The Montagna dei Fiori Anticline is intersected by two major fault categories (Storti et al., 2018), which based on the chronological order include the following. Firstly, ∼ E–W- and ∼ N–S-striking fault zones showing extensional kinematics bounding the major outcrops of Calcare Massiccio and dominantly affecting the Jurassic rocks older than the Maiolica Formation (Fig. 2a; e.g., sites 1 to 4). Overprinting relations indicate that ∼ E–W deformation structures are systematically younger than the ∼ N–S ones. Similar trends were observed in syn-rift fault zones in other anticlines of the Central Apennines (e.g., Cooper and Burbi, 1986; Alvarez, 1989; Chilovi et al., 2002). Such a tectonosedimentary inheritance was involved in the growth of the Montagna dei Fiori Anticline, which initiated during the Late Miocene (Mazzoli et al., 2002; Artoni, 2003) and progressively evolved into the upper thrust sheet of a well-developed antiformal stack until Plio-Pleistocene times (e.g., Mazzoli et al., 2002; Calamita et al., 1994; Artoni, 2013). The second set of faults is a major structural feature trending parallel to the Montagna dei Fiori Anticline and dissecting it is the Montagna dei Fiori Fault, a NNW–SSE-striking extensional fault system cutting at a high angle through the folded footwall rocks, typically at the forelimb–crest transition (Figs. 1, 2). This fault consists of two partially overlapping main fault zones with an extensional stratigraphic separation exceeding 900 m, and juxtaposes intensely deformed Late Miocene sediments in the hanging wall against dolomitized and undolomitized Lower Jurassic and Cretaceous limestones in the footwall (Figs. 1 and 2). The development of the Montagna dei Fiori Fault has been alternatively interpreted as either a pre-folding (e.g., Calamita et al., 1994; Mazzoli et al., 2002; Scisciani et al., 2002) or late-folding (Ghisetti and Vezzani, 2000) feature. More recently, the origin of the Montagna dei Fiori Fault has been ascribed to the mutual interaction between horizontal shortening and uplift and episodic gravitational re-equilibration during antiformal stacking underneath the anticline during Plio-Pleistocene times (Storti et al., 2018). The dolomitized intervals are exposed in the damage zones of both aforementioned fault categories.

There is no evidence of dolomitization in the overlying and immediate surrounding successions of the Calcare Massiccio, Bugarone, and Corniola formations (e.g., Maiolica and Scaglia formations), though the base of Maiolica Formation is reported as dolomitized in the Central Apennines onshore (e.g., Pierantoni et al., 2013) and offshore areas (Murgia et al., 2004).

Dolomitized intervals are folded in the forelimb of the Montagna dei Fiori Anticline and are abruptly truncated by the Montagna dei Fiori Fault, which juxtaposes them against intensely foliated Scaglia, Bisciaro, and Marne con Cerrogna formations (Figs. 2 and 3). The distribution of dolomitized intervals is wider in the Salinello creek (Figs. 1b, 2a), perhaps due to a better exposure. In the Corano Quarry location, dolomitization occur in the Calcare Massiccio and Bugarone formations only as meter-sized dolostone geobodies in the footwall of the Montagna dei Fiori Fault (Fig. 4). The map pattern (Fig. 2) of dolostones indicates that their distribution is maximized in the Castel Manfrino to Osso Caprino hill area and fades out both southward and eastward.

The lateral extent of dolomitization is gradual. In some outcrops, dolomitization fronts show irregular outlines following, but also cross-cutting, bedding surfaces (Fig. 5). Dolomitized intervals vary in thickness from a few meters to hundreds of meters affecting the totality of the exposed Calcare Massiccio and the lower part of Corniola Formation, where no clay interlayers are present. In the Calcare Massiccio Formation, dolomitization does not follow a systematic pattern. In the northern side of the Osso Caprino hill (Fig. 2), the top of formation is dolomitized but moving toward the Salinello creek and a thick non-dolomitized limestone is exposed. The same situation occurs on the opposite side of the creek and to the east of Castel Manfrino.

Dolomitized intervals in the Corniola Formation have a darker color relative to the host rock and are systematically more fractured than the hosting limestone. High-amplitude (> 1 mm) bed-parallel stylolites are clearly visible in both limestones and dolostones (Fig. 5). However, in some dolostones only ghosts of stylolite traces can be seen. No apparent porosity could be observed in host rock limestones but the dolostones locally contain porosity, appearing as millimeter- to centimeter-sized pores. Dolostone breccias in fault cores are typically clast supported, with angular and millimeter- to centimeter-sized fragments (Fig. 3c), changing to crackle breccia (Woodcock and Mort, 2008) away from the main slip surface. In the proximity of the main slip surface, dolostone fragments are sporadically cross-cut by millimeter-sized dolomite veins. The breccia fragments, where cemented, are commonly surrounded by calcite cement.

The overview of microthermometry measurements is given in Table 1 and Fig. 13a to c. All the measured fluid inclusions are primary and occur in growth zones. Based on optical and fluorescence microscopy analysis of wafers, all the inclusions are aqueous monophase (liquid) or have two phases (liquid and vapor) with relatively consistent L:V ratios of 10 %–15 % within a single FIA (fluid inclusion assemblage). Special care was taken to avoid the samples that occasionally displayed scattered mottled luminescence that may indicate recrystallization.

D1 contains dominantly monophase aqueous inclusions with sizes greater than 5 µm. All the inclusions froze at -65 to -49 ∘C. The first melting (Te) was detected between -22 and -19.3 ∘C. The final ice melting (Tm) appeared at temperatures between -7.7 and -2 ∘C. Applying the equation of Bodnar (1993), the obtained final melting temperatures correspond to salinity ranges of 3.5 to 11.3 eq. wt % NaCl.

D2 is characterized by the presence of monophase and infrequent two-phase inclusions generally within their growth zones. The homogenization temperature of two-phase inclusions varies between 58 and 71 ∘C. Upon cooling, a complete freezing of the fluid phase is reached at -56 to -40 ∘C. The first ice melting temperature was distinguished at -22 ∘C. The final ice melting temperatures fall within -17.5 and -5 ∘C, corresponding to salinities between 7.9 and 20.5 eq. wt % NaCl.

D3 is commonly inclusion poor. The measurable inclusions were detected and examined only in saddle dolomite crystals. These crystals contain only two-phase aqueous inclusions. Their homogenization temperatures are within the narrow range of 70 to 73 ∘C. The complete freezing and first ice melting temperatures could not be distinguished but the final ice melting temperature occurred at temperatures between -13 and -6 ∘C equal to salinity ranges of 9.2 to 16.9 eq. wt % NaCl. The first melting temperatures of fluid inclusions in D1, D2, and D3 were about -21 ∘C, suggesting a H2O-NaCl fluid system.

D4 contains only two-phase aqueous inclusions. The homogenization temperatures in D4 vary between 79 and 105 ∘C. Complete freezing of inclusions occurred at temperatures between -86 and -54 ∘C. The first ice melting was detected from -35 to -40 ∘C indicating the possible presence of divalent cations such as Ca2+ and/or Mg2+ in the fluids (Shepherd et al., 1985; Goldstein and Reynolds, 1994). The final ice melting temperatures fall within a range of -15 and -9 ∘C corresponding to salinities of 12.8 to 18.6 eq. wt % NaCl. A couple of inclusions show homogenization temperatures exceeding 120 ∘C with salinities higher than 20 eq. wt % NaCl.

No measurable fluid inclusion could be identified in C1 and C2 due to intense deformation twinning. C3 and C4 contain only primary monophase aqueous inclusions, indicating an entrapment temperature of below about 40–50 ∘C (Goldstein and Reynolds, 1994). A complete freezing of the inclusions in C3 occurred at temperatures between -40 and -52.5 ∘C. The first melting temperature was detected at about -21 to -22 ∘C, suggesting a H2O-NaCl composition. The final melting temperatures range between -6.4 and -2.7 ∘C, corresponding to salinities between 9.7 and 4.5 eq. wt % NaCl. The majority of the values cluster between 7.8 and 5 eq. wt % NaCl.

The complete freezing temperatures of the inclusions in C4 fall within -46 and -35.5 ∘C. The first melting temperature could not be determined with confidence but the final melting temperatures were reached at about -0.1 to -1.8 ∘C, corresponding to salinities of 0.17 to 3.0 eq. wt % NaCl.

The δ13C values of all dolomite types mimic the range of host rock and Jurassic marine limestones and, consequently, they can be interpreted as largely rock buffered. Their δ18O values are partly similar to those of their respective host rocks as well as Jurassic marine reference values but more depleted when compared to the presumable Jurassic marine dolomites. The relatively depleted δ18O dolomite values could indicate the contribution of heated fluids in dolomitization process, although they could also relate to recrystallization of a precursor dolomite by fluids at higher temperature or 18O-depleted dolomite (Land, 1980, 1985). The absence of distinctive textural evidence in the analyzed samples such as enlarged crystal size and/or systematic mottled cathodoluminescence pattern, and their covariation with δ18O values, do not confirm recrystallization (Mazzullo, 1992, and references therein).

The oxygen isotope fractionation relation between water and dolomite (Land, 1983) was used to determine the most plausible parental fluids. In order to avoid erroneous results due to rock-buffered δ18O values, only the δ18O values of dolomite cements, especially from the bed-parallel veins containing D4 were used. These values may provide the closest approximation to the δ18O signature of the parental fluids (Barker and Cox, 2011). Accordingly, a δ18O value of ≈+2.5 ‰ to +4 ‰ V-SMOW (Vienna Standard Mean Ocean Water) was calculated for D3, while this values increase to ≈+5 ‰ to +7.5 ‰ V-SMOW for D4 (Fig. 13d). The higher δ18O composition of the dolomitizing fluids relative to the Mesozoic seawater, which is estimated at -1.2 ‰ to -1 ‰ V-SMOW (Shackleton and Kennett, 1975; Marshall, 1992; Saelen et al., 1996), is compatible with fluids derived from or that had interacted with siliciclastic rocks, crystalline basement (Taylor, 1997) and/or evaporite-derived brines.

The 87Sr/86Sr ratios obtained for all dolomite types are higher than the Lower Jurassic marine carbonate values (0.70704–0.70768; McArthur et al., 2012). Since marine carbonates have very low rubidium (Rb) concentrations they produce negligible in situ radiogenic 87Sr after their deposition (Stueber et al., 1972; Burke et al., 1982). Therefore, the higher 87Sr/86Sr ratios can be explained by the contribution of fluids originated or interacted with potassium rich siliciclastic rocks (K feldspars), crystalline basement, and/or stratigraphic levels with higher 87Sr/86Sr ratios (Emery and Robinson, 1993; Banner, 2004). Taking into account that the Upper Triassic Burano Formation underlying the studied intervals as the basal detachment has 87Sr/86Sr ratios between 0.70774 and 0.70794 (Boschetti et al., 2005), the 87Sr/86Sr ratios (D1 and D4) can partially be explained by their contribution. However, this contribution cannot justify much higher 87Sr/86Sr ratios recorded in D3, being higher than values reported for Phanerozoic seawater (McArthur et al., 2012), and the values obtained for the adjacent basinal deposits (i.e., Corniola and Scaglia formations). Therefore, parental fluids most likely originated from or had interacted with the siliciclastic rocks underlying the Burano Formation (Verrucano Formation), if present, and/or with the crystalline basement with common elevated 87Sr/86Sr ratios (0.71500–0.72650; Del Moro et al., 1982). The significantly higher 87Sr/86Sr ratios in D3 in comparison with other studied dolomites indicates a higher influence of 87Sr/86Sr-rich fluids either due to major changes in the permeability architecture of faults or availability of such fluids. The lack of any ferroan diagenetic phase minimizes the interaction of fluids produced by clay transformation or dewatering (i.e., smectite to illite transformation; Boles and Franks, 1979).

C1 is characterized by δ13C and δ18O values lower than the host limestones (i.e., Calcare Massiccio), while its 87Sr/86Sr ratio is similar to them. The salinity and composition of the parental fluids cannot be inferred here since no measurable fluid inclusions were found within this cement. The 87Sr/86Sr ratio being within the range of the corresponding host rocks and the reference values, points to a rock-buffered system for 87Sr/86Sr.

The δ13C and δ18O values obtained for C2, as well as 87Sr/86Sr ratios, fall within the range of the Scaglia host rocks, thus reflecting their rock-buffered nature. This interpretation is further supported by the similar luminescence characteristics of C2 with that of encasing Scaglia host rocks. The fluids from which C2 calcite precipitated, as expected for tension gashes, were most likely derived from carbonate dissolution during pressure solution and stylolitization of host rock, pointing to a closed fluid system in contrast with the subsequent vein generations.

C3 is characterized by δ13C values within the Jurassic marine values but are generally lower than the host rocks, while their δ18O values partially overlap both the hosting limestones and dolostones. Microthermometry of fluid inclusions revealed only monophase aqueous inclusions and thus precipitation at relatively low temperature (≤40–50 ∘C) with moderate salinity (4.5–9.7 eq. wt % NaCl). Such levels of salinity can be assigned to evaporated seawater, residual brines or fluids derived from evaporite dissolution, and thus makes it difficult here to interpret their exact origin with the available data.

C4 is the latest calcite phase, and records the δ13C and δ18O values, respectively enriched and significantly depleted when compared to their hosting rocks and preceding diagenetic products. Generally, the enrichment of 13C could suggest CO2 outgassing due to evaporation or pressure changes (Friedman, 1970; Hendry et al., 2015) or bacterial fermentation (methanogenesis) of organic matter (Hudson, 1977) in low-temperature diagenetic environments. The homogenization temperature of C4, being below about 40–50 ∘C, could support any of these processes. Their low δ18O values and fluid inclusions with salinities similar to, but also significantly lower than, seawater reflect the contribution of meteoric fluids during precipitation of this calcite.

The contribution of brines that derived from highly evaporated seawater or evaporites is suggested by the elevated salinity values obtained from microthermometry of the fluid inclusions (3.5 to 20.5 eq. wt % NaCl). Accordingly, two sources that could potentially provide such fluids can be proposed: (1) fluids related to the Late Messinian evaporites, associated with the overlying Upper Miocene Laga Formation, and their possible downward percolation through fault zones by density driven flow and/or seismic pumping mechanisms (Sibson, 1981; McCaig, 1988, 1990); or their tectonic involvement into the Apenninic thrust wedge during its propagation (underthrusting; Lobato et al., 1983); and (2) fluids related to the underlying detachment horizon of the Burano evaporites (Upper Triassic) and their upward flow through fault zones during development of the Montagna dei Fiori Anticline. The first scenario is valid if the dolomitization would have occurred only from the Upper Miocene time onwards. Moreover, several researchers (e.g., Vai and Ricci Lucchi, 1977; Bassetti et al., 1998; Roveri et al., 2001) have shown that the occurrence of primary shallow-water evaporites, which were dominantly gypsum, was limited only to the western and central parts of the northern Apennines consisting of thrust-top marginal basins (Roveri et al., 2001). Hence, the evaporitic horizons existing within the Laga Formation corresponds to resedimentation (gypsum debris) of those previously precipitated in the marginal basins. This interpretation makes the Messinian evaporites an unlikely source of Mg-rich brines. Taking into account that the maximum burial-related temperature of the Calcare Massiccio Formation did not exceed 80 ∘C in the Montagna dei Fiori region (Ronchi et al., 2003), it is unlikely that the downward percolation of relatively low-temperature brines derived from the Messinian evaporites, located at higher stratigraphic levels, could reach or exceed the high temperatures recorded in fluid inclusions of the studied dolomites (D4; up to 105 ∘C), given that the homogenization temperatures reflect the minimum entrapment temperatures (Goldstein and Reynolds, 1994). Deep circulation of these brines, if they existed, can also be excluded by the fact that their limited involvement in the thrust wedge was confined merely to the offshore side of the Montagna dei Fiori region (Artoni, 2013).

Accordingly, the Upper Triassic Burano Formation – the basal detachment – appears as the most plausible source for the high-salinity brines recorded in fluid inclusions, and likewise the Mg-rich fluids could have originated from postevaporite brines associated with them (Carpenter, 1978; McCaffrey et al., 1987). The fluctuations in salinity may argue for a diverse range of fault connectivity, different degrees of rock–water interaction, and contribution of pore waters of lower salinity (e.g., marine or meteoric).

A generalized paragenesis and the relative chronology of dolomitization in relation to the structural evolution of the Montagna dei Fiori Anticline are illustrated in Figs. 14 and 15. The structural episodes are based on the evolutionary stages of the Montagna dei Fiori Anticline suggested by Storti et al. (2018). The paragenesis is constructed on the basis of direct evidence recorded during observations at outcrop scale and microscopic observations (e.g., cross-cutting relationships between diagenetic phases, stylolites, fractures, and other structural kinematics), and indirect evidence (e.g., regional geodynamics and burial history).

The occurrence of micritic envelopes and fibrous calcite cements (FC), in grain-supported stratigraphic levels of the Calcare Massiccio Formation, is interpreted to be of eogenetic origin (i.e., marine phreatic diagenesis; Moore, 1989), reflecting an early diagenesis shortly after deposition. The well-developed brown-and-orange concentric cathodoluminescence pattern of the succeeding mosaic calcite cement (MC) suggests a progressive shift to more reducing conditions during precipitation in a phreatic diagenetic environment (as shown in Li et al., 2017). High-amplitude bed-parallel stylolites postdate both cements, which confirm their precipitation before significant burial. The observations made here are in agreement with earlier work by Giacometti and Ronchi (2000), interpreting that the Calcare Massiccio Formation was cemented during the early diagenetic stages.

D1, C1, and D2 are cut by well-developed high-amplitude bed-parallel stylolites. The presence of D1 and C1 in bed-perpendicular veins typically cut by these stylolites (see Fig. 6e to h) support the interpretation that the first dolomitization event (D1 and D2) took place before significant burial and stylolite development. The latter and bed-perpendicular veins are dynamically compatible within the same stress field, which is characterized by a vertical load-related maximum principal axis of the stress ellipsoid (Fig. 15a). The dominantly monophase fluid inclusions within D1 and D2 are in agreement with precipitation temperatures below about 40–50 ∘C, suggesting a relatively shallow to intermediate burial environment and hence supporting a pre-Apenninic orogeny age of precipitation from a mix of formational and extra-formational fluids with elevated 87Sr/86Sr ratios. The distribution of D1 and D2 localized near the rifting-related ∼ N–S- and E–W-striking extensional faults and even their displacement along them (Fig. 2a; e.g., site 1) point to the possible contribution of these faults in occurrence of D1 and D2. These faults dominantly affect the Jurassic rocks older than the Maiolica Formation which is attributed to post-rift deposits, therefore suggesting a pre-Maiolica age for these dolomite types. Although an absolute age cannot be provided, based on the evidence discussed above, the circulation of Mg-rich fluids during this dolomitization event was most likely controlled by rifting-related Jurassic extensional fault zones cutting through the crystalline basement. Precipitation of D1 and D2 at the lower part of Corniola Formation, which is known as the syn-rift deposit, discards a pre-rift origin for these dolomites. The displacement of dolomites along the aforementioned faults is possibly related to their prolonged activation during the Early Jurassic to Late Jurassic. In addition to the role of these faults in channelizing the fluids, their mobilization must have been intensified by some deriving mechanisms. A thermal convection system derived from high hit flux during rifting was interpreted by Hollis et al. (2017) to be responsible for circulation of seawater in a syn-rift dolomitization case in the Hammam Faraun fault block (Suez Rift, Egypt). In such a scenario, the salinity of the fluids and their 87Sr/86Sr ratios are expected to be more or less within the range of seawater. Furthermore, this scenario seems unlikely in the studied area given the lack of a deep aquifer to accommodate the fault tips and promote the lateral fluid flux from basin to the rift shoulders and vice versa. Taking into account that D1 and D2 are the volumetrically more relevant dolomites within the studied intervals, and assuming the likely role of syn-rift extensional faults (Early Jurassic to Late Jurassic) in their precipitation, a dominantly syn-rift dolomitization process is proposed for the dolostones in the Montagna dei Fiori Anticline. Although the CL zonation pattern observed in D2 may indicate changes in flow condition or fluid composition, the lack of physical disruptions such as multiple fracturing suggests external regional controls rather than slip along the same faults (Eichhubl and Boles, 2000). The absence of pervasive syn-dolomitization fracturing and brecciation as well as zebra fabrics in these dolomites perhaps indicate a relatively calm tectonic period during dolomite development (e.g., Hollis et al., 2017).

D3 and D4 both record elevated 87Sr/86Sr ratios which account for their fault-controlled origin. However, their occurrence at the top of the Calcare Massiccio and overlaying Bugarone Formation (Corano Quarry site), which is < 1 m thick in the Montagna dei Fiori region and is marked as the final rift deposit (Cardello and Doglioni, 2015), discards a syn-rift origin for these dolomites. Moreover, D3 and D4 postdate the development of high-amplitude bed-parallel stylolites. The formation of stylolites requires an approximate overburden of 600 to 1500 m (Lind et al., 1993; Machel, 1999; Mountjoy et al., 1999; Schulz et al., 2016), corresponding to a late to post-Maiolica deposition time (Early Cretaceous time onwards). The presence of D3 and D4 dolomites in bed-parallel fractures and as shear veins (D4) (Fig. 9a and b) suggests their association with contractional deformations, i.e., the most likely tectonic regime for explaining bed-perpendicular dilation. Therefore, the volumetrically minor second stage of dolomite precipitation may possibly be related to the Late Miocene to post-Miocene compressional tectonics recorded in this region (e.g., Mazzoli et al., 2002; Artoni, 2013; Storti et al., 2018).

Dolostones containing D3 and D4 appear commonly as clast-supported breccias along fault zones pertaining to the Montagna dei Fiori Fault, then overprinted by fault-parallel stylolites (Figs. 3 and 7). Accordingly, the occurrence of these dolomites was probably synchronous with the incipient stages of fault development, predating fault buttressing (Storti et al., 2018). Homogenization temperatures recorded in D4 (up to 105 ∘C), much higher than the maximum temperatures recorded in the host rocks (below about 80 ∘C; Ronchi et al., 2003), suggest hydrothermal fluid circulation. The development of the Montagna dei Fiori Anticline at the toe of the Late Miocene Central Apennines thrust wedge could have favored the forelandward migration of hydrothermal fluids expelled from the more internal regions of the belt, similarly to what has been proposed for the Rocky Mountains foreland (i.e., squeegee flow model; Machel and Cavell, 1999). Such a migration may have possibly favored the precipitation of D4 in bed-parallel veins, generally considered evidence for syn-compressional fluid overpressure (Sibson, 2001; Hiemstra and Goldstein, 2015). At this stage, in addition to dilation of the pre-existing ∼ N–S- and E–W-striking rift-related extensional faults and their possible role in fluid migration, the excess of pore pressure at the base of the thrust ramp in the fold hinge and during fold tightening could promote the localization of the fractures (Smith and Wiltschko, 1996; Ghisetti and Vezzani, 2000), with fluid migration within this zone and eventually dolomitization. These fractures could have been corridors that later on formed the incipient NW–SE Montagna dei Fiori Fault. Their localization at the back limb, cross-cutting the core, best explains the distribution of D3 and D4 at this locality. The presence of only D5 within the damage zone of the Montagna dei Fiori Fault, postdating dolostone brecciation and, in places, cementing breccia fragments, suggests that D5 dolomite precipitation was associated with the late-stage evolution of the Montagna dei Fiori Fault, predating late-stage calcite precipitation. The shift from dolomite to calcite precipitation can be ascribed to attenuation of Mg-rich fluids and/or calcite saturation. This condition was perhaps initiated during the late stages of anticline evolution due to changes in fault conductivity sealing the upward migration of Mg-rich fluids.

The presence of several generations of bed-perpendicular stylolites bounding and intersecting C2 veins (Fig. 10) supports the postulation that late-stage calcite cements precipitated in close association with the deformation history of the Scaglia Formation in the hanging wall of the Montagna dei Fiori Fault (Fig. 3). This deformation occurred during buttressing against Calcare Massiccio and Corniola formations in the footwall and related to the positive inversion event induced by thrust-sheet stacking at depth (Storti et al., 2018). Precipitation of C3 and C4 is interpreted to have occurred during uplift and cooling as revealed by their relatively low homogenization temperatures (≤40–50 ∘C) of fluid inclusions trapped within these cements. Deformation twinning is either absent or weakly developed, reflecting the lack of significant tectonic deformation after calcite precipitation. These cements postdate the dolomitization events, high-amplitude bed-perpendicular and parallel stylolites, and are precipitated as cements bounding the breccia fragments within the damage zone of the Montagna dei Fiori Fault. Salinities calculated from their fluid inclusions, particularly in C4, suggests precipitation from meteoric waters, which should have been favored during the late evolutionary stages of antiformal stacking beneath the Montagna dei Fiori Anticline and eventual late extensional slip along the Montagna dei Fiori Fault (Storti et al., 2018). The results obtained in this study are in relative agreement with the earlier work by Ronchi et al. (2003) and Murgia et al. (2004) in the Central Apennines, assigning dolomitization phases to the pre- and syn-orogenic deformations, although they did not specify the direct relation between the local structures and the different types of dolomite.

The textures of the studied dolomites vary from planar-e to nonplanar; the preponderance of planar dolomite, as in D4, creates a rock with interesting poroperm characteristics (e.g., Woody et al., 1996; Wilson and Ruppel, 2007; Wenzhi et al., 2012). This case-study is certainly relevant for many potential reservoirs elsewhere in the world. Similar multistage burial dolomitization events enhancing the reservoir quality have been reported from the carbonate successions of the Jurassic in the Kopet Dagh Basin, north eastern Iran (Adabi, 2009), and Devonian of the Rainbow subbasin, western Canada (Qing and Mountjoy, 1989; Lonnee, 1999).