4.a. The geochemical imprint of the Cycladic plutons

Several geochemical and isotopic studies were made in the last two decades trying to decipher the source of several Cycladic granitoids (Altherr et al. Reference Altherr, Henjes-Kunst, Matthews, Friedrichen and Hansen1988; Robert, Foden & Varne, Reference Robert, Foden and Varne1992; Pe-Piper, Reference Pe-Piper2000; Altherr & Siebel, Reference Altherr and Siebel2002; Stouraiti et al. Reference Stouraiti, Mitropoulos, Tarney, Barreiro, McGrath and Baltatzis2010). The principal geochemical contrast is between large plutons close to the MCL and the small intrusions far to the east on the East Aegean–Anatolian block (i.e. Samos, Kos and Bodrum). These small intrusions of the eastern Cyclades appear to have a largely mantle derivation, are consistently enriched in K, Sr and Ba, and have primitive Nd and Sr isotopes (e.g. Robert, Foden & Varne, Reference Robert, Foden and Varne1992). High-field-strength element (HFSE)-enriched mafic enclaves from the central and western Cyclades have higher Nb and lower Sr and Ba than the mafic rocks in the east (Pe-Piper, Piper & Matarangas, Reference Pe-Piper, Piper and Matarangas2002). In the Pliocene–Quaternary South Aegean Arc, volcanic rocks in the east (Nisyros, Santorini) have more primitive Nd isotopes and higher Sr and Ba than volcanic rocks in the west (Pe-Piper, Piper & Matarangas, Reference Pe-Piper, Piper and Matarangas2002; Altherr & Siebel, Reference Altherr and Siebel2002).

Previously proposed petrogenetic models for the formation of the Miocene Cyclades I-type granites have been formulated mostly in terms of two-component mixing, either, between mantle-derived magmas (arc-type basalt) and/or mantle-derived igneous rock (i.e. obducted ophiolites) and upper crust end-members (Juteau, Michard & Albarede, Reference Juteau, Michard and Albarede1986; Altherr et al. Reference Altherr, Henjes-Kunst, Matthews, Friedrichen and Hansen1988). Recently, Stouraiti et al. (Reference Stouraiti, Mitropoulos, Tarney, Barreiro, McGrath and Baltatzis2010) investigated the involvement of the different basement rocks of the Attic–Cycladic Massif in the generation of the I- and S-type plutons. Initial isotopic compositions of all the granitoids are typical of crust-derived magmas from heterogeneous metasedimentary sources (I-type: ⁸⁷Sr/⁸⁶Sr = 0.7091–0.712, ε_Nd = −6.4 to −10.4; S-type:⁸⁷Sr/⁸⁶Sr = 0.710–0715, ε_Nd = −7.5 to −10.1). They also argued that compared with the isotopic Sr–Nd–O isotopic composition of the contemporaneous extrusive and intrusive rocks in the eastern Aegean, it would appear that a contribution from a high-K isotopically enriched mantle-derived magma source is not required. Their results coincide with the previously proposed suggestion that the I-type intrusions were possibly generated by dehydration melting of metaluminous older crustal sources, e.g. of metagreywacke type (Altherr & Siebel, Reference Altherr and Siebel2002).

4.b. Naxos pluton

The Miocene I-type hornblende–biotite granite of Naxos, in the western part of the island, crops out against the metamorphic basement or the Miocene molassic sediments along its border (Fig. 5b; Altherr et al. Reference Altherr, Kreuzer, Wendt, Lenz, Wagner, Keller, Harre and Hohndorf1982; Pe-Piper, Reference Pe-Piper2000; Keay, Lister & Buick, Reference Keay, Lister and Buick2001). In the north part of the island, small peraluminous S-type leucogranites are also exposed. The metamorphic basement rocks make up most of the eastern part of the island and are composed of meta-pelite, marbles and metavolcanic rocks (Durr et al. Reference Durr, Altherr, Keller, Okrusch, Seidel, Closs, Roeder and Schmidt1978). The migmatitic core, in the central part of the island, is the lowermost structural unit and comprises gneiss and dismembered marble, amphibolites and metapelite slices (Keay, Lister & Buick, Reference Keay, Lister and Buick2001).

Figure 5. Geological map of Naxos Island (modified from Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003) with fault-slip data from the major faults. NF – Naxos Fault, SF – Stellida Fault, GF – Glinadon Fault. (a) The Stellida area in more detail showing the NE-striking main Stellida Fault and a secondary NW-striking left lateral fault bounding the molasse deposits. (b) The Naxos pluton in the western part of island. Radiometric ages shown on map are from Bolhar, Ring & Allen (Reference Bolhar, Ring and Allen2010). Section 1 is shown on map, as well as the approximate location of the stereonets (1–4) sch. – schists. (c) Magmatic foliation trajectories for the Naxos pluton (modified from Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003). (d) Flow path trajectories for the magmatic stage of the Naxos pluton (modified from Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003).

Geochronological data for the I-type granite of Naxos provide ages of 12.2 Ma (Ar–Ar, hornblende; Henjes-Kunst et al. Reference Henjes-Kunst, Altherr, Kreuzer and Hansen1988) and 12.4–11.3 Ma (SHRIMP U–Pb, zircon; Keay, Lister & Buick, Reference Keay, Lister and Buick2001) for the emplacement and early cooling, while U–Pb ages range from 15.4–12.2 Ma for the S-type leucogranites (Keay, Lister & Buick, Reference Keay, Lister and Buick2001). Recently, Brichau et al. (Reference Brichau, Ring, Carter, Ketcham, Brunel and Stockli2006) reported 11.8 ± 0.8 Ma and 9.7 ± 0.8 Ma zircon fission track (ZFT) ages for the Naxos pluton, while U–Pb ages calculated by Bolhar, Ring & Allen (Reference Bolhar, Ring and Allen2010) range between 13 and 12 Ma, for the I-type, and 14.97 ± 0.32 Ma for the S-type leucogranite (Fig. 5b). Previous work on Naxos suggests a tectonic evolution that involves long-lived crustal extension (from ~ 23 Ma until ~ 8 Ma; Gautier, Brun & Jolivet, Reference Gautier, Brun and Jolivet1993; Jolivet, Reference Jolivet2001; Keay, Lister & Buick, Reference Keay, Lister and Buick2001) and formation of ductile shear zones that initially resulted in the partial exhumation of the migmatitic core in Naxos. Final exhumation involved formation of a generation of detachment faults that began to operate at ~ 10 Ma (Keay, Lister & Buick, Reference Keay, Lister and Buick2001). Others suggest that net shortening that operated on a crustal scale shaped the Cyclades islands and surroundings, while extension played an important role but later, around late-Middle Miocene time (Boronkay & Doutsos, Reference Boronkay and Doutsos1994; Avigad, Ziv & Garfunkel, Reference Avigad, Ziv and Garfunkel2001; Xypolias et al. Reference Xypolias, Iliopoulos, Chatzaras and Kokkalas2012).

The contact with the metamorphic basement is marked by a NNE–SSW-trending syn-magmatic shear zone, the Naxos Fault (NF, Fig. 5b; Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003). Mafic enclaves, which are common near the margin of the pluton, show foliation oblique to the solid-state foliation. Fault zone rocks within the central segment of the Naxos Fault are variably deformed, displaying high-to-low temperature solid-state fabrics (i.e. plastic deformation, recrystallization, aplite veins, etc.). All the magmatic structures are cross-cut by a younger mylonitic steeply dipping to vertical foliation, sub-parallel to the pluton margin. Strain ratio estimates from deformed mafic enclaves reveal an increasing ratio of simple shear to pure shear towards the pluton's eastern margin, while the x-axes of strain ellipsoids trend sub-parallel to the Naxos Fault (Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003). Mylonitic zones strike NNE and are intruded by pegmatitic and aplitic dykes. Quartz veins and pegmatitic dykes are closely spaced trending NE to ENE and display slickenlines indicating strike-slip motion during or shortly after their formation. Kinematic analysis and geological evidence (Fig. 5b, cross-section 1) suggest that the Naxos Fault Zone has a prominent dextral horizontal movement in the north and a significant reverse component in the south in addition to the dextral movement (Fig. 5b, stereonet 1). Thus, transpressional kinematics apparently control the eastern margin of the Naxos pluton (Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003).

The deformation within the pluton is characterized by ENE-trending and WNW- to NW-trending sets of steeply dipping or almost vertical faults that show evidence of right-lateral and left-lateral sense of shear, respectively (Fig. 6b, d, e). In some cases, parallel ENE-trending right-lateral strike-slip faults deform the pluton, while in the overlap zone, branching splays are formed (Fig. 6a). Furthermore, S-C type foliation formed in the mylonitic zone bordering the Naxos granite shows a similar right-lateral sense of shear, as is indicated by the offset of a K-feldspar crystal (Fig. 6c).

Figure 6. (a) Plan view of paired ENE-trending right-lateral strike-slip faults that form branching splays in their overlap zone (Naxos granite). Geological hammer for scale is 32 cm long. (b) Cataclastic zone (lower part of photo) with left-lateral slip accompanied by splays of the same kinematics displacing a dyke in the Naxos granite. Geological hammer for scale is 32 cm long. (c) S-C type foliation in the mylonitic zone bordering the Naxos granite with dextral sense of motion as is indicated by the offset of a K-feldspar crystal. Coin for scale is 2.5 cm diameter. (d) Curved foliation traces indicating right-lateral sense of slip in the Stellida Fault (SF) zone. The lower margin of the fault zone is marked by a 4 cm thick mylonitic (my) zone. Photograph looking down; width of photograph is 25 cm.

(e) Dextral shearing and cataclastic flow in the solid-state stage of the Naxos pluton. Diameter of lens cap is 6 cm. (f) Aplitic dyke in the host rock of the granite along the northwestern contact zone of the Tinos pluton. The dyke forms along a dextral mylonitic zone, which gradually curves the foliation (lower left part of photo). Diameter of lens cap is 6 cm. (g) Mesoscopic characteristics of the brittle-ductile fault zone in the Panormos area (Mykonos), comprising several anastomosing faults and fractures, trending sub-parallel to the main fault zone (PF – Panormos Fault). In the more densely fractured western part, sinistral sense of shear is observed. (h) View of NW-trending left-lateral Panormos Fault (PF) in the Mykonos pluton, placing granitic rocks (gr) over the Miocene molassic sediments (Mioc.).

The NE-trending Stellida Fault (SF; Fig. 5a, b) is a moderate-dipping normal fault that runs along the northwestern pluton rim. The fault zone, which is ~ 300 m wide, comprises 10 cm thick mylonites, a composite 20 m thick damage zone displaying several 1.5 cm thick pseudotachylites and a 15 m thick cataclastic zone (Fig. 5a). The Miocene molassic sediments, north of the fault trace, are also affected by NW- and ENE-trending faults with transtensional kinematics (Fig. 5a, stereonet 2), indicating that these faults were operated during pluton emplacement or even earlier, since molassic sediments have an Early Miocene age (Boger, Reference Boger1983). Kinematics of the Stellida Fault, based on folded aplite veins, fault-slip data (Fig. 5, stereonet 3), curved magmatic foliations (Fig. 5c) and offset of mafic enclaves and solid-state foliations, suggest a predominant dextral slip (Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003). Furthermore, NW-trending faults showing left-lateral strike-slip kinematics are also associated with the Stellida Fault Zone (Fig. 5a, b and stereonet 3). From all the above observations it seems that the Stellida Fault was initiated during the solid-state deformation stage.

In the central part of the pluton, the NE-trending Glinadon Fault (GF; Fig. 5b) is 200 m wide and divides the pluton into two structural domains: an onion-skin foliation pattern in the north and a tangential to the rim pattern in the south (Fig. 5c; Koukouvelas & Kokkalas, Reference Koukouvelas and Kokkalas2003). The pluton's magmatic lineation pattern comprises measurements of amphibole and biotite, which are interpreted as pre-full crystallization markers. These lineations abruptly change their trend from NW into N–S on either side of the Glinadon Fault, while their plunges are consistently directed towards the fault trace (Fig. 5d). The deflection of the magmatic foliations and flow path trajectories in map view along this steeply to almost vertically dipping fault, as well as fault-slip data along this wide zone, indicate a dextral shear along the Glinadon Fault (Fig. 5c, d; stereonet 4).

Summarizing, the syn-magmatic and solid-state fabrics within the main body of the Naxos pluton and its rim provide additional information about the interplay between plutonism and regional deformation at the upper crustal level. The parallelism of syn-magmatic and solid-state foliation, as well as the flow path trajectories, with the bordering Naxos Fault calls for the syntectonic emplacement of the pluton (Fig. 5c, d). Additionally, the different foliation patterns on either side of the Glinadon Fault, as well as the deflection of fabric orientations along the steeply dipping fault trace, support the dextral horizontal motions along it (Fig. 5c, d). The deepest part of the pluton, displaying steeply dipping foliation and magmatic lineation, is along the central and southern part of the Naxos Fault, as well as along the eastern side of the Glinadon Fault, close to their intersection. These areas can be interpreted as possible marginal feeder zones from which magma upwelled.

4.c. Tinos pluton

The tectonostratigraphic framework of Tinos comprises three main units: (i) The greenschist, locally amphibolitic, facies Upper unit is built up by a disrupted meta-ophiolite sequence consisting of serpentinites, meta-gabbros, ophicalcites and phyllitic rocks (Melidonis, Reference Melidonis1980). (ii) The underlying Blueschist unit, separated from the Upper unit by a shallow-dipping, ductile extensional shear zone (Gautier & Brun, Reference Gautier and Brun1994; Jolivet & Patriat, Reference Jolivet, Patriat, Durand, Jolivet, Horvath and Seranne1999), consists of marbles, schists and metavolcanic rocks (Melidonis, Reference Melidonis1980). Relics of the earlier HP-stage are preserved in many places (Bröcker et al. Reference Bröcker, Kreuzer, Matthews and Okrusch1993). (iii) The lowermost Basal unit is exposed only in NW Tinos and consists of calcite-rich marbles that are underlain by a dolomite sequence (Avigad & Garfunkel, 1989). Recently, Bröcker & Franz (Reference Bröcker and Franz2005) questioned the existence of the Basal unit on Tinos and argued that this unit is an integral part of the overlying Blueschist unit, as originally suggested by Melidonis (Reference Melidonis1980).

In the eastern part of Tinos Island (Figs 2, 7a), the Upper and greenschist–blueschist units were affected by contact metamorphism caused by an I-type Miocene granite (Bröcker & Franz, Reference Bröcker and Franz2000; ZFT age range 14.4–12.2 Ma; Brichau et al. Reference Brichau, Ring, Carter, Monie, Stockli and Brunel2007) that intruded the metamorphic basement (Fig. 7a). A small occurrence of S-type granite, at the southwestern end of the pluton, yielded U–Pb zircon ages of 14.4 ± 0.2 Ma (S. Keay, unpub. Ph.D. thesis, Australian National Univ., Canberra, 1998). The granodiorite intrusion cuts through the detachment between the Blueschist unit and Upper unit, which is considered to be extensional (Jolivet & Patriat, Reference Jolivet, Patriat, Durand, Jolivet, Horvath and Seranne1999; Mehl, Jolivet & Lacombe, Reference Mehl, Jolivet and Lacombe2005).

Figure 7. (a) Geological map of eastern Tinos showing the pluton's structural characteristics and contact aureole. Fault slip data are also shown. Position of inset (b) is shown as a white rectangle on the northwestern border of the pluton. Map modified after Melidonis (Reference Melidonis1980) and Faure, Bonneau & Pons (Reference Faure, Bonneau and Pons1991). (b) Schematic map view illustration showing the faulting configuration in the pluton border zone comprising paired dextral slip surfaces associated with cataclasite and pseudotachylite zones. (c) and (d) show cross-sections from the pluton's northwestern and southeastern borders, showing the intrusion of aplitic dykes and pseudotachylites along the strike-slip faults.

The exposed pluton body has an almost half-oval shape with a long axis trending NE–SW parallel to the foliation trajectories of the wall rocks (Fig. 7a). A weak magmatic foliation and lineation is observed, both having a NE–SW orientation. The solid-state foliation in the granite trends parallel to the magmatic foliation. The NW margin of the intrusion comprises a 1 km wide, NE-trending high-strain zone of mylonites and pseudotachylites, which possibly acted as a conduit for pegmatitic and aplitic fluids and possibly for the granitic magma. At this contact zone, ‘flame-type’ injections, curved foliation traces and apophyses of granite in the metasediments indicate magmatic intrusion during left-lateral shear. Similar results can be drawn from meso- and microscopic structures in the granite border zone, such as imbrication of K-feldspar crystals, asymmetric σ-clasts, and S-C structures (K. Boronkay, unpub. Ph.D. thesis, Univ. Patras, 1993).

Additionally, along the NW border zone, pseudotachylites are observed, which cut and displace the solid-state foliation planes (Fig. 7c, cross-section A–A′). Pseudotachylites accompany cataclasite zones (Fig. 7b) and both are confined at fault intersections, in Riedel shears, close to shear lenses or normal to the shear planes forming injection veins (Fig. 7b). This concomitance suggests a formation under an abrasive wear mechanism at a depth of 6–10 km (Sibson, Reference Sibson1983; Swanson, Reference Swanson1992). The pseudotachylites are closely associated with faults that formed under a strike-slip stress regime (Fig. 7a, stereonet 1). Furthermore, the dyke swarm that intruded the host rock at the NW border zone has a NE–SW orientation, parallel to the foliation of the border zone, and a NNW–SSE orientation inside the granitic body (Fig. 7a, stereonet 3). The dyke swarm shows a complex geometry comprising flow apophyses patterns and bridges, indicating that dilational deformation was accompanied by shearing during the intrusion of the magmatic fluids (Fig. 6f; see also Nicholson & Pollard, Reference Nicholson and Pollard1985). This is clearly also evidenced by the sub-horizontal lineations along the aplitic dykes and locally within mylonitic zones that are parallel to the dykes (Fig. 6f).

The southeastern border zone is characterized by granitic apophyses along NNW–SSE-trending left-lateral strike-slip faults (Fig. 7a, d, cross-section B–B′). Stress analysis of fault-slip data suggests an oblique-tensional stress regime with a NW orientation of σ₁ and NE orientation of σ₃ that was operating during Late Miocene time (Fig. 7a, stereonet 2)

4.d. Mykonos and Delos plutons

The Mykonos I-type monzogranite crops out over most of the island (Fig. 8) and forms a strongly asymmetric laccolith-like intrusion with a N70°E long axis, having an outlying root zone to the SW (cropping out on Delos and Rhenia islands; Fig. 2) and a major body mainly developed to the NE cropping out on Mykonos Island (Denèle et al. Reference Denèle, Lecomte, Jolivet, Lacombe, Labrousse, Huet and Le Pourhiet2011). The pluton intruded into micaschists at the top of migmatitic gneisses belonging to the basement (Faure, Bonneau & Pons, Reference Faure, Bonneau and Pons1991) that crops out only on the islands of Rhenia, Delos and in the westernmost part of Mykonos (Fig. 8; Lecomte et al. Reference Lecomte, Jolivet, Lacombe, Denèle, Labrousse and Le Pourhiet2010; for Rhenia Island see location on Fig 2). Lee & Lister (Reference Lee and Lister1992) and Avigad, Baer & Heimann (Reference Avigad, Baer and Heimann1998) suggested that the Mykonos monzogranite is topped by a low-angle (~ 30°) normal fault system, exposed in the northern part of the island. Lecomte et al. (Reference Lecomte, Jolivet, Lacombe, Denèle, Labrousse and Le Pourhiet2010) showed that this detachment is a two-branch detachment system comprising a lower branch (correlated to the detachment also cropping out in Tinos), corresponding to a low-angle ductile shear zone that reworked the intrusive contact of the granite with the Upper unit, and an upper branch (Mykonos detachment), which represents a brittle low-angle normal fault. The hanging wall comprises rare blocks of Permo-Triassic limestones and relics of the Upper unit, composed of Upper Miocene conglomerates and sandstone (Sanchez-Gomez, Avigad & Heiman, Reference Sanchez-Gomez, Avigad and Heiman2002). The footwall granite exhibits thin mylonitic zones, which are overprinted by brecciation and cataclasis close to the contact (Lee & Lister, Reference Lee and Lister1992).

Figure 8. (a) Simplified geological map of Mykonos and Delos islands (modified from Alther et al. 1982; Faure & Bonneau, Reference Faure and Bonneau1988; Faure, Bonneau & Pons, Reference Faure, Bonneau and Pons1991; Pe-Piper, Piper & Matarangas, Reference Pe-Piper, Piper and Matarangas2002; Lecomte et al. Reference Lecomte, Jolivet, Lacombe, Denèle, Labrousse and Le Pourhiet2010) with two cross-sections (marked A–A′ on inset (a) and B–B′ on map). Stereonet plots of fault planes with slickenlines and calculated stress tensor for the brittle-ductile B/D (strike-slip stress regime) and brittle B (extensional stress regime) stages of deformation are also shown. Inset (a) shows the Panormos Fault (PF) in more detail.

Granite emplacement and the tectonically controlled cooling are of a similar age, suggesting a close relationship between plutonism and faulting (U–Pb zircon ages of ~ 13.5 Ma and ZFT ages of ~ 13 Ma, Brichau et al. Reference Brichau, Ring, Carter, Bolhar, Stockli and Brunel2008). Palaeomagnetic studies suggest a ~ 22° clockwise rotation about a vertical axis (Avigad, Baer & Heimann, Reference Avigad, Baer and Heimann1998), implying a contrasting sense of rotation with the neighbouring Naxos pluton, since the emplacement of the granitoids. The latter authors also demonstrated that the (now) low-angle fault probably had a steeper dip (~ 54°) due to tilting before the vertical axis rotation occurred.

The most prominent structural feature in the granite is a NW-trending low-angle solid-state foliation that dips to the NE, while kinematic indicators, such as S-C structures and asymmetric K-feldspar crystals, suggest a NE-directed motion similar to that in the Naxos pluton (Faure, Bonneau & Pons, Reference Faure, Bonneau and Pons1991). Mylonitic foliation of the Mykonos granite is cut by wide NNW-trending oblique-normal and sinistral strike-slip fault zones (Fig. 8, inset (a) and stereonet 1), which probably developed under brittle-ductile conditions based on the low-temperature mineral assemblages (chlorite + epidote) found within the cataclasites. NW-trending, intermediate-dipping, mylonite zones up to 3 m thick are also present, and are displaced by the NW-trending steeply dipping faults. These NW-trending faults affect a zone of about 1.5 km in width and several kilometres in length going through the entire island (Fig. 8). The eastern side of this zone is marked by the NW-trending Panormos Fault (Fig. 8; PF and Fig. 6g, h).

Mesoscopic-scale characteristics of the brittle-ductile fault zone show an anastomosing fracture network oriented parallel, as well as at a low angle, to the main NW-trending fault zone (Fig. 6g). Displacements along mesoscale fractures, inside this zone and in the molassic deposits above, display sub-horizontal slickenlines on the fault surfaces, indicating a dominantly sinistral sense of motion (Fig. 8, cross-section B–B′ and inset (a), cross-section A–A′; Boronkay & Doutsos, Reference Boronkay and Doutsos1994). All these structures are cross-cut and displaced by normal faults with NW–SE to N–S orientation, forming thin cataclastic zones along them, and are formed under an extensional stress regime with an ENE–WSW tension direction (Fig. 8, stereonet 2). The stress analysis of the mesoscale faults shows that the local stress regime changed progressively from transtensional strike-slip (brittle-ductile) to pure extensional (brittle) in Late Miocene time, leading to increased dilation along the main NW fault trend (Fig. 8, stereonets 1, 2).

The neighbouring island of Delos is built up of Miocene granitoid rocks and some minor exposures of the host metamorphic basement (Fig. 8; Pe-Piper & Piper, Reference Pe-Piper and Piper2002). The characteristics of deformation include a magmatic foliation and a NE-plunging lineation, but the solid-state foliation, observed on Mykonos Island, is not evident here. Some major sub-vertical shear zones extend across the island, trending almost ENE–WSW similar to the screens of the country rocks. Sparse observations suggest a dextral movement along these zones. These zones often delimit different types of igneous rocks and appear to control magma supply (Pe-Piper, Piper & Matarangas, Reference Pe-Piper, Piper and Matarangas2002). The emplacement of plutonic rocks appears to occur with the progressive widening of space between the country rocks caused by lateral translation in shear zones, at their intersection with a detachment system, providing a pathway to the surface for magma. This is also evidenced by the steeply dipping sheets of country rocks flanked by igneous rocks (Pe-Piper, Piper & Matarangas, Reference Pe-Piper, Piper and Matarangas2002). Late brittle faults, striking almost N–S, displace the previously mentioned shear zones and show left-lateral motion along them (Fig. 8).

4.e. Ikaria pluton

On Ikaria Island, three tectonic units can be distinguished, which from bottom to top are: (i) The basal Ikaria unit, which comprises a series of metapelitic gneisses, as well as minor quartzites, amphibolites, marbles and metapegmatite (Kumerics et al. Reference Kumerics, Ring, Brichau, Glodny and Monie2005; Ring, Reference Ring, Lister, Forster and Ring2007). The Ikaria unit shows no HP-metamorphism and was correlated with the Menderes Bozdag nappe of western Turkey (Kumerics et al. Reference Kumerics, Ring, Brichau, Glodny and Monie2005; Ring, Reference Ring, Lister, Forster and Ring2007). (ii) The Messaria unit consists of greenschist-facies marbles, phyllites and calc-mica schists, and is interpreted as a Mesozoic platform succession (Altherr et al. Reference Altherr, Kreuzer, Wendt, Lenz, Wagner, Keller, Harre and Hohndorf1982; Pe-Piper & Photiades, Reference Pe-Piper and Photiades2006); (iii) The Upper unit in the central part of the island consists of a large klippe of marbles overlying a tectonic ophiolitic mélange (Altherr et al. Reference Altherr, Kreuzer, Lenz, Wendt, Harre and Durr1994). Coarse-grained terrestrial conglomerates (possibly Upper Miocene molasse; Photiadis, Reference Photiades2002) and transgressive lower Pliocene littoral sediments, which unconformably overlie the molasse, complete the stratigraphic succession. According to Kumerics et al. (Reference Kumerics, Ring, Brichau, Glodny and Monie2005), the Messaria and Upper units are parts of the passive-margin sequence of the Cycladic Blueschist unit and are separated, from the Ikaria nappe and from each other, by extensional detachments (Messaria and Fanari detachments).

The lowermost Ikaria unit/nappe was intruded by three granites (Fig. 9): (i) the large I-type Raches granodiorite intrusion (U–Pb ages ~ 15–13 Ma, Bolhar, Ring & Allen, Reference Bolhar, Ring and Allen2010), exposed in the western half of Ikaria Island, (ii) the S-type Karkinagrion granite (U–Pb ages ~ 17–13.5 Ma; Bolhar, Ring & Allen, Reference Bolhar, Ring and Allen2010), occurring within the Raches granite on the southwest coast of Ikaria, and (iii) the S-type Xylosirtis granite (U–Pb ages ~ 17–14 Ma; Bolhar, Ring & Allen, Reference Bolhar, Ring and Allen2010), exposed in the eastern part of the island.

Figure 9. (a) Structural map of Ikaria Island with cross-sections of the Raches granite (A–A′) and Xylosirtis granite (B–B′). Stereonet plots of quartz veins with stretching lineations (1), faults with slickenlines (2) and aplitic dyke orientations (3) are also provided. Note that fault-bounded intrusion bodies and aplitic dykes along faults are sub-parallel to the major faults. IF – Ikaria Fault, RG – Raches granite, KG – Karkinagrion granite (Ring, Reference Ring, Lister, Forster and Ring2007), XG – Xylosirtis granite. (Map modified from K. Boronkay, unpub. Ph.D. thesis, Univ. Patras, 1993; Photiades, Reference Photiades2002; Ring, Reference Ring, Lister, Forster and Ring2007.)

The lower part of the Raches granodiorite, exposed only on the southern coast, is weakly deformed and displays a roughly N–S-trending magmatic foliation. In contrast, the upper part of the pluton shows a pervasive low-angle (10–40°) NW-dipping solid-state foliation, containing a NNE-trending stretching lineation with a top-to-the-N sense of shear (Boronkay & Doutsos, Reference Boronkay and Doutsos1994). The eastern margin of the Raches granodiorite (RG) is controlled by the N–S-striking Ikaria Fault (IF; Papanikolaou, Sakellariou & Leventis, Reference Papanikolaou, Sakellariou and Leventis1991; Boronkay & Doutsos, Reference Boronkay and Doutsos1994). Higher strain is concentrated within a 3 km wide zone of ductile deformation parallel to the fault, where the strongly foliated granodiorite comprises mylonitic zones, folds and S-C structures (Faure, Bonneau & Pons, Reference Faure, Bonneau and Pons1991). Inside the mylonitic zone two lens-shaped tectonic slices of basement rocks are observed (Fig. 9, map and cross-section A–A′). Kinematic analysis on these structures shows a top-to-the-NNE sense of shear, while ductile WNW–ESE extensional shear bands display the same shear sense. Most aplitic dykes and quartz veins strike parallel to the solid-state foliation of the granodiorite and display horizontal lineations, probably formed during the later stages of intrusion (Fig. 9, stereonet 1). The wall rocks in the contact zone are strongly deformed by chevron and kink folds with axial planes parallel to the fault zone. In the same place, aplitic dykes are affected by a system of W-dipping shear planes forming S-C structures, which show top-to-the-NNE sense of shear (Fig. 9, cross-section A–A′ inset). Fault-slip data in the granodiorite body suggest deformation under a strike-slip/contractional regime (Fig. 9, stereonet 2). Based on all these data, it appears that the Ikaria Fault is an important structure that operated as a right-lateral, oblique thrust (Boronkay & Doutsos, Reference Boronkay and Doutsos1994; Photiades, Reference Photiades2002) during, and in the late stages, of granite emplacement.

Furthermore, in the eastern part of the island, the small S-type Xylosirtis granite (XG) is almost undeformed, with an absence of solid-state deformation, displaying only a weak NNE-striking magmatic foliation and NE–SW-trending lineation (Fig. 9). The pluton body is cut by a series of NE-striking, steeply dipping, right-lateral strike-slip faults (Fig. 9, cross-section B–B′). Aplitic dykes (0.3–5 m thick), observed on either sides of the pluton, strike N–NE, parallel to the fault planes, and have sub-horizontal lineations suggesting a transtensional stress regime (Fig. 9, cross-section B–B′, stereonet 3).

5.a. Milos volcanic centre

This volcanic centre comprises Milos, Kimolos and Polyegos islands, which are part of a series of compound volcanoes (Fig. 10; Fytikas et al. Reference Fytikas, Innocenti, Kolios, Manetti, Mazzuoli, Poli, Rita and Villari1986; Stewart & McPhie, Reference Stewart and McPhie2003; Francalanci, Fytikas & Vougioukalakis, Reference Francalanci, Fytikas and Vougioukalakis2003; Francalanci et al. Reference Francalanci, Vougioukalakis, Perini, Manetti, Fytikas and Vougioukalakis2005). Volcanic activity in the area started about 3.5 Ma ago and continued up to recent times with hydrothermal explosions (Fytikas et al. Reference Fytikas, Innocenti, Kolios, Manetti, Mazzuoli, Poli, Rita and Villari1986; Francalanci et al. Reference Francalanci, Vougioukalakis, Perini, Manetti, Fytikas and Vougioukalakis2005). The period between 3.5 Ma and 1.6 Ma was characterized by alternations of explosive and extrusive activity building most of the present islands in the region (Fytikas et al. Reference Fytikas, Innocenti, Kolios, Manetti, Mazzuoli, Poli, Rita and Villari1986; Stewart & McPhie, Reference Stewart and McPhie2003). The Milos volcanic succession, which has a maximum thickness of 700 m, is composed of calc-alkaline volcanics, mainly of rhyolitic and dacitic compositions (Fytikas et al. Reference Fytikas, Innocenti, Kolios, Manetti, Mazzuoli, Poli, Rita and Villari1986; Francalanci, Fytikas & Vougioukalakis, Reference Francalanci, Fytikas and Vougioukalakis1994, Reference Francalanci, Fytikas and Vougioukalakis2003). The succession also includes some sedimentary rocks, the stratigraphic record of which shows an upward progression from submarine to subaerial depositional environments (Stewart & McPhie, Reference Stewart and McPhie2006).

Figure 10. Tectonic map of the Milos volcanic centres (modified from Fytikas et al. Reference Fytikas, Innocenti, Kolios, Manetti, Mazzuoli, Poli, Rita and Villari1986; Stewart & McPhie, Reference Stewart and McPhie2006; radiometric ages are compiled from the same authors). Small grey circles along the Gulf of Milos represent aftershocks of the 20 March 1992 earthquake event (epicentre of major event is marked with the beach ball). Z.G. – Zephyria graben, ACHF – Achivadolimni Fault, FF – Filacopi Fault, KF – Kontaro Fault, KLF – Kleftiko Fault. Stereonet 1 shows fault slip data of normal faults mainly from the south-central part of the island (broader area of Fyriplaka, see 1 on map).

The major fault strikes identified in the islands are (see also Stewart & McPhie, Reference Stewart and McPhie2006):

(i) ENE-striking normal faults, which also seem to control the uplifted area in the SE part of Milos Island (Fig. 10, KLF and stereonet 1), exposing the metamorphic basement, as well as Neogene sediments. This fault strike also cuts the underlying molassic sediments. The ENE-striking faults show displacements in the order of 20–80 m and form cataclastic zones that reach a thickness of 1.5 m. These faults operated mainly during the late Pliocene period and seem to control many of the dacitic domes and lavas (Fig. 10), while their activity ceased gradually close to the Pliocene–Pleistocene boundary. Volcanic vent distribution and the shape of the volcanic edifices (subaerial and submarine lavas) seem to be controlled also by NE–SW- and E–W-trending lineaments, with a subordinate role of the NNE–SSW-trending lineaments (Fig. 10).

(ii) Between 1.1 and 0.38 Ma, and even more recently during the Holocene (19 ka; Principe, Arias & Zoppi, Reference Principe, Arias and Zoppi2003), rhyolitic extrusive activity was focused along the NNE–SSW-trending zones of central Milos or at their northern tips, as well as along the southern tip of the Filacopi Fault (FF) that bounds the active N–S-trending Zephyria graben (ZG, Fig. 10). The younger activity of the N–S-striking faults is confirmed also by offshore seismic reflection profiles, showing a transition from the E–W to the N–S fault strike at about 1.5–0.7 Ma (Piper & Perissoratis, Reference Piper and Perissoratis2003). Seismic profiles in the Milos and Folegandros offshore basins indicate that the N- to NNE-trending faults show displacements of more than 100 m and the progressive offset of the graben fill suggests that it is a long-lived structure (Piper & Perisoratis, Reference Piper and Perissoratis2003; Anastasakis & Piper, Reference Anastasakis and Piper2005).

(iii) Finally, NW-trending oblique-normal faults appear to be important for the recent shape of the island. This fault strike formed the Gulf of Milos (Fig. 10, i.e. ACHF), while many recent (0.38–0.1 Ma) subaerial volcanic centres (Fig. 10, i.e. Trachylas and Fyriplaka) are formed at the tip zone of this NW–SE-trending graben structure. This tectonic structure remains still active, as is indicated by the young phreatic centres, active fumaroles and solfataras and the alignment of seismic epicentres of weak earthquakes, with shallow focal depths in the order of 1–4 km (Fig. 10; Papanikolaou et al. Reference Papanikolaou, Lekkas, Syskakis and Adamopoulou1993). The focal mechanism of the moderate-sized (Ms = 5.3) 20 March 1992 earthquake, which also caused surface ruptures, indicates an E–W tensional direction along the NW-trending rupture zone with a slight right-lateral component of motion.

5.b. Santorini volcanic centre

On Santorini Island, early volcanic activity (~ 750–350 ka) was restricted to the Akrotiri Peninsula on the SW edge of the island and to the Peristera volcano in the NE part, comprising basaltic to andesitic lavas, tuffs and breccias (Fig. 11a). Above these units the products of two eruptive cycles were deposited, comprising andesitic to rhyodacitic pyroclastic deposits in the first cycle (360–180 ka) and basaltic to rhyodacitic lavas and minor pyroclastic deposits in the second cycle (180–3.6 ka). After the devastating caldera-forming Minoan eruption (~ 1600 bc), submarine volcanic eruptions recommenced at the centre of the caldera, consisting of dacitic flows lasting from 197 bc to 1950 and forming the Kameni Islands (Fig. 11 a, b; Heiken & McCoy, Reference Heiken and McCoy1984; Druitt et al. Reference Druitt, Edwards, Lanphere, Sparks, Davis, Casale, Fytikas, Sigvaldason and Vougioukalakis1998, Reference Druitt, Edwards, Mellors, Pyle, Sparks, Lanphere, Davis and Barriero1999).

Figure 11. (a) Simplified tectonic map of the Santorini volcanic centre showing the distribution of dacite domes, cinder cones and submarine volcanoes with regard to the volcano-tectonic lineaments. Stereonet 1 displays fault slip data in the northern part of the island close to the Columbo line. (b) Map of the Santorini–Amorgos Fault Zone. Areas of higher earthquake activity (microseismicity) are marked by the two light grey shaded areas with dashed outline. (c) Inset shows fault pattern in the Christianoi Islands (modified from Mountrakis et al. Reference Mountrakis, Pavlides, Chatzipetros, Meletlidis, Tranos, Vougioukalakis, Kilias, Casale, Fytikas, Sigvaldason and Vougioukalakis1998). (d1) Seismic reflection profile (profile 1 on map) across the Columbo submarine volcano (taken from Sakellariou et al. Reference Sakellariou, Sigurdsson, Alexandri, Carey, Rousakis, Nomikou, Georgiou and Ballas2010). (d2) Composite seismic profile (profile 2 on map) along the Columbo volcano and Anydros basin (from Sakellariou et al. Reference Sakellariou, Sigurdsson, Alexandri, Carey, Rousakis, Nomikou, Georgiou and Ballas2010). The bright spot area below volcano caldera suggests the presence of fluids or gas about 200 m beneath the caldera.

In the older volcanic units of SW Santorini near Akrotiri, cinder cones form resistant NE-trending spatter ridges (Fig. 11a). Dykes, exposed in the caldera walls and on lava shields, are almost vertical and have a mean orientation of N28°E (standard deviation ~ 14°; Heiken & McCoy, Reference Heiken and McCoy1984). Most of the dykes are feeders for the small volcanoes. Submarine volcanoes and dacitic domes are aligned along this volcano-tectonic lineament as far southwest as the Christianoi Islands, 25 km SW of Santorini (Fig. 11 a–c). These islands consist mainly of lavas and pumice flows and display WNW- and NNE-trending normal fault sets.

The Anafi Fault Zone, extending from Santorini to Amorgos, marks a major transtensional structural boundary with a right-lateral component of sense of shear, as is indicated by recent GPS data (on a local reference frame, see also Fig. 13), focal mechanism solutions of earthquakes and seismic profiling data (McClusky et al. Reference McClusky, Balassanian, Barka, Demir, Ergintav, Georgiev, Gurkan, Hamburger, Hurst, Kahle, Kastens, Kekelidze, King, Kotzev, Lenk, Mahmoud, Mishin, Nadariya, Ouzounis, Paradissis, Peter, Prilepin, Reilinger, Sanli, Seeger, Tealeb, Toksöz and Veis2000; Bohnhoff et al. Reference Bohnhoff, Rische, Meier, Becker, Stavrakakis and Harjes2006). The fault zone has a NE-orientation and bounds the Anafi basin from the Santorini–Amorgos ridge (Fig. 11b). To the northwest, another significant fault zone, the Ios Fault Zone is a segmented NE-trending linear zone that bounds the Anydros basin and also displays right-lateral shearing, as is indicated by seismic profiles (Sakellariou et al. Reference Sakellariou, Sigurdsson, Alexandri, Carey, Rousakis, Nomikou, Georgiou and Ballas2010). Inside the basin and towards its southeastern margin, en échelon arranged fault segments also strike NE to ENE. Along the seafloor of the basin, more than 20 volcanic cones are aligned on this orientation (Sakellariou et al. Reference Sakellariou, Sigurdsson, Alexandri, Carey, Rousakis, Nomikou, Georgiou and Ballas2010). The fault defining the southeastern margin of the Anydros basin (Fig. 11d2) passes from the Columbo submarine volcanic seamount (which erupted in ad 1650) to the Peristera submarine summit, and crosses through the northern part of the Santorini caldera, where a small graben is formed (Fig. 11a, b). The presence of this fault system, in a N45–50°E orientation, seems to facilitate the ascent and concentration of the magma along this trend (Fig. 11d1). The highest earthquake activity has been recently observed beneath the submarine Columbo volcano and northeast of the volcano along the Santorini–Amorgos ridge (Fig. 11b, grey-shaded area on map), while it terminates south of the island of Amorgos (Hubscher et al. Reference Hubscher, Hensch, Dahm, Dehghani, Dimitriadis, Hort and Taymaz2006; Bohnhoff et al. Reference Bohnhoff, Rische, Meier, Becker, Stavrakakis and Harjes2006; Dimitriadis et al. Reference Dimitriadis, Karagianni, Panagiotopoulos, Papazachos, Hatzidimitriou, Bohnhoff, Rische and Meier2009).

On a smaller scale, the dykes observed in Skaros and Oia villages show a main NE–SW orientation, while a subordinate population concentrates in a NNE and NW orientation (Mountrakis et al. Reference Mountrakis, Pavlides, Chatzipetros, Meletlidis, Tranos, Vougioukalakis, Kilias, Casale, Fytikas, Sigvaldason and Vougioukalakis1998). The cinder cones on the Kameni Islands are also aligned along this N45°E fault strike. Fault-slip data from NE-trending mesoscale faults cutting the pre-Minoan volcanic deposits show a right-lateral oblique-normal motion along them (see also Mountrakis et al. Reference Mountrakis, Pavlides, Chatzipetros, Meletlidis, Tranos, Vougioukalakis, Kilias, Casale, Fytikas, Sigvaldason and Vougioukalakis1998), suggesting a NNW-trending σ₃ stress axis (Fig. 11a, stereonet 1). This stress direction seems to have been active up to the present, as is confirmed by the mixed normal and dextral-normal focal mechanisms of recent earthquakes along the Santorini–Amorgos ridge (Dimitriadis et al. Reference Dimitriadis, Karagianni, Panagiotopoulos, Papazachos, Hatzidimitriou, Bohnhoff, Rische and Meier2009).

Towards the south, with a separation of 3 km, another major tectonic lineament lies approximately parallel to the regional N50°E fault system (Heiken & McCoy, Reference Heiken and McCoy1984), concentrating a NE–SW zone of surface faulting, vent alignment and gas emissions (Druitt et al. Reference Druitt, Edwards, Lanphere, Sparks, Davis, Casale, Fytikas, Sigvaldason and Vougioukalakis1998). This lineament (Kameni line) intersects the caldera wall, where oblique-normal faulting is evident, passes through the Kameni Islands and extends offshore southeast of the island (Fig. 11a). The Kameni lavas have been erupted from NE–SW-trending fissures during 12 observed periods of activity over the last 2ka, with the latest eruption occurring in 1950 (Georgalas, Reference Georgalas1962). Additionally, the Kameni line, relative to the local bathymetry, separates the caldera floor into a deeper northern part (max. ~ 380 m below sea level (bsl)) and a southern shallower part (max. ~ 280 m bsl; Fig. 11a).

Inference of recent strike-slip deformation in the northeast of Amorgos Island is based on morphotectonic criteria and recent fault surfaces in a few outcrops of Pleistocene and recent alluvial deposits (Papadopoulos & Pavlides, Reference Papadopoulos and Pavlides1992). Focal mechanisms are also consistent with a dextral component of motion along the Amorgos–Santorini ridge border faults (Fig. 13). The rupture zone between the islands of Santorini and Amorgos has a composite length of about 50–60 km, as is determined from the distribution of earthquake epicentres, which corresponds well to the length of a fault responsible for the 9 July 1956 earthquake of Ms = 7.5 (Pavlides & Caputo, Reference Pavlides and Caputo2004). This event and its strongest aftershock (Ms = 7.2) close to Santorini Island (Fig. 11b) caused severe damage to Santorini, as well as tsunami inundation.

Furthermore, interpreted offshore seismic profiles suggest that in the last 0.2 Ma, an abrupt basin inversion occurred in the Santorini–Anafi area, where basinal-facies units are now part of the prominent narrow ridges, and self-progradation units are now at a depth of 600 m (Piper & Perissoratis, Reference Piper and Perissoratis2003). These abrupt morphological changes can be attributed to the strike-slip component of motion along the oblique-normal NE–SW-trending fault segments of the Santorini–Amorgos ridge. The ENE-trending normal faults and lineaments arranged oblique to the main NE-trending fault, along the Santorini–Amorgos ridge (Fig. 11b), provide clues about extensional deformation due to the right-lateral oblique-normal motion between the paired right-stepping Ios and Anafi fault zones, which form an extensional strike-slip duplex structure (Fig. 11b).

In summary, we can conclude that the Santorini volcanic field has been developed in an area that has undergone extensional deformation accommodated by en échelon arranged, NE–SW-trending oblique-normal fault segments. The volcanic activity occurred along the inferred lineaments and at the overlap zone between these fault segments, whereas this deformation appears to be associated with the major extensional strike-slip duplex structure formed between the Ios and Anafi fault zones.

5.c. Nisyros volcanic centre

Nisyros Island is a Quaternary stratovolcano with a well-developed caldera, 4 km wide and 300 m deep; a second caldera, the one of Yali, lies partly submerged ~ 5 km off Nisyros's northern coast (Fig. 11b). The exposed eruptive products on Nisyros range from basaltic andesites to rhyodacites with ages of 100–150 ka and < 20 ka, respectively (Seymour & Vlassopoulos, Reference Seymour and Vlassopoulos1992; Vougioukalakis, Reference Vougioukalakis1993; Francalanci et al. Reference Francalanci, Vougioukalakis, Perini, Manetti, Fytikas and Vougioukalakis2005). The main volcanic infrastructure of Nisyros consists of lava flows, pyroclastic rocks and feeder dykes of andesitic composition (Seymour & Vlassopoulos, Reference Seymour and Vlassopoulos1992). The formation of the present caldera was related to Plinian eruptions of < 20 ka (Limburg & Varekamp, Reference Limburg and Varekamp1991).

The major fault strikes identified on Nisyros are the NE-, NW- and almost E–W-trending faults. The NE strike is part of the major NE- to ENE-trending structure that shaped the south coast of Kos Island and Datca peninsula in western Turkey. The post-caldera activity was characterized by the emplacement of dacitic lava domes aligned in a NE–SW orientation and located inside and outside the Lakki Plain (LP, Fig. 12a), which represents the floor of the Nisyros caldera. Post-caldera domes and craters of hydrothermal eruptions are also aligned in narrow zones of NE–SW orientation (Fig. 12a). These faults, although segmented, run across the whole island, cut the southern rim of the caldera and reach the southern coast of the island. This fault system delimits a NE–SW elongated graben-like structure in the caldera, which is the most prominent feature on the island. Another striking characteristic of the island is that the hydrothermal craters have an elliptical shape and are elongated in a NE orientation (Fig. 12a, Polyvotis (PV) and Stefanos (ST)). They are located at the intersection between the NE- and NW-trending faults (Fig. 12a; Lagios et al. Reference Lagios, Sakkas, Parcharidis and Dietrich2005).

Figure 12. (a) Fault–lineament map of Nisyros Island with the three main fault orientations shown with different shades of grey. The hydrothermal explosion craters of Polyvotis (PV) and Stefanos (ST) are shown inside the caldera rim. LP – Lakki plain; LR – Lakki rupture, MF – Mandraki Fault; PE – Profitis Elias. (b) Map of the broader area of the Nisyros–Kos volcanic centre. Stereonet plots provide fault slip data from Kos (stereonets 1, 2) and Nisyros (stereonets 3, 4) islands. (Data from Limburg & Varekamp, Reference Limburg and Varekamp1991 and Kokkalas & Doutsos, Reference Kokkalas and Doutsos2001.)

The NW-trending faults seem to have higher dip angles, offset the southern caldera rim and run through the hydrothermal eruption craters (Brombach et al. Reference Brombach, Hunziker, Chiodini, Cardellini and Marini2001). The NW- to NNW-trending Mandraki Fault (Fig. 12a, MF; P. Nomikou, unpub. Ph.D. thesis, Univ. Athens, 2004; Volentik et al. Reference Volentik, Vanderkluysen, Principe, Hunziker, Hunziker and Marini2005) is an active feature, which possibly extends from the northwestern end of the caldera to the offshore area south of Kos Island, which is probably responsible for the two moderate-sized earthquake events of M = 5.3 and M = 5.2 (27 August 1997), causing damage in the village of Mandraki (Papadopoulos et al. Reference Papadopoulos, Sachpazi, Panopoulou and Stavrakakis1998). Detailed studies showed that maximum uplift deformation and seismic activity was concentrated on the NW flank of the volcanic structure, based on seismicity and InSAR data (Sachpazi et al. Reference Sachpazi, Kontoes, Voulgaris, Laigle, Vougioukalakis, Sikioti, Stavrakakis, Baskoutas, Kalogeras and Lepine2002; Lagios et al. Reference Lagios, Sakkas, Parcharidis and Dietrich2005).

Mesoscale structural data are in agreement with the far field picture, similar to that from the neighbouring Kos Island (Kokkalas & Doutsos, Reference Kokkalas and Doutsos2001; Fig. 12, stereonets 1, 2) indicating that faults with NE–SW strike show oblique-normal slip with a left-lateral component of motion (Fig. 12b, stereonet 3; see also Caliro et al. Reference Caliro, Chiodini, Galluzzo, Granieri, La Rocca, Saccorotti and Ventura2005). Secondary N–S-trending faults with pure normal slip represent tension fractures that occur often in the tip of the main faults (Limburg & Varekamp, Reference Limburg and Varekamp1991), accommodating extension associated with the left-lateral component of motion on the major NE-trending faults. A regional reactivation of the ENE-trending faults as sinistral strike-slip faults led to extension on the older NW-trending lineaments, which are located on the relay zone. At the tips of the NE-striking faults onshore, hot springs occur at sea level forming a spatial cluster at or near the faults (Fig. 12a). Veins and hydrothermally altered zones also concentrate along this fault orientation. Recently, Tibaldi et al. (Reference Tibaldi, Pasquare, Papanikolaou and Nomikou2008) reported that the NE- to ENE-trending lineaments all over the caldera floor are faults and fractures. This fault strike controls offshore tectonic grabens with a left-lateral component of slip (Nomikou & Papanikolaou, Reference Nomikou and Papanikolaou2000; Pe-Piper & Piper, Reference Pe-Piper, Piper, Fytikas and Vougioulakakis2005). The presence of NE-striking fault scarps offshore, affecting the deposits below the Holocene lava domes, indicates that significant motion along the NE–SW faults occurred before dome emplacement and that they possibly acted as feeders for magma rise (Tibaldi et al. Reference Tibaldi, Pasquare, Papanikolaou and Nomikou2008).