez-Millan et al., 2008; Miyoshi et al., 2013; Reolid and Abad, 2014). Saponitic smectites, characterised by a predominance of Ca in the interlayer (Table 2), were identified not only in the host rock but also in the white clasts, in the matrix of the phreatomagmatic breccia, and in the altered surface of lamproitic pyroclasts. Their presence suggests a smectite crystallisation caused by a process involving the circulation of Mg-rich fluids from the igneous rocks. The surrounding host carbonate rocks that were simultaneously altered could be a source of Ca. The textural features of the phreatomagmatic breccia, with high porosity and low cementation, facilitated the circulation of hydrothermal fluids. Indeed, the matrix of the phreatomagmatic breccia with <2 mm grain size and components from sedimentary and volcanic nature is where most mineralogical changes occurred. Saponitic smectites have been described in hydrothermal alterations of marls related to the intrusion of basalt dykes (Henry et al., 2007) and the contact metamorphism of limestones affected by basic sills (Kemp et al., 2005). In the Betic Cordillera, Abad et al. (2003) record saponite as the main phyllosilicate in the contact between the marl-limestone rhythmite and doleritic laccolith of Sierra de Priego (Cordoba), as a consequence of a late hydrothermal process after contact metamorphism produced by the intrusion of subvolcanic rocks. In this research, given the high Mg content characterising the lamproites, the record of saponite in the phreatomagmatic breccia and in the host rocks affected by the alteration processes was clearly promoted by the emplacement of the lamproitic body.

Finally, kaolinite and serpentine minerals were also found in the host rocks and in the matrix of the phreatomagmatic breccia. The presence of kaolinite in some samples of the marls closest to the phreatomagmatic breccia could indicate a fluid-mediated process under low temperatures (<200[degrees]C) and probably favoured by post-intrusion hydrothermal fluids. Its presence is usually considered as evidence of low-temperature alteration of other aluminosilicates, especially feldspars. With regard to the serpentine minerals, their record in the matrix of the phreatomagmatic breccia, the white carbonate clasts and the xenoliths could be the result of interactions between Si and Mg-rich fluids and dolomite, as described Deer et al. (1992). Specifically in the case of the phreatomagmatic breccia and altered lamproitic pyroclasts, the serpentine minerals are probably the result of forsterite and enstatite alteration; in turn, vermiculite in the altered lava layers of the phreatomagmatic-effusive complex would be the result of alteration of phlogopite, but probably associated with weathering processes as described by Toksoy-Koksal et al. (2001).

7. Conclusions

This study was intended to characterise the phreatomagmatic phase of the eruption of the Volcano of Cancarix and the interactions between the host rocks and the lamproites. The main results may be summarised as follows:

a.--Two episodes of material emission can be distinguished: 1) explosive volcanism due to the interaction between magma and groundwater from the karstic system of the host carbonate rocks, generating a phreatomagmatic-effusive complex (breccias and lavas); and 2) emission of crystal-rich magma responsible for the main lamproitic body.

b.--The crystalline lamproitic dome of the volcano is surrounded by a ring constituted by the phreatomagmatic-effusive complex, composed by contact breccia, phreatomagmatic breccia (with pyroclast ash, lapilli and bombs, and white carbonate clasts from the host rock) and lava interlayers (massive clastogenics and banded vesiculars).

c.--The thickness and composition of the phreatomagmatic ring is conditioned by the mechanical behaviour of the country rocks. It is minimal in contact with the massive limestone-dolostones, where contact breccias predominate (Fig. 8). In contrast, phreatomagmatic deposits dominated by pyroclastic deposits are reaching a maximum thickness towards the marl-limestone rhythmites.

d.--Superimposed hydrothermal events took place in the host rocks close to the lamproites, mainly detected by textural and mineralogical changes. Porous oncolitic limestone shows the effects of circulating hydrothermal fluids up to 20 m away from the contact, with increasing contents in dolomite, amorphous silica and saponite. Yet relatively impervious marly intervals exhibit no significant mineralogical changes except for the substitution of calcite by dolomite. Limited chemical changes do occur, with enrichment in Mg, Ni, Rb, Cs, Ba, U, Th, Be, Cr, Zr and Co present in the first 8 m from the contact.

e.--Hydrothermal alteration is pervasive in the phreatomagmatic breccia--both in white carbonate clasts and matrix of the breccia--with an abundance of saponite and lizardite.

f.--Study of the phreatomagmatic complex in five detailed sections, in view of the tectonic features of the Sierra de las Cabras, led us to situate the lava conduit below the western part of the lamproite building. In this sector, contact breccia and host rock xenoliths are more common. The banded lava indicates the advance of lava in a preferently west-east direction favoured by the depressed palaeo-topography produced by the marl-limestone rhythmites.

In summary, this work reveals that spatiotemporally limited eruptions, typical of ultrapotassic magmas, are strongly conditioned by the country rock lithology and structure. Hydrothermal processes result in a variety of low T minerals, depending on the nature of the affected rocks and the fluids from the volcano, which also impose a unique geochemical signature upon the surrounding sedimentary rocks, in turn conditioned by their permeability.

http://dx.doi.org/10.5209/rev_JIGE.2015.v41.n2.46696

Acknowledgements

This research was supported by Projects RYC-2009-04316 (Ramon y Cajal Program, Ministerio de Ciencia e Innovacion + FSE), CGL2011-30153-C02-01 (Spanish Ministry of Science and Innovation) and RNM-7408 (Junta de Andalucia), and the Research Groups RNM-325 and RNM-370 (Junta de Andalucia). We are grateful to Antonio Piedra, technician at the Laboratorio de Geologia of the Universidad de Jaen, for preparation of thin sections and polished slabs. The authors thank Jean Sanders for reviewing the grammar. We also thank the Delegacion Provincial de Medio Ambiente y Desarrollo Rural de Albacete for their authorization to work in the Monumento Natural del Piton Volcanico de Cancarix. We are also grateful to Prof. loan Seghedi (Univ. Bucharest) and an anonymous reviewer for their detailed and constructive reviews.

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M. Reolid *, I. Abad, M. Sanchez-Gomez

Departamento de Geologia and CEACT, Universidad de Jaen, Campus Las Lagunillas, 23071 Jaen, Spain.

e-mail addresses: mreolid@ujaen.es (M.R., * Corresponding author); miabad@ujaen.es (I.A.); msgomez@ujaen.es (M.S.-G.)

Received: 22 March 2014 / Accepted: 22 November 2014 / Available online: 20 July 2015
Table 1.--Chemical composition (wt % and ppm) of the studied
materials, lamproites, matrix of the phreatomagmatic breccia and marls
at different distances to the contact (0.9 m,8m and 40 m).

                                  Phreat.    Marls    Marls   Marls
                      Lamproite   breccia   (0.9 m)   (8 m)   (40 m)

Si[O.s ub.2]             38.56      24.48     16.66    32.13   12.58
[Al.sub.2][O.sub.3]     5.85       7.39      5.19     7.17     4.28
[Fe.sub.2][O.sub.3]     3.44       2.42      1.76     2.13     1.41
MnO                     0.051      0.028     0.02     0.031   0.016
MgO                     14.97      21.25     14.94    10.75    2.43
CaO                     12.1       33.49     23.89    17.76   41.46
[Na.sub.2]O             0.51        0.3      0.07     0.16     0.07
[K.sub.2]o               2.4       1.17      1.47     2.05     0.86
Ti[O.sub.2]             0.951      0.348     0.251    0.393   0.168
[P.sub.2][O.sub.5]      0.58       0.16      0.12     0.07     0.04
L.0.I.                    -        8.52      27.18     27     26.23

Sc                        8          6         4        6       4
Be                       11          3         2        2       1
V                        51         55        41       55       31
Cr                       400        50        30       30       30
Co                       22          3         2        5      < 1
Ni                       390        80        70      < 20      20
Cu                       20         30        10      < 10      20
Zn                       70         30        40       30      <30
Ga                       12         12         9        9       6
Ge                       1.2        0.9       0.9      0.7    < 0.5
As                       65          5         5       < 5     < 5
Rb                       150        94        87       100      51
Sr                       486        184       131      121     390
Y                       17.3        22       15.5     20.9     10.6
Zr                       501        110       80       118      44
Nb                      24.1        7.4       5.1      7.4     3.8
Sn                        8          3         2        3       1
Sb                       5.9        0.5       3.7      7.3     0.5
Cs                      27.3       12.1      18.5     15.5     4.6
Ba                       537        265       189      218     108
La                      61.5       26.3      17.9     23.3     12.4
Ce                       164       43.8      30.1     44.7     18.7
Pr                      23.3       5.17      3.46     5.46     2.25
Nd                      88.1        19       12.8     18.2     8.11
Sm                      19.5       4.06      2.75      3.8     1.77
Eu                      3.18       0.803     0.559    0.76    0.364
Gd                      10.1       3.39      2.39     3.52     1.52
Tb                      0.91       0.56      0.37     0.57     0.25
Dy                      3.65       3.08       2.1     3.07     1.47
Ho                      0.57       0.61      0.41     0.57     0.29
Er                      1.48       1.79      1.22     1.64     0.84
Tm                      0.209      0.26      0.182    0.242    0.12
Yb                      1.22       1.56      1.15     1.57     0.71
Lu                      0.173      0.227     0.16     0.238   0.101
Hf                      14.7        2.8        2       3.1     1.1
Ta                      2.03       0.79      0.53      0.8     0.34
W                        3.2        5.3       3.8      3.6      1
Tl                      0.85       0.43      0.67     0.79     0.26
Pb                       31          6         7        7       7
Th                      65.2       7.66      5.11     6.93     3.48
U                       12.2       3.96      2.69     2.54     2.18

L.O.I.: Loss of ignition.

Table 2.--Structural formula for smectites of the sedimentary material
normalized to [O.sub.10][(OH).sub.10].

            Si    [sup.IV]Al   [sup.VI]Al    Fe     Mg     Ti

CX-6b2_4   3.38      0.62         0.00      0.42   2.54   0.04
CX-6b2_5   3.18      0.81         0.00      0.50   2.67   0.00
CX-6b2_6   3.14      0.86         0.11      0.33   2.69   0.01
CX-6b2_7   3.16      0.84         0.12      0.34   2.54   0.03

           [summation]    K      Na     Ca    [summation]
              oct.                              inter.

CX-6b2_4      3.00       0.14   0.01   0.20      0.35
CX-6b2_5      3.17       0.01   0.00   0.25      0.27
CX-6b2_6      3.15       0.01   0.02   0.20      0.23
CX-6b2_7      3.03       0.01   0.02   0.28      0.31

EDX-SEM microanalyses

Table 3.- Structural formula for phyllosilicates in the marl-limestone
rhythmite.

                 Si    [sup.IV]Al   [sup.VI]Al    Fe     Mg     Ti

CX-8b_1 1   1   3.12      0.88         1.80      0.06   0.21   0.00
CX-8b_6 3   2   3.05      0.95         1.82      0.06   0.11   0.06
CX-8b_6 4   3   3.05      0.95         1.83      0.04   0.13   0.02
CX-8b_1 3   4   2.62      1.38         0.66      1.07   0.98   0.15
CX-8b_1 4   5   2.70      1.30         0.62      1.09   0.97   0.12
CX-8b_1 7   6   3.07      0.93         0.47      0.98   1.19   0.10
CX-8b_6 7   7   3.95      0.05         3.84      0.01   0.14   0.00

                [summation]    K      Na     Ca    [summation]
                   oct.                              inter.

CX-8b_1 1   1      2.07       0.89   0.04   0.01      0.94
CX-8b_6 3   2      2.05       0.83   0.08   0.01      0.92
CX-8b_6 4   3      2.03       0.85   0.09   0.01      0.98
CX-8b_1 3   4      2.85       0.61   0.03   0.04      0.68
CX-8b_1 4   5      2.79       0.74   0.06   0.03      0.83
CX-8b_1 7   6      2.74       0.67   0.05   0.04      0.76
CX-8b_6 7   7      4.00       0.17   0.00   0.04      0.20

1-3: muscovites, 4-6: biotites normalized to [O.sub.10][(OH).sub.2];
7: kaolinite normalized to [O.sub.10][(OH).sub.8].

                 Si    [sup.IV]Al   [sup.VI]Al    Fe     Mg     Mn

CX-8b_3 1   1   2.75      1.25         1.19      2.71   2.10   0.02
CX-8b_4 3   2   2.86      1.14         1.29      2.73   1.85   0.04

                 Ti    [summation]
                          oct.

CX-8b_3 1   1   0.00      6.03
CX-8b_4 3   2   0.01      5.92

chlorites normalized to [O.sub.10][(OH).sub.8].

EDX-SEM microanalyses
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Title Annotation:texto en ingles
Author:Reolid, M.; Abad, I.; Sanchez-Gomez, M.
Publication:Journal of Iberian Geology
Date:Jul 1, 2015
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