Structural Styles and Dolomites Field Trip

Carlo  Doglioni

Structural Styles and Dolomites Field Trip

2008, Memorie Descrittive della CARTA GEOLOGICA d’ITALIA vol. LXXXII

Abstract

During the last 30 years, structural geology had relevant advances thanks to a wealth of data provided by geological and geophysical investigations. In this volume we propose a sort of handbook for basic tectonic features. We review the fundamental mechanics, and the most common features associated to brittle deformation. The main geometries and kinematics of compressional, transpressive, strike-slip, transtensive and extensional tectonic environments are presented. Moreover the migration of rupture along faults and the elementary evolution of diapirs are discussed. Few instances on how tectonics influences sedimentation, compaction, and the sedimentary architecture controls tectonics are highlighted. The structural features are then inserted in a wider geodynamic scenario in which the basic element is the basal decollement, with its depth, temperature, pressure and strain rate. The regional subsidence or uplift is combined with the growth rate of single tectonic features, which may locally increase or decrease the regional motions. This points to the computation for example of the fold total uplift, e.g., the uplift of an anticline minus the regional subsidence rate of the foredeep, that can even exceed the vertical growth rate of the fold. The combination with variable sedimentation rates may further differentiate the growth structures. The geodynamics of subduction zones and rift zones is discussed in the frame of the westward drift of the lithosphere, generating a worldwide asymmetry, which can be recognized as well as between the subduction zones in the Mediterranean realm. Unlike the Apennines, the Alps do not have a coaxial rifting in the hangingwall of the subduction, i.e., a back arc basin as the Tyrrhenian Sea. In this context, the Southern Alps and the Dolomites in particular, are rather the compressive retrobelt of the Alpine orogen, which is associated to the right-lateral transpressive subduction along the E-W trending segment of the belt. The Dolomites were undergoing rifting episodes during the Permo-Mesozoic. Fast subsidence rates during the Ladinian would support a backarc origin in the hangingwall of a W-directed subduction zone, possibly located in the present Pannonian Basin. The Dolomites are located half on the N-trending Trento Horst to the west, and the other half on the Belluno Graben to the east, both extensional features of Permo-Mesozoic age. The Trento Horst determined an undulation in the Alpine thrusting, being located in a recess since the early phases of shortening, generating left-lateral transpression along the western border (Giudicarie system), and right-lateral transpression on its eastern margin (Paleogene WSW-verging thrusts of the central-eastern Dolomites). The adjacent Lombard Graben (and basin) to the west, and the Belluno Basin to the east have rather been the seat of salients. The description of a five days field trip illustrates well exposed examples of tectonic features generated by the Mesozoic rifting and the later Cenozoic Alpine inversion that generated a classic thrust belt with imbricate fan geometry.

Figures (442)

ted to the generation of folds, stretching and thinning of layers, generation of planar (folia- tions) and linear (lineations) tectonic fabrics. It is however stressed that not necessarily folding of a sedimentary succession may be accommo- dated by plastic deformation mechanisms. More frequently, folding in foreland fold-and- thrust belts is accommodated both at meso- and micro-scales by brittle fracturing. In this case the term brittle-ductile deformation is adopted to indicate that geometries of structures may be similar to those occurring in ductile regimes but deformation is accommodated by brittle deformation mechanisms. The strenoth of rocks in the brittle reoime

racterised by very low strengths) and increases

Fig. 3 - The brittle-ductile transition depth (BDTD) is strongly dependent on the kind of structural regime active in a giver region, on the strain rate and on the thermal gradient. In this figure, the sensitivity of the BDTD on these parameters is schema- tically shown. In all the panels the lithology is kept constant. Areas characterised by thrust faulting are characterised by large1 brittle strength with respect to areas characterised by strike-slip and normal faulting. As a consequence, keeping constant the other controlling parameters, the brittle-ductile transition depth in contractional areas will be shallower than in strike-slip (inter- mediate depth) and extensional (deeper depth) areas. Strain rate strongly controls the ductile strength. Typical strain rates in deforming areas vary between 10° s' (slow deformation) and 10° s" (fast deformation). Increasing strain rates induce highet ductile rock strengths. This is mainly due to the fact that, for fast deformations, migration of lattice defects, recrystallization anc other plastic deformation mechanisms are not capable to keep the pace with deformation. Keeping constant the other controlling parameters, areas characterised by faster deformations will be characterised by deeper brittle-ductile transition depths. The ther- mal gradient of an area is controlled by the regional mantle heat flow, by the crustal radiogenic heat production (that is in turr controlled by crustal lithologies) and by local effects (e.g., magmatic intrusions, heat production generated along major regional faults). An area characterised by larger thermal gradient will be characterised, at a fixed depth, by higher temperatures with respect to areas characterised by low thermal gradients. Since ductile strength is inversely proportional to temperature, at the same depth areas characterised by high gradients will also be characterised by low ductile strength. Therefore, keeping constant the other controlling parameters, larger thermal gradients will be associated to shallower brittle-ductile transition depths.

Fig. 4 - Fault rocks associated to brittle (breccias, fault gouge and foliated/non-foliated cataclasites) and ductile (mylonites) structural settings (left panel). Notice that the transition from brittle to ductile behavior is rather ditributed. This is mainly due to the fact that the brittle ductile transition for different rock forming minerals occurs at different depths. In the right panel the intersection between the britt- le strength curves (for normal faults, N, for strike slip, SS and for reverse faults, R) and the ductile strength curves for quartz and feld- spar is shown. Quartz and feldspar are typical mineralogical components of upper crustal rocks, such as gneisses. In this example, at depths of about 15-18 km, quartz may deform plastically whereas feldspar may be still characterised by fracturing (after Sisson, 2001).

Fig. 5 - Ductile strength of different rocks and rock forming minerals calculated for a cold continental geothermal gradient. Rs, rock salt; Gr(w), wet (i.e., hydrated) granite; Gr, granite; Qz, quartz; Qz(w), wet quartz; Pg, plagioclase; Db, diabase; QzD, quartz-diorite; Ol, olivine. Calcite and limestones have strength intermediate between salt and quartz. Quartz, plagio- clase and olivine are the main constituents of upper crust, lower crust and lithospheric mantle. Notice that the ductile strength of the lithospere increases with depth. The minerals of the crust have lower strenght than the minerals of the mantle (after RANALLI & Murpuy, 1987).

Fig. 6 - Lithospheric scale rheological profiles are characterised by the alternance of brittle and ductile structural levels. Ductile layers may act as major detachment levels along deforming plate margins (e.g., rift zones and collisional belts). The profiles were calcula- ted for different geodynamic setting: C(s1), precambrian shields with 150 km thick lithosphere and unlayered (1.e., quartz.granitic) 4( km thick crust; C(s2), precambrian shields with 150 km thick lithosphere and layered (i.e., quartz-granitic over intermediate basic) 40 km thick crust; C(c1) continental convergence zones with 150 km thick lithosphere and unlayered 60 km thick crust; C(c2) continental convergence zones with 150 km thick lithosphere and layered 60 km thick crust; H(1), continental extensional zones with 50 km thick lithosphere and unlayered 30 km thick crust; H(2), continental extensional zones with 50 km thick lithosphere and layered 30 km thick crust; O, mature oceanic lithosphere, 75 km thick and with a 10 km thick basic crust. The lithospheric rheological profile is even more complicated if one considers sedimentary cover as well. As discussed in the previous figures, evaporites and limestones are charac- terised by ductile strengths lower than crystalline rocks. Additional brittle-ductile transitions may therefore occur within the sediment succession (after RANALLI & Murpuy, 1987). consequence, the lithological layering of the lithospere (sedimentary covers, upper crustal crystalline basement, lower crust, lithospheric mantle) will determine a rheological layering within the lithosphere. The rheological laye- ring of the lithosphere varies according to the type of lithosphere (continental vs. oceanic), to the thickness of the various layers and to the parameters controlling brittle and ductile strength (fig. 6). Typical lithospheric rheologi- cal profiles are characterised by the alternance of brittle and ductile levels and have a christ- mas-tree (or saw-tooth) shape. The occurrence of ductile low strength levels between stronger brittle layers is a major factor controlling the deep geometry of extensional (e.g., rift zones) and compressional areas (e.g., mountain belts). Weak ductile layers, that may be located both in the sedimentary covers and in the crustal crystalline rocks, will, in fact, behave as litho- spheric scale detachment levels.

Fig. 7 - Plumose structure on a fracture in lavas at Colcuc (Dolomites). Note on the left the center of nucleation of the fracture.

Fig. 9 - Generic Mohr diagram showing a composite Griffith- Coulomb failure envelope for intact rocks. The three shown cri- tical stress circles represent different failure modes: tensional fracturing, transitional-tensional fracturing (also called in lite- rature extensional-shear fracturing) and compressional failure (also called in literature compressional shear failure). Fig. 8 - Mohr diagram showing the state of effective stress at failure for various experiments (varying the confining and the axial load) with intact Fredrick Diabase. Each circle represents the state of stress at failure at a different mean stress ([o1+02+03]/3). The locus of stress states that bounds the field of stable and unstable stresses is called the Mohr envelo- pe (after Suppe, 1985).

Fig. 10 - Mohr-Coulomb failure criterion for isotropic intact rocks (left panel). The point of tangency of the Mohr circle represents the state of stress on the plane which is at an angle © from the o1 axis of the right panel. thickness of the fault gouge is direct function linear Amonton's law and, for faults characte- rised by cohesive strength, is a form of the Coulomb envelope (fig. 15). The slope of the Amonton envelope is controlled by the sli- ding coefficient of friction of the fault, which has been shown to have typical values as high as 0.85 in the shallow crust (BYERLEE, 1978). The frictional resistence along pre-existing planes is expressed by the Byerlee's law: T=0.850. The Byerlee's law is broadly litho- logy-insensitive. Fault surfaces that contain a layer of gouge also obey Byerlee's law and dis- play a frictional behavior similar to clean rock surfaces involving either strong and weak lithologies. The only silicatic minerals that display significant lower friction than Byerlee's law are montmorillonite, vermiculi- te and illite. Other platy minerals (e.g., chlori- te, kaolinite and serpentine) display normal frictional properties. The existence of near optimally oriented faults

Fig. 13 - Hydrothermal extensional vein systems associated to Anderson style fault systems (after SIBSON, 1990).

Fig. 12 - Mohr stress diagram showing the effects of pore fluid pressure on the state of stress of a rock volume. An increase of pore fluid pressure p lowers the principal stresses acting on the body and shifts the corresponding Mohr circle to the left. In this example, the dry rock is stable (i.e., it does not fracture since it does not reach the Mohr envelope) whereas the same rock volu- me, when subject to fluid pressure, may be unstable, i.e, cha- racterised by a Mohr circle tangent to the Mohr envelope. Oy is the normal stress.

Fig. 15 - Measurements of maximum friction in sandstone, limestone and graywake. The frictional resistence along pre-existing pla- nes is expressed by the Byerlee's law: T=0.850. Notice that, in the shallow crust (for lithostatic pressure lower than 0.2 GPa) the cohe- sion is null. The Byerlee's law is broadly lithology-insensitive and is valid for fault surfaces involving either strong and weak litholo- gies that may contain a layer of gouge as well. The only silicatic minerals that display significant lower friction than Byerlee's law are montmorillonite, vermiculite and illite (after BYERLEE, 1978). enters the belt at the front. If, at a specific point, the critical taper of a wedge is less than the critical value (because of erosion, for example), then the material will deform in more internal sectors (for example with the development of out of sequence thrusts) until the critical value is once again attained. At this

Fig. 16 - Fracturing in anisotropic rocks. The effect of the orientation of the plane of slaty cleavage on a fracture orientation is inve- stigated in compressive fracture experiments on Martinsburg Slate. The orientation of the forming fractures is generally influenced by the foliation. For inclination of the slaty cleavage to the o1 of as much as 45°, the fracture always lay parallel to the cleavage. Only in cases for which the plane of anisotropy is nearly perpendicular to the o1 direction, anisotropy is unimportant in localizing fractures (after SuppE, 1985).

Fig. 18 - Maximum length (1), as a function of fluid pressure ratio (A=Pf/P1, where Pf is the fluid pressure and P1 is the litho- static pressure) of a rectangular thrust sheet that can be pushed along a horizontal decollement without fracturing. ) is the fluid pressure within the thrust sheet and Af is the fluid pressure along the fault (after SUPPE, 1985). Fig. 17 - Fault gouge thickness (T) vs. total slip (D) for faults mainly developed in crystalline rocks (after SCHOLZ, 1990).

thrust or the faults occur along an unconfor- however, transterred also by dittuse strain. In

where displacement dies to zero; 3) detach-

Fig. 22 - Upper panel. Glarus thrust, Swiss Alps. The dark rocks above the thrust (the Permian Verrucano Fm.) have been emplaced on the lighter Mesozoic limestones. Lower panel. Although the age and nature of the lithologies was recognised since the early beginning, the tectonic nature of the contact (by thrust fault) was only later recognised. To explain the regional geology of the Glarus area, Arnold Escher (1807-1872) and his student Albert Heim (1849-1937) proposed a strange "double fold" structure. Heim's clear descriptions in his classic 1878 text on mountain building led, in 1884, to the re-interpretation of the Glarus structure as a single, southward-derived thrust sheet by Marcel Bertrand (1847-1907) - although Bertrand had never visited the area. (after Butler, 2006; earth. leeds.ac.uk/tec- tonics/thrust_tectonics/index.htm).

Fig. 23 - Possible 2-D geometries of thrust faults. The thrust can either reach the Earth's surface or not (blind faults). Note the staircase (ramp and flat) geometry of the fault plane. Redrawn after McCLay (1992).

Fig. 24 - Generally older thrusts are carried on younger thrusts (in sequence thrusting). Higher older thrusts become folded as younger thrusts climb ramps. The opposite occurs when out-of-sequence thrusts develop. Redrawn after McCray (1992).

Fig. 30 - a) The thrust faults of the Foppo Alto area are severely misoriented (far too steep; see previous figures) for fault reactivation in a sublithostatic fluid pressure regime. Thrust motion likely started when the faults were less steep and the faults were progressively rotated up to the present day dip by domino tilting. When the faults became inclined beyond the fault lock-up angle, no further thru- sting was accommodated along them. At later stages regional shortening was accommodated by newly formed lower angle shear pla- nes (dipping around 30°-40°), consistently with predictions from fault mechanics. b) Dip distribution for the thrust faults and for the later (D4) shear planes of the Foppo Alto area (after CARMINATI & SILETTO, 2005). Fig. 29 - Plot of differential stresses needed in order to reactivate thrust faults with different dips at a depth 10 km for variable pore fluid pressure ratios (A=Pf/P1, where Pf is the fluid pressure and PI is the lithostatic pressure). The coefficient of friction is 0.35. Notice that, even for very high fluid pressure, thrust faulting along planes more inclined than ca. 70° would need far too high differential stresses. Such faults are therefore severely misoriented to be re-sheared in a compressional regime (after CARMINATI & SILETTO, 2005).

Fig. 33 - Thrust systems may be extremely complicated and all the structures so far described may coexist in fold-and-thrust belts (after DE PAor, 1988). Fig. 32 - The different spacing of thrust fault ramps in the lower panel may be misinterpreted as the result of a variable spacing of fault nucleation points. When the section is retro-deformed, in this case the original spacing was constant, which is not the rule. The varia- ble spacing in the deformed section is related to the variable displacement along the faults.

‘ig. 37 - The spacing of thrust faults is controlled by the thickness of the foreland sedimentary cover above the basal decollement. The larger the thickness, the larger the spacing (after Huiql, et alii 1992).

Fig. 36 - The progressive development of fold- and-thrust belts has been studied extensively (since the pioneering work of CADELL 1889, above picture) with analogue models that simulate thrust faults formation pushing laterally a multilayer of sand (plus additional materials). Notice in the left panel the foreland propagation of thrust faulting. When new thrust faults form, older thrusts are abandoned, rotated and transported passively toward the foreland (after Huiqi et alii, 1992).

Fig. 38 - The geometry of the wedge is controlled by the friction along the basal decollement. Highly frictional decollements are asso- ciated to narrow belts with a very steep topographic slope. Low friction decollements are associated to wide and flat mountain belts. In experiments a-c, the friction coefficient of the decollement is 0.55, 0.47 and 0.37 respectively. Compare these geometries with the fol- lowing two figures.

Fig. 39 - Three cross sections through the Zagros show different fold-and-thrust geometries. The basal decollement of sections A and C runs through the Cambrian Hormoz Fm. (salt) and the chain is wide. Along section B no salt occurs along the basal decollement and the surface slope is steeper and the belt is narrower with thrust faults characterised by larger displacement (after MCQUARRIE, 2004).

Fig. 41 - Cleavage formation is often associated to folding. When cleavage is subparallel to fold axial planes it is termed axial plane cleavage. However, the axial plane cleavage is often refracted between competent and incompetent layers and the spacing is dependent on lithology (larger in more competent lithologies). Cleavage may represent an important source for rock permeability. Modified after Twiss & Moores (1992).

Fig. 42 - Folded volcaniclastic rocks of the Permian Collio Fm. (the original stratigraphic layers are indicated by the black line). Notice the development of a strong axial plane cleavage (red lines) that is unevenly distributed. Some layers (normally those characterised by higher percentages of phyllosilicates, such as micas for example) are characterised by a sharp cleavage, whereas in other layers (more competent) cleavage is absent. In some layers, cleavage may have obliterated continuously the original strata. Val Belviso, Central Southern Alps.

Fig. 43 - Microphotograph of an axial plane cleavage developed in marly limestones from the Alpi Apuane, Central Italy. Notice tha one of the main processes controlling cleavage development in this sample was pressure solution.

Fig. 44 - In terrains involved in more than one phase of folding, different generations of cleavage may develop. For example in these Cretaceous marly limestones from the Mt. Marguareis area (Ligurian Brianconnais, Western Alps) two generations of folds and corre- sponding axial plane cleavages (labelled as S1 and S2) can be observed. The original stratigraphic surfaces are labelled SO.

Pigs. TO - Uldssl€ PClauONs DELWCCH UNUSt sadpe, prOpdedllOn dalla 101d GEVElOpMent VOUPFE, LAGI, LAGI). Lt DOU Cas@s Ul© DACK- HTT of the fold is parallel to the footwall ramp. Fault-bend folds are generated when the hangingwall of a thrust sheet passes over a step ped (ramp-and-flat) thrust fault. Above each kink of the thrust fault surface folds are generated to accommodate slip. Fault-propagation folds are generated by the folding strain at the tip of the ramp of a blind thrust, where displacement dies to zero Above this point, an asymmetric fold with a steep or overturned forelimb (indicating the vergence of the thrust) is formed. With th progressive upward migration of the thrust fault tip, the fold may be dissected and in part transported in the hangingwall of the thrus sheet, with a relict footwall syncline. The forelimb is less inclined in the fault-bend fold, whereas it is often overturned and subject t elongation and flattening in fault-propagation folds.

Fig. 47 - The trishear model (ERSLEV, 1991, ALLMENDINGER, 1998) explains most of the deformation occurring at the tip line of faults. httn://www geo. cornell edu/RWA /trishear/default htm]

Fig. 51 - Fault trajectories may be characterised by two or more bends with the same concavity or convexity. These features produce multi-bend fault-bend folds. After SHAW et alii (2005) AAPGO©, reprinted by permission of the AAPG whose permis- sion is required for further use. Fig. 50 - Fault-bend fold in the Dachstein Limestone (Da) of the southern cliff of the Piz Boe, Sella Massif, Dolomites. Note the fore- limb inclined toward the vergence. Dolomia Principale (DP). The cliff is about 20 m high.

Fig. 53 - Growth-fold associated to fault bending with low sedimentation rates (i.e., sedimentation rates slower or equal to uplift rates). After SHAW et alii (2005) AAPG©, reprinted by permission of the AAPG whose permission is required for further use. Fig. 52 - Fold interference structures form when two or more fault-bend kink bends intersect. This produces for example anticlines over synclines. After SHAW et alii (2005) AAPGO©, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 55 - Growth fold characterised by sedimentation exceeding uplift. Notice that the growth strata, bounded by two seismi reflections, thin towards the structural high. After SHAW et alii (2005) AAPG©, reprinted by permission of the AAPG whose per mission is required for further use.

Fig. 56 - Growth fold characterised by uplift exceeding sedimentation. Growth-strata typically thin toward and onlap the structural high. Growth strata do not occur on the crest of the anticline. After SHAW et alii (2005) AAPGO©, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 57 - Seismic section through a fault-propagation fold in the Tarim Basin, China. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 58 - Fault-propagation fold, southeastern cliff of Col Bechei, Dolomites. This type of structure is the most common for britt le regime in the orogens worldwide.

Fig. 59 - Natural section of the Mt. Chiampon, near Venzone, Friuli, eastern Southern Alps. Two major thrusts (Musi Thrust, eastward prolongation of the Valsugana Thrust, and Periadriatic Thrust) contain an intervening folded and internally sheared slice of Mesozoic rocks (platform facies up to Liassic, and deep water up to the Cretaceous).

Fig. 60 - Detail of the previous figure. The upper decollement is older and has been tilted and folded by the deeper decollements. Note the gradual counter-clockwise rotation of the axial planes of the detached folds due to piggy back forward propagation of the imbricate fan.

Fig. 62 - Cartoon showing the possible breakthrough structures associated to fault-propagations folds. a) and b) decollement breakth- rough; c) synclinal breakthrough; d) anticlinal breakthrough; e) high-angle-forelimb breakthrough; f) low-angle breakthrough. After SHAW et alii (2005) AAPG©, reprinted by permission of the AAPG whose permission is required for further use. Fig. 61 - Fault-propagation fold in a 25 cm wide specimen of the Lagonegro Jurassic Scisti Silicei, Southern Apennines. Notice the transfer of shortening from the thrust to the fold. Unlike the fault-bend fold mechanism, the forelimb tends to be vertical or overturned. The backlimb has the inclination of the ramp.

Fig. 63 - Fault-propagation fold in the Siusi Member, within the olistholith of Werfen Fm at Varda (Dolomites). Note the partitioning of the strain associated to the shortening. In the transfer zone thrusting gradually decreases and passes into folding. In the lower part, small backthrusts disappear at depth into vertical slaty cleavage accommodating the shortening.

Fig. 64 - Seismic section showing a forelimb breakthrough pattern. After SHAW et alii (2005) AAPG©, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 65 - Fault-propagation fold at the hinge of a chevron fold in the Ladinian basinal Livinallongo Fm. 1 km west c Livinallongo del Col di Lana, Dolomites.

Fig. 68 - Growth fault-propagation folding with sedimentation rates greater than uplift rates. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use. Fig. 67 - Growth fault-propagation folding with sedimentation rates slower than uplift rates. Notice that in submarine conditions the crest of the growing anticline may be characterised by non-deposition, wherear it may be eroded in subaerial conditions. After SHAW ef alii (2005) AAPG©, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 69 - Growth fault-propagation fold in Pleistocene conglomerates. Alianello, Southern Apennines. Notice the thinning of syn- tectonic sediments toward the crest of the anticline.

Fig. 72 - Detachment fold cropping out in the Opal Mt., Canadian Rocky Mountains. Modified after Bons, 2007 (http://homepages.uni- tuebingen.de/paul.bons/paul/lectures/structure/index.html). Fig. 71 - Seismic line of a detachment fold in the Gulf of Mexico. After SHAW et alii (2005) AAPGO, reprinted by permission of th AAPG whose permission is required for further use.

Fig. 74 - Models of growth detachment folds. Different ratios of sedimentation vs. uplift rates (increasing toward the bottom of the figu- re) are shown. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use. Fig. 73 - Possible geometries of detachment folds. After SHAW et alii (2005) AAPGO©, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 75 - Chevron folds in the Livinallongo Fm. Note the decoupling surface above where the thickness of the beds increase and the underlying tight folds disappear. The fold axis is NNW-trending. Gasoline station, Pieve di Livinallongo.

Fig. 76 - Example of triangle zone at the Alpine front in Bavaria (after BERGE & VEAL, 2005)

Fig. 78 - The décollement depth (Z) is the fundamental information for tectonic reconstructions. The depth can be estimated once the length of a marker bed (L) and the uplifted area (A) is measured. The uplifted area is equal to the shortened area. The area (A) divided by the difference between the original marker length (L) and the final shortened length (L') gives the calculated depth, after CHAMBERLIN, (1910) and MiTRA & NAMSON (1989).

Fig. 79 - Folds in the Zagros. The large scale wavelength and amplitude of the folds indicate a deep decollement (10-14 km), which is located in early Paleozoic salt. Note three diapiric intrusions (darker rocks). USGS-Nasa picture.

Fig. 80 - Seismic line through the Nankai Trough, Japan, showing the active front of the accretionary wedge associated to the Nankai subduction zone. Notice the ramps generating from a nearly subhorizontal basal decollement. Fault-bend folds accommodate the dis- placement along the ramp-flat thrust faults. After SHAW et alii (2005) AAPG©, reprinted by permission of the AAPG whose permis- sion is required for further use.

‘ig. 81 - Seismic sections through the front of the Northern Apennines beneath the Po Plain. Notice that the two central growth fault- ropagation folds are characterised by sedimentation faster than uplift (after Preri, 1983).

Fig. 82 - Structural wedges contain two connected fault segments that bound a triangular block. The two faults may be either two ramps or aramp and a detachment. Structural wedges nucleate in agreement with the conjugate fault theory. Once nucleated, slip occurs along both faults allowing propagation of the wedge and determining the folding of the hangingwall. The lower thrust is named roof-thrust whereas the upper thrust is the back- (or roof-) thrust. At large scales wedges are referred to as triangle zones. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 83 - Seismic line showing a triangle zone imaged in the Alberta Foothills, Canada. The various structures highlighted by numbers are interpreted in the inset. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 84 - Triangle zone in the Liassic Calcari Grigi, Tofana II, Dolomites. The outcrop is about 20 m wide.

Fig. 85 - WSW-verging chevron folds in Liassic Calcari Grigi in the hangingwall of the basal thrust of the Civetta klippe, Dolomites

Fig. 87 - Detail of the central part of the previous picture. Triangle structure in Liassic Calcari Grigi, Van Dei Sass, Civetta Massif. The backthrust on the right side is folded and tilted.

Fig. 88 - Triangle structure developed in Carboniferous Rocks. Front Ranges of the Canadian Rockies. After SHAW et alii (2005) AAPGO, reprinted by permission of the AAPG whose permission is required for further use.

Fig. 89 - 2D geometric classification of normal faults. The clas- sification is based on the effects that faults have on beds and on other faults.

lopment of a wedge geometry of hangingwall

Fig. 92 - Domino set of planar rotational normal faults, Liassic Medolo Fm, Lake Maggiore, Lombardy. The faults are decoupled olong the marly intrabed layers. Courtesy of Daniel Bernoulli.

Fig. 94 - Listric fault cropping out in the Sinai Peninsula, Egypt (after KHALIL, 1998).

Fig. 95 - Seismic line showing a rotational, mainly planar domino set of extensional faults. Gulf of Mexico.

Fig. 97 - Listric faults are characterised, in the hangingwall, by the occurrence of rollover anticlines. Such structures accommodate the slip along the curved fault plane and are possibly characterised at their top by a crestal collapse, with the development of graben like structures.

Fig. 99 - Analogue modeling showing the development of a rollover anticline and of the associated crestal collapse grabe: (after PLATT et alii, 1993).

Fig. 100 - Synsedimentary listric fault. The syn-rift strata show in the hangingwall a wedge-like thickening towards the fault.

Fig. 102 - Seismic section showing a listric fault associated with a rollover anticline characterised by crestal collapse (after XIAO & SuppE, 1992).

Fig. 103 - Seismic line and interpretation showing syn-sedimentary listric faults from the Cascadia subduction zone (after MCNEILL et alii, 1997,

Fig. 105 - Growth normal fault in the Devonian Hornelen Basin, Norway (after DAviEs & REYNOLDS, 1996)

Fig. 106 - Syn-sedimentary normal faults in lacustrine sediments of the Pianico Basin, Lombardy. Courtesy of Fabrizio Berra.

Fig. 109 - The development of mainly strike-slip (transfer) faults may accommodate normal fault segmentation. In this case the displa- cement between two normal faults is transmitted by hard linkage. The panel to the right shows a tectonic map from the Viking graben showing extensional faults offset by transfer faults (after GipBs, 1984 and 1990).

Fig. 108 - Seismic section showing a normal fault-propagation fold developed in the Usungu Flats, Tanzania (after MORLEY, 2002).

Fig. 110 - Hard linkage structures may link normal faults or semi-grabens with opposing or common dips. After MILANI & DAVISON (1988),

Fig. 113 - Relay ramps develop at very different scales. In this case kilometre-scale relay ramps are imaged by an aerial photograph (the inset is the interpretation) taken in the Canyonlands National Park, Utah, USA. At the terminations of individual grabens, the main faults branch into a series of splays that form relay ramps between overlapping, like-dipping faults (after MCCLAy et alii 2002).

Fig. 112 - Relay ramps develop at very different scales. In this case decimeter-scale relay ramps are exposed in limestones, East Quantoxhead, Someset, UK (after PEACocK & SANDERSON, 1994).

Fig. 114 - Conceptual models of displacement transfer in rift systems. (a) Synoptic model of a low-strain intracontinental rift with along-axis segmentation. Individual half grabens are separated by soft-linked accommodation zones formed by overlapping fault seg- ments. (b) Synoptic model of hardlinked, strike-slip, rift transfer fault system (after MCCLAy et alii, 2002).

Fig. 117 - Typical geometry of strike-slip fault systems. EE SS ———— <_.eee tractional duplexes. Strike-slip faults are often segmented and fault stepovers occur when single fault segments are arranged in en-echelon geometry. Stepovers of fault segments may drive to the development of contractional areas (restrai- ning stepovers) and extensional areas (relea- sing stepovers). The most common structures associated to releasing stepovers are pull-apart basins (figs. 124 -128). At the tips of major stri- ke-slip faults displacements die to zero. Slip is here accommodated by the formation of (both extensional or compressional) faults fans (fig. 129).

Fig. 119 - Undulated strike-slip faults give rise to local extensional (releasing bends; upper panel) or compressional (restraining bends; lower panel) zones. Releasing bends are associated to negative flower structures, whereas restraining bends to positive flower structu- res. In map view, these features are also called strike-slip duplexes. Modified after Twiss & Moores (1992). Fig. 118 - Structures associated to strike-slip faults. PDZ (prin- cipal displacement zone) or Y-Shear is the principal strike-slip zone. R1 (synthetic Riedel, at 15-20° to the PDZ) and R2 (anti- thetic Riedel at 70-75° to the PDZ) shears develop in response to Navier-Coulomb shear failure with respect to 01 and 03. P syntheic shear develops at low angle (15°) to the PDZ as a con- sequence of the rotation of the principal stress axes at the tips of R1 shears as faulting develops. Minor thrust and normal faulting also may develop. Fig. 120 - Strike slip segmentation and en-echelon stepover geo- metries may cause local shortening (restraining stepovers) or stretching (releasing stepovers). Stretching areas associated to releasing stepovers are called pull-apart basins, whereas the opposite restraining stepovers generate push-ups.

Fig. 122 - Negative flower structure related to sinistral strike-slip alpine tectonics in the SettSass Massif (Dolomites)

Fig. 123 - Transtensional N-S trendind left-lateral fault in the Caotico Eterogeneo. Slickensides are more developed in conglomerates thar in more homogeneous rocks. The striations have a pitch of 12° left (northward). The left-lateral sense of motion is indicated by the step: in the fault plane, and by the crystals of calcite grown in their shadow. Road between Caprile and Selva di Cadore (Dolomites). Casey Moore for scale.

Fig. 126 - Stages of formation of a pull-apart basin (after KEAREY & VINE, 1992).

Fig. 127 - Along the El-Salvador fault zone, a major strike-slip fault system, the Rio Lempa pull-apart basin, developed in Quaternan times. (a) Digital elevation model of the region. (b) Line-drawing of the active structures forming the El Salvador Fault Zone (ESFZ in the area, with also reported palaeostress orientations. Inset shows the interpretation of the regional fault pattern. (c) Geological may superimposed on a DEM of the region. Arrows indicate the fault trace. After Corti et alii (2005).

Fig. 128 - Terminology and stress state for bends and stepovers along strike-slip fault systems (after Twiss & Moores, 1992).

depths and early altershocks are expected to Fig. 131 - Cross-section of the Friuli 1976 earthquake sequen- ces occurred in a compressive environment. The seismicity occurred in two main episodes, May 6 and September 15, with magnitude 6.4 and 6.1 respectively. The focal mechanisms are mainly compressional, with a slight transpressional component. They were associated with N-dipping thrust planes in the eastern part of the Southern Alps, at the intersection with the Dinarides thrust belt in north-east Italy. The maximum energy released by the aftershocks of the May 6 earthquake was very shallow, between 2 and 8 km. The highest energy release densi- ty has been evaluated to be located at a depth of about 4 km. The epicentres of the later September quakes and related aftershocks were about 8 km northward. The September hypocentres were deeper (down to 16-17 km) and characterised by a maximum energy release density at about 5-6 km. The timing of groups of events clearly shows a deepening of seismic activity with time. Modified after CARMINATI et alii (2004). Fig. 130 - Cross-section of the Irpinia 1980 normal fault-related earthquake sequence. The largest red square is the mainshock. The blue squares are aftershocks. The earthquake was generated by two normal faults dipping north-eastward and one conjugate, and was charac- terised by Ms 6.9. The earthquake showed the largest moment release at a depth of 8-13 km and a possible nucleation at about 10 km. Most of the aftershocks were generated in the hanging wall of the main fault, between this and an antithetic normal fault dipping to the south-west. All the aftershocks occurred at depths (3-10 km) shallower than the mainshock. Modified after CARMINATI et alii (2004).

Fig. 132 - Left, sketch and Mohr diagram of the variation of o1 (vertical load) and o3 (horizontal load) along normal faults. Right, sketch and Mohr diagrams of the variation of 03 (vertical load) and o1 (horizontal load) along thrust faults. For the sake of simplicity orizontal stress is assumed to remain constant with depth. It is shown in the next figure that such a rough assumption does not hamper the validity of the model. Along thrust faults, shallow initial ruptures and downward migration of seismicity is more likely. On the con- trary, along normal faults, rupture should generally occur at deeper levels and aftershocks are expected to be shallower.

Fig. 133 - Left, sketch and Mohr diagram of the variation of 01 (vertical load) and 03 (horizontal load) along normal faults. Right sketch and Mohr diagrams of the variation of 03 (vertical load) and o1 (horizontal load) along thrust faults. The horizontal stres: increases with depth according to the Poisson effect. Along thrust faults, shallow initial ruptures and downward migration of seismi city is more likely. On the contrary, along normal faults, rupture should generally occur at deeper levels and aftershocks are expec: ted to be shallower. Modified after CARMINATI et alii (2004).

Fig. 134 - Classic model of salt (red layer) evolution when loa- ded. First it forms a pillow uplifting the overlying cover (2). Then the diapiric piercing starts (3) and the mother salt layer is eventually entirely evacuated (4) to generate a subsidence around the diapir and producing a basin surrounding the diapir (rim syncline). Note the yellow layer which is first uplifted (pil- low stage) determining a thinning of the growth (gray) strata. Then it subsides (diapiric stage) forming an overlying thicke- ning of the cover. An unconformity separates the two stages at the base of the azure (after TRUSHEIM, 1960).

Fig. 135 - Example of diapir from the Permian Zechstein salt in northern Germany after Monr et alii (2006). Note the two major unconformities, one at the base of the Cretaceous (K), and a deeper one at the base of the kmS-kmW, marking the transition betwe- en the pillow stage and related uplift producing erosion, followed by the rim syncline which occurred with the evacuation of the salt from the mother layer.

chment level characterised by the occurrence

Fig. 139 - Gravity salt tectonics in the Kwanza Basin, offshore Angola. Notice the extensional tectonics in the internal sectors (close to the continent; raft domain) with listric normal faults rooted in a shallow detachment level characterised by the occurrence of remnant bodies of salt. The central part of the section is characterised by wide and numerous diapirs, while the external part is a compressional domain characterised by thrusts and folds. Shortening is due to accumulation of sliding sedimentary cover at the transition between the continental ramp and the abyssal plain (after Fort et alii, 2004).

Fig. 141 - Seismic sections from the diapir domain of the Kwanza Basin, offshore Angola. After HUDECK & JACKSON (2004).

Fig. 143 - Salt diapirs with crestal collapse normal faults from the US Gulf coast (after SUNWALL ef alii, 1983) Fig. 142 - Seismic section from the salt nappe domain of the Kwanza basin, offsho- re Angola. Notice the development of folds and thrust faults. The bottom panel is an enlargment of box c of the upper panel (after HUDECK & JACKSON, 2004).

Fig. 146 - Same sections as in the previous figure with some interpretation (modified after TRANSALP, 2002; DOGLIONI et alii, 1999c) Note the Alpine double vergence versus the single Apennines vergence, where the prism is even deeper than the foreland and followed by extension to the left. L, lower plate, H, subduction hinge, U, upper plate. In the Alps H migrated toward the upper plate, whereas 1 moves toward the lower plate in the Apennines. The Adriatic plate is the upper plate U in the Alps, while it is the lower plate L in the Apennines. The E-W segment of the Alps formed under right-lateral transpression (After DOGLIONI et alii, 2007). topmost basement) because the accretionary

Fig. 147 - Schematic sections showing how in an Alpine setting, the subduction rate is decreased by the migration of the hinge H toward the upper plate U, and the orogen in the final collisional stage is composed both by the upper and lower plate L rocks. In the opposed Apenninic setting, the subduction rate is rather increased by the migration of H away from U, and the accretionary prism is made of shallow rocks of the lower plate. Note also the shallower asthenosphere in the hangingwall, which is typical of W-directed subduction zones (after DOGLIONI et alii, 2007). W-directed subduction zones on Earth appear as short lived. If the Marianas or Japan subduc-

The aggradation of a carbonate platform is controlled by subsidence and carbonate pro-

Fig. 148 - System of grabens with N-S direction in Dogger sediments (synchronous to the deposition of the Fonzaso and Calcare de Vajont Fms.) cropping out along the northern flank of Mt. Mangart (area of Tarvisio). It can be considered as a field example of thi Mesozoic Southalpine extensional tectonics. Generally, inversion 1s considered 1n two dimensions (e.g., a graben inverted to produce a pop-up). However, the direction of compres- sion can vary widely with respect to the prece- ding extension. In the Southern Alps, for example, alpine compression (N0°-30°W) is oblique to Mesozoic extensional paleostructu- res (i.e., extensional growth faults oriented N10°E-N10°W) (fig. 148). Extensional paleo- faults dipping to the east were easily cut and transported passively along thrust faults, whe- reas those dipping to the west were easily re- sheared as sinistral transpressional faults. As already mentioned, inversion (or re-shear) of a fault occurs when peculiar conditions are respected (e.g., low angle both for strike and dip between the thrust fault and the pre-exi- sting normal fault; high fluid pressure). The shortening of a platform-basin system can pro- duce different interference geometries for dif- ferent directions of compression relative to the direction of the platform margin. The concept fault occurs when peculiar conditions are

Fig. 149 - The Mesozoic horts and grabens and facies transitions controlled the geometry of the Venetian Alps fold-and-thrust belt, producing undulations of the structures that are characterised by non-cylindrical shape. Transpressional fault systems and transfer faults are localised along inherited Mesozoic extensional faults and along thickness and facies variations of the sedimentary cover. The shaded pink zone represents the morphological and structural relief of the area. The darker colors are toward the more internal thrust sheets. Note that such a relief diminishes rapidly from the Trento Platform towards the east where compressional tectonics affected the sediments of the Belluno Basin. At the eastern margin of the basin, a major sinistral transpressional zone occurs at the intersection with the Friuli carbonate platform. The structural relief is lower but more diffuse within the Belluno Basin which is richer in shaly content, allowing lower friction in the decollement zones, and forming a salient of the thrust belt front (after DOGLIONI, 1992). In the salient, ramp distance is wider and the topography is lower. totally compensated by the eustatic rise. As a result the relative sea-level may remain appro- ximately constant along the flanks of a growth anticline and erosion may be minimal. The subsequent high stand represents an interme- diate situation, in which the structural uplift is only in part (or not at all) compensated by the eustatic rise. Consequently also erosion is intermediate between that occurring in low stand and transgressive phases. Therefore, each fold has its own tectonic and eustatic history. The resulting syntectonic stratigraphy is there- fore highly dependent on when, where and

Fig. 150 - Satellite image of the Venetian Alps. Notice that the dimension of the morphological high of the Asiago Plateau (A), cor- responding to the Mesozoic Trento Platform, drastically reduces to the east within the Belluno Basin, broadly correspondent with the Belluno syncline (S) located between Feltre (F) and Belluno (B). To the right the undulation of the Mt. Grappa-Visentin Anticline can be observed. This anticline developed in a transpressional system, at the intersection between the Belluno Basin and the Friuli Platform. V, Vittorio Veneto. tectonic Pulse. The geometry of sediments onlapping normal fault escarpments or the flanks of anticlines, is normally considered to be controlled by the relationship between tectonics and sedimenta- tion but it may be additionally controlled by differential sediment compaction. For exam- ple, in the Apennines the basinal strata onlap- ping platform sediments along Mesozoic nor- mal fault paleoescarpments commonly dip considerably (as much as 20°-30°) toward the basin when platform strata are retrodeformed to horizontal so to eliminate the geometrical effects of Tertiary compressional tectonics. These geometries are usually interpreted, in the field and in seismic lines, as the effects of normal drag along synsedimentary faults. Similarly, basinward dip affecting clastic sedi- ments onlapping anticlines flanks in the Po Plain has been interpreted as the effect of syn- sedimentary growth of the anticline. Although

Fig. 153 - Schematic section of a carbonate platform (violet) and of the adjacent basin covered and levelled by later sediments. The mar- gin of the platform becomes the area where compressional stresses are focussed. For this reason, the chance that ramp geometries of the thrust faults may form in this inherited structure is rather high. The klippen (K) of the Dolomites are interpreted as remaining parts of the hanging-wall of thrust faults similar to that shown in this figure. The evolution of the belt will be therefore registered by punctual information provided by the unconformities. The dip of the base- ment monocline varies between 0° and 5° in belts associated to subductions that follow the flow of the mantle (1.e., directed to the E-NE), and between 5° and 10° in the belts associated to subductions contrasting the mantle flow (i.e., directed to the W). In this latter case, cha-

Fig. 152 - Potential control of the stratigraphic succession on the localization of ramps and flats (decollement surfaces) of thrust faults in the Dolomites region. The basement includes the Gardena Sandstone, the older Permian porphyrites and the underlying Hercynian crystalline basement. Notice the occurrence of margins of Ladinian and Carnian carbonate platforms (Marmolada Limestone and Cassian Dolomite respectively) that can be the loci of thrust ramps. The Bellerophon Fm (Upper Permian), consisting mainly of evapo- ritic sediments such as gypsum, shales and marls, is characterised by flat geometries of thrust faults, similarly to other relatively less competent formations (e.g., Andraz Horizon of the Werfen Fm and Raibl Fm).

Fig. 155 - Northern wall of the Croda del Vallon Bianco (Ampezzo Dolomites), consisting of Liassic Calcari Grigi in which a W-ver- ging thrust cuts a pre-existing extensional fault of Upper Jurassic - Lower Cretaceous? age. Fig. 154 - The Sasso di Richthofen (central Dolomites) is a prograding edge of the Lower Cassian Dolomite (lower Carnian, D1) that was doubled by a thrust that partially used as a ramp the clinostratification of the megabreccias. Notice within the San Cassiano Fm the thinning of the body of megabreccias towards the right that onlaps the same body. D2, Upper Cassian Dolomite; S1, Lower San Cassiano Fm; $2, Upper San Cassiano Fm.

Fig. 156 - WSW-verging thrust of the Remeda Rossa in the Ampezzo Plateau. The thrust surface is sub-horizontal, the hanging-wal consisting of folded and partly overturned Calcari Grigi (G) and the footwall consisting of Cretaceous marls (K; in the meadows) extremely deformed by the shear zone.

Fig. 157 - Detail of the previous figure. In the hanging-wall of the Remeda Rossa thrust the Calcari Grigi sediments are overturned along the broken flank of a fault-propagation fold. In these sediments, a shear zone can be observed. If the sediments are retro-deformed to the pre-overturning position, the shear zone is interpretable as a syn-sedimentary extensional fault. Notice that the fault is sutured at the bottom of the picture. It may, however, be argued that the flexural slip necessary to accommodate the fold formation may have oblite- rated such an inherited shear zone. In this case the shear zone could be interpreted as an antithetic Riedel fault of the main thrust plane.

lateral ramp along the inherited hinge zone

Fig. 160 - The sinistral transpression along the Giudicarie fault system represents the strongest structural undulation of the Alps anc developed as a sinistral transpressional reactivation of a system of extensional faults dipping to the west (one of the major fault is th« Ballino Line; CASTELLARIN, 1972). This undulation developed at the border between the Lombardian Basin to the west and the Trent Platform to the east (DOGLIONI & BOSELLINI, 1987).

Fig. 162 - A thrust obliquely affecting a platform-basin margin will produce an apparent dextral displacement. For instance the Trento Platform was thrust towards southeast over the Belluno Basin in the Venetian Prealps along the Belluno Thrust, due to the southalpine N20°-30°W compression oblique to the Mesozoic N-S trending margin.

Fig. 163 - The front of the fold-and-thrust belt is characterised by clastic syntectonic deposition. With the progressive growth of the belt, the molassic deposits tend to onlap the flank of the frontal growth anticline. If the frontal anticline is cylindrical, the resulting unconformity will interest only the dip angle. In anticlines characterised by axial undulations, the resulting unconformity will occur both along dip and strike.

Fig. 164 - Three tectonic cycles may occur during three different segments of a eustatic curve. The depositional sequences may thus b« characterised by different interference patterns between tectonics and eustatism. The fault-propagation folds of an accretionary prism produce episodic and localised uplifts of the belt, which migrate regularly towards the foreland together with the foredeep basin and the depositional sequences associated to the growth of any anticline.

Fig. 165 - The growth of an anticline may be episodic or constant (like the slow steady-state plate motions). In this figure it is show a scenario characterised by an uplift rate of 1 mm/yr. The uplift is superimposed to a eustatic cycle of third order, represented to th right. During the low stand, with a eustatic sea level fall of 1 mm/yr, the relative sea-level fall at the growing anticline will occur a rates of 2 mm/yr, producing a significant erosion of the anticline. During a transgressional stage with rates of 1 mm/yr, the tectoni uplift will be compensated by the eustatic rise and the relative sea-level will remain constant, inducing minimum rates of erosion 0 the fold. The assumed uplift and eustatic rates are hypothetical. However, they may help in understanding the problem of the rela tionship between tectonics and sequence stratigraphy.

Fig. 167 - This figure shows the results of numerical models of compaction of the basinal sediments onlapping a normal fault paleoe- scarpment. Such results may be, however, extended to sediments onlapping the flanks of anticlines in compressional settings. Notice that the geometry and dip of basinal sediments vary consistently with the variation of the sediment nature and of the fault dip. (1) (A Dip and (B) geometry predicted for a 1200-m-thick succession of chalk onlapping an escarpment dipping 45°. (2) Dip predicted fot 1200-m-thick chalk (black), shale (red), and sandstone (blue) successions onlapping an escarpment dipping 45°. (3) (A, C) Dip and (B. D) geometry predicted for chalk section (1200 m and 600 m thick, respectively) onlapping escarpments with dips ranging between 25‘ and 65°. In A and C, dip is referred to level 3. Colors indicate simulations with different escarpment dips: cyan, 25°; brown, 35°; blue. 45°; green, 55°, red, 65° (after CARMINATI & SANTANTONIO, 2005). Fig. 166 - Panoramic view of platform-to-basin transition at Castiglione (Sabina region, Central Apennines). Notice that the onlapping basinal sediments are at an angle of approximately 20° with respect to the sediments deposited onto the structural high. Such an angle is interpreted as the result of differential compaction affecting basinal sediments in the hangingwall of the paleoescarpment. PCP- pela- gic carbonate platforms, (after CARMINATI & SANTANTONIO, 2005).

Fig. 168 - The fold-total uplift may be defined as the single fold uplift minus the regional subsidence. This value can be either positive or negative. This last case may occur along the front of W-directed subduction zones where the subsidence rates may be faster than the fold uplift rate. folds rise slower than the regional subsidence. In the first positive case the folds are deeply eroded and the onlap of the growing strata moves away from the fold crest if the sedimen- tation rate is lower than the fold uplift rate in marine or alluvial environments. In the second negative case, where folds rise at lower rates than the regional subsidence, the onlap moves toward the fold crest if the sedimentation rate is higher than the fold uplift rate. This is also the more favorable case for hydrocarbon traps where growth strata can seal the fold. Moreover the fast subsidence rates in the fore- deep provide a higher thermal maturation. This second tectonic setting is commonly asso- ciated with accretionary wedges forming along west-dipping subduction zones, which are cha- racterized by high subsidence rates in the fore- deep or trench, due to the fast "eastward" roll- back of the subduction hinge (e.g. Apennines, Carpathians, Banda arc).

Fig. 170 - Normal faults form in almost all geodynamic settings. The subsidence they generate competes with regional thermal subsi- dence, lithospheric thinning or uplifting orogens along subduction zones. and back-arc basins), in contrast with negative total subsidence (e.g. orogens). Where the han- gingwall subsidence rate is faster than the sedi- mentation rate in case of both positive and negative total subsidence, the facies and thick- ness of the syntectonic stratigraphic package may vary from the hangingwall to the foot- wall. A hangingwall subsidence rate slower than sedimentation rate only results in a larger thickness of the strata growing in the hangin-

Fig. 169 - Comparison between the Apennines and Dinarides-Hellenides front in the southern Adriatic-Ionian seas. Note the deepest foredeep along the Apennines subduction and the higher structural elevation of the Hellenides front. M, Messinian. The Apenninic front has negative fold-total uplift, whereas in the Hellenides it is positive.

Fig. 172 - Cross-section at the end of the Ladinian along the Padon Massif (Dolomites). Synsedimentary N-S trending normal faults for- med with subsidence rates slower than sedimentation rate and no facies changes occur across the faults. Fig. 171 - Relationship between tectonics and stratigraphy across an extensional syn-sedimentary fault. Subsidence rates and sediment supply are independent factors that have different values in different geodynamic settings. When the subsidence rate along a fault is lar- ger than the sedimentation rate, there is a morphological step at the fault and this may generate lateral facies changes (upper panel). When the subsidence rate is lower than the sedimentation rate, the fault only controls the thickness of the syntectonic sediments and the morphological gradient is absent in any kind of sedimentary environment, e.g., both subaerial and shallow or deep marine facies (lower panel). In the former case, a higher footwall uplift should be expected due to the mass deficit. Interbedded debris flows may be sourced both by the fault scarp and by the limb of the rollover anticline. Debris-flow deposits (megabreccias) of this origin intercalated in a sedi- mentary successions should not be confused with low-stand wedge deposits generated by global eustatic falls.

Fig. 174 - Example of the Apennines history. Pizzo Intermesoli, Gran Sasso Range, central Apennines. Jurassic transtensional tectonics sealed by Cretaceous basinal facies are preserved in a thrust sheet presently cross-cut by active normal faults to the left. Active thru- sting is presently displaced eastward in the Adriatic coast. 40 km from the leading edge of the accretiona-

Fig. 176 - The dip of the regional foreland monocline is a significant signature of thrust belts. In spite of the higher elevation, the Alps have in average lower dip of the monocline with respect to the Apennines. This asymmetry might be ascribed to the different polarity of the subduction (after MARIOTTI & DOGLIONI, 2000). Fig. 175 - Mean vertical movements along the main geodynamic settings, i.e., foredeeps, passive continental margins, and rift: Foredeeps subside for the flexure of the lithosphere generated by the loading of the upper plate along E- to NE-directed subductio zones, whereas they are possibly generated by the flexure generated by the opposite mantle flow along W-directed subduction zone: Oceanic embayments are subsiding mostly for the cooling and thickening of the lithosphere, where z is the depth of the ocean floor rela tive to the ridge, and K is a constant (320). Backarc rifts, passive margins and intraplate rifts are generated primarily by thinning of th lithosphere, plus thermal and loading effects. The red dashed line indicates the inferred trend of the isotherm marking the lithospher base.

Fig. 177 - Two examples of the regional foreland monocline dip in red (6) from the Alps, above, and the Apennines, below. The envelope of the fold crests in blue () rises toward the interior of the belt in the Alps, whereas it may dip toward the interior of the prism in the Apennines. This last geometry occurs if the regional subsidence is faster than the fold uplift. Seismic sections after BACHMANN & Kocu (1983) and PIERI (1983). accretionary wedges show high relief and broad outcrops of metamorphic rocks. They are associated with shallow foredeeps with low subsidence rates. In contrast west-dipping sub- duction-related accretionary wedges show low relief and involve mainly sedimentary cover. The related foredeeps are deep and have high subsidence rates (figs. 180, 181). This differen- tiation is useful both for oceanic and continen- tal subductions, e.g., eastern vs. western Pacific subductions, or east-dipping Alpine vs. west-dipping Apenninic subductions (figs. 182, 183, 185). In a cross-section of the Alps the ratio of the area of the orogen to the area of the foredeeps is at least 2:1, whereas this ratio is 0.22:1 for the Apennines. These ratios explain why foredeeps related to east- or nor- theast-dipping subduction are quickly filled and bypassed by clastic supply, whereas fore- deeps related to west-dipping subduction maintain longer deep water environment. These differences support the presence of an "eastward" asthenospheric counterflow relati- ve to the "westward" drift of the lithosphere detected in the hot-spot reference frame, even in the Mediterranean where no hot spots are present (fig. 184). In this interpretation, the Apennines foredeep was caused by the "eastward" push of the mantle acting on the subducted slab, whereas the foredeeps in the Alps were caused by the load of the thrust she- ets and the downward component of move- ment of the upper Adriatic plate; these forces contrast with the upward component of the "eastward" mantle flow.

Fig. 181 - Strong asymmetry of the geologic parameters is observed between foredeeps related to west-directed subduction (left), and foredeeps related to east-northeast-directed subduction (right). In the west-directed case, the area of foredeep (filled or unfilled) is wider than the area of mountainous ridge, whereas foredeeps areas are smaller in opposite directed subductions (after DOGLIONI, 1994). Fig. 180 - Alps and Apennines are thrust belts with very different characters. The Alps have a shallow foredeep, broad outcrops of meta- morphic rocks, and high structural and morphologic relief. The Apennines have deep foredeep, few outcrops of basement rocks (partly inherited from earlier Alpine phase), low structural and morphologic elevation, and the Tyrrhenian back-arc basin. These differences mimic asymmetries of Pacific east-directed Chilean and west-directed Marianas subductions. For E-NE-directed subduction foredeep and trench, the origin may be interpreted as controlled by the lithostatic load and the downward component of upper plate push, minus the upward component of the "eastward" mantle flow (upper left panel). For the W-directed subduction where subsidence rate is much higher, the foredeep origin may rather be interpreted as due to the horizontal mantle push and the lithostatic load, panel in the lower right (after DOGLIONI, 1994).

Fig. 182 - Average values of the topographic envelope (c), dip of the foreland monocline ((), and critical taper (=a +) for the two classes of subduction zones, i.e., W-directed and E- or NE-directed. Note that the "western" classes show lower values of a and steeper values of 6 (after LENcI & DoGLIONI, 2007).

Fig. 183 - Assuming fixed the upper plate U, along west-directed subduction zones the subduction hinge H frequently diverges relative to U, whereas it converges along the opposite subduction zones. L, lower plate. Note that the subduction S is larger than the convergence along W-directed slabs, whereas S is smaller in the opposite case. The two end-members of hinge behavior are respectively accompanied in average by low and high topography, steep and shallow foreland monocline, fast and slower subsi- dence rates in the trench or foreland basin, single vs. double verging orogens, etc., highlighting a worldwide subduction asymme- try along the flow lines of plate motions indicated in the inset (after DOGLIONI et alii, 2007).

Fig. 184 - Tectonic undulated mainstream along which plates move with a few cm/yr "westward" delay relative to the underlying mantle.

Fig. 185 - The orogens are structured by the combination of a number of parameters. The most important are listed in the figure (after LENcI & DOGLIONI, 2007).

Fig. 186 - NASA-Shuttle view of the central western Mediterranean. the Carpathians subduction and the Pyrenees. The Mediterranean orogens show two distinct signatures similar to those occurring in the opposite sides of the Pacific Ocean. High mor- phologic and structural elevation, double ver- gence, thick crust, involvement of deep crustal rocks, and shallow foredeeps characterize E- or NE-directed subduction zones (Alps-Betics, Dinarides-Hellenides-Taurides). On the other hand, low morphologic and structural eleva- tion, single vergence, thin crust, involvement of shallow rocks, deep foredeep and a widely developed backarc basin characterize W-direc- ted subduction zones (Apennines, Carpathians). This asymmetry can be ascribed to the "W"-ward drift of the lithosphere relati-

Fig. 187 - Present geodynamic framework. There are four subduction zones with variable active rates in the Mediterranean realm: the W-directed Apennines-Maghrebides; the W-directed Carpathians; the NE-directed Dinarides-Hellenides-Taurides; the SE-directed Alps The Apennines-Maghrebides subduction-related backarc basin of the western Mediterranean stretched and scattered into the segmentec basins most of the Alps-Betics orogen (after CARMINATI & DOGLIONI, 2004). A characteristic feature of the western Mediterranean is the large variation in litho- spheric and crustal thickness. The lithosphere has been thinned to less than 60 km in the basins (50-60 km in the Valencia trough, 40 km in the eastern Alboran sea, and 20-25 km in the Tyrrhenian) while it is 65-80 km thick below the continental swells (Corsica-Sardinia and Balearic Promontory). The crust mimics these differences with a thickness of 8-15 km in the basins (Valencia trough, Alboran sea, Ligurian sea and Tyrrhenian sea) and 20-30 km underneath the swells (Balearic Promontory and Corsica-Sardinia), as inferred by seismic and gravity data. These lateral variations in Mediterranean areas, which are mainly relics

Fig. 188 - Paleogeodynamics at about 15 Ma. Note the "E"-ward vergence of both Apennines-Maghrebides trench and the backar extensional wave. The Liguro-Provengal basin, the Valencia Trough and the North Algerian basin were almost completely opened a 10 Ma. The Dinarides subduction slowed down, due to the presence of the thick Adriatic continental lithosphere to the west, where as to the south, the Hellenic subduction was very lively due to the presence in the footwall plate of the Ionian oceanic lithosphere The Carpathians migrated E-ward, generating the Pannonian backarc basin, with kinematics similar to the Apennines (afte CARMINATI & DOGLIONI, 2004). thickness and composition are related to the rifting process that affected the western Mediterranean, which is a coherent system of interrelated irregular troughs, mainly V-sha- ped, which began to develop in the Late Oligocene-Early Miocene in the westernmost parts (Alboran, Valencia, Provengal), beco- ming progressively younger eastward (Eastern Balearic, Algerian basins), up to the presently active E-W extension in the Tyrrhenian Sea. Heat flow data and thermal modelling show that the maximum heat flow values are encountered in the basins: 120 mW/m?’ in the eastern Alboran, 90-100 mW/m’ in the Valencia trough, and more than 200 mW/m’ in the Tyrrhenian Sea. All these sub-basins appear to be genetically linked to the backarc opening related to the coeval "E"-ward rol- Iback of the W-directed Apennines- Maghrebides subduction zone. Extreme stret-

Fig. 189 - Paleogeodynamics at about 30 Ma. The location of the subduction zones is controlled by the Mesozoic paleogeography. The Alps-Betics formed along the SE-ward dipping subduction of Europe and Iberia underneath the Adriatic and Mesomediterranean pla- tes. The Apennines developed along the Alps-Betics retrobelt, where oceanic or thinned pre-existing continental lithosphere was present to the east. Similarly, the Carpathians started to develop along the Dinarides retrobelt (i.e., the Balkans). The fronts of the Alps-Betics orogen are cross-cut by the Apennines-related subduction backarc extension (after CARMINATI & DOGLIONI, 2004). Sardinia to its position prior to rotation, during the early Cenozoic the Alps were pro- bably joined with the Betics in a double ver- gent single belt. The Western Alps, which are the forebelt of the Alps, were connected to Alpine Corsica; the Alps continued SW-ward into the Balearic promontory and Betics. The retrobelt of the Alps, i.e., the Southern Alps, were also continuing from northern Italy toward the SW. In a double vergent orogen, the forebelt is the frontal part, synthetic to the subduction, and verging toward the subduc- ting plate; the retrobelt is the internal part, antithetic to the subduction, and verging toward the interior of the overriding plate. The W-directed Apennines-Maghrebides sub- duction started along the Alps-Betics retrobelt, where oceanic and thinned continental litho-

Fig. 190 - Paleogeodynamics at about 45 Ma. The Alps were a continuous belt into the Betics down to the Gibraltar consuming an ocean located relatively to the west (after CARMINATI & DOGLIONI, 2004).

Fig. 191 - Main tectonic features of the Mediterranean realm, which has been shaped during the last 45 Ma by a number of subduction zones and related belts: the double vergent Alps-Betics; the single "E"-vergent Apennines-Maghrebides and the related western Mediterranean backarc basin; the double vergent Dinarides-Hellenides-Taurides and related Aegean extension; the single "E"-vergent Carpathians and the related Pannonian backarc basin; the double vergent Pyrenees (after CARMINATI & DOGLIONI, 2004).

eastward migrating asthenospheric wedge at the subduction hinge of the retreating Adriatic plate. The flip of the subduction, from the Alpine E-directed to the Apennines W-direc- ted subduction could have been recorded by the drastic increase of the subsidence rates in the Apennines foredeep during the Late Oligocene - Early Miocene. In fact, west-direc- ted subduction zones such as the Apennines show foredeep subsidence rates up to 10 times faster (> 1 mm/yr) than to the Alpine forede- eps. The flip of the subduction could also be highlighted by the larger involvement of the crust during the earlier Alpine stages with Fig. 194 - Cross-section of the Northern Apennines. The Alps are interpreted as a boudinaged relict stretched in the Tyrrhenian Sea and western Northern Apennines. The retrobelt of the Alps is considered the prolongation of the Southern Alps. The numbers are interpre- ted vertical and horizontal velocity rates. It shows how in the same tectonic system different geodynamic mechanisms may concur. Green and red arrows indicate subsidence and uplift mechanisms respectively. Modified after CARMINATI et alii (2004b).

Fig. 195 - Phyllites recording Hercynian deformation in the NW-ward dipping basement of the Dolomites. 3 km south of Agordo Chris Golding for scale.

Fig. 198 - Transition (platform break plane, PB) between internal lagoon sediments (back reef, L) and clinostratified escarpment sedi- ments (fore-reef, F) of the Ladinian Marmolada Limestone (or Sciliar Dolomite). The plane is steeply dipping, thus indicating strong aggradation. Photo taken in the Catinaccio Massif. Fig. 197 - Detail of the Permian-Triassic boundary (the largest extinction) in the Dolomites. From the bottom to the top: marls, gyp- sum and bituminous limestones of the upper Permian Bellerophon Fm (B); the first cliff is constituted by Triassic marly limestones of the Mazzin Member (M) of the Werfen Fm (Scythian); a yellow-orange level marks the occurrence of the evaporitic dolomites of the Andraz Horizon of the Werfen Fm (A); the topmost cliff (Siusi Member of the Werfen Fm, S) is made by marly and oolitic lime- stones with frequent intercalations (toward the top) of reddish siltstones. The picture was taken at the western end of the Pale di San Martino massif, near Passo Rolle.

Fig. 199 - Latemar Massif, western Dolomites. The carbonate platform is showing its backreef or lagoon. V, late Ladinian volcanic dikes. There are different datings and genetic interpretations on the evolution of this platform (GAETANI et alii, 1981; GOLDHAMMER e: alii, 1990; KENT et alii, 2004). Whatever the interpretation, the subsidence rate can be higher than 0.6 mm/yr, a rate typical of backarc basins associated to W-directed subduction zones.

Fig. 201 - View of the southern outcropping termination of the Catinaccio prograding carbonate platform. C, Marmolada Limestone; L, Livinallongo Fm; Co, Contrin Fm; W, Werfen Fm. See next figure.

Fig. 203 - Above: seismic reflection profile in the Adriatic Sea (Crop Atlas, Scrocca et alii, 2003) crossing the Mesozoic slope of a carbonate platform. Below: outcropping example of prograding margin of a Ladinian carbonate platform (C, Marmoladz Limestone), over the basinal equivalent (L, Livinallongo Fm); W, Werfen Fm, Scythian; Co, Contrin Fm, Upper Anisian. Sass dé Putia, northern Dolomites.

Fig. 206 - Reconstruction of the Ladinian Pale di San Lucano platform; NNW is to the right. Example of divergent platform break (PB) and downlap (DP) planes. Note that the basinal Livinallongo Formation shows a deepening upward sequence (1, black, laminated sili- ceous mudstone, starved basin; 2, nodular limestone; 3, graded calcarenites organized in small thickening-coarsening upward sequen- ces). The Sciliar Dolomite (or Marmolada Limestone) is represented by megabreccias of the slope, whereas the Latemar Limestone con- sists of peritidal back-reef sediments (after DOGLIONI & BOSELLINI, 1989). The steep platform break is controlled by fast subsidence rate. It cannot be generated only by sea-level rise since the platform sequence is thicker than 1 km. Fig. 205 - The platform break and downlap planes are the ideal planes intersecting respectively all the platform break and the downlap hinges in a progradational system (after DOGLIONI & BOSELLINI, 1989).

Fig. 207 - Reconstruction of the western margin of the Sella Carnian carbonate platform (Upper Cassian Dolomite). It is an example of convergence between the platform break plane (PB) and the downlap plane (DL). The coeval basin (San Cassiano Formation) conse- quently shows a shallowing upward sequence (1, graded calcarenites and breccias alternating with shale and mudstone; 2, nodular and wavy bedded limestones with traction-current structures; 3, wave bedded calcarenites). The flat and gentle dip of the platform break has to be related to absent or low subsidence rate.

Fig. 210 - Section to the west of Raibl, Tarvisio area, between Cinque Punte and Mt. Guarda. DM, Dolomia Metallifera, Ladinian carbonate platform; B, Buchenstein (Livinallongo) Fm, Ladinian, basinal equivalent of the Dolomia Metallifera. P, Predil Limestone, lower Carnian; R, Rio del Lago Fm, lower Carnian, consisting of marls and marly limestones onlapping the slope of the underlying platform (Dolomia Metallifera); C, Conzen Dolomite, middle Carnian; T, Tor Fm, upper Carnian, basinal equivalent of the upper Carnian-Norian Dolomia Principale (DP), consisting of classical peritidal dolomite to the right and of clinostratified megabreccias and calcarenites prograding towards the north onto the Tor Fm; DL, Downlap plane; PB, Platform break plane. The picture of the previous figure represents the central-right part of the section.

Fig. 209 - Natural section west of Raibl, in the Tarvisio area, between Portella and Mt. Guarda, to the right (eastern Southern Alps). DM, Ladinian Dolomia Metallifera, equivalent to the Sciliar Dolomite; R, Rio del Lago Fm, lower Carnian, consisting of marls and marly limestones onlapping the slope of the underlying platform (Dolomia Metallifera); C, Conzen Dolomite, middle Carnian; T, Tor Fm, upper Carnian, basinal equivalent of the upper Carnian-Norian Dolomia Principale (DP), consisting of classical peritidal dolomite to the right and of clinostratified megabreccias and calcarenites prograding towards the north in the central part of the pic- ture. DL, Downlap plane; PB, Platform break plane.

Fig. 211 - Clinostratifications of the prograding megabreccias of the Ladinian Sciliar Dolomite, observable along the northern flank of the Cime di Focobon. Some latite-basaltic dykes (V) referred to the upper Ladinian volcanic event cut the platform sediments.

Fig. 212 - The famous onlap of the volcanic and volcaniclastic deposits (V) of Ladinian age onto the northern slope of the lower-mid- dle Ladinian carbonate platform (D) of the Pale di San Lucano prograding towards the north, at the head of the Gares Valley.

Fig. 213 - Pillow lava of the Ladinian shoshonitic volcanic event in the Dolomites. Locality Colcuc, André Droxler for scale.

Fig. 214 - Ladinian latit-basalt pillows covered (left) or lying on (right) and squeezing volcaniclastic turbidites. Locality Colcuc.

Fig. 215 - Caotico Eterogeneo, Ladinian mass flow deposit containing pebbles from the Upper Permian to the Ladinian. Volcanic rocks, basinal and platform fragments are included. Some elements are sheared or folded, and contained in the undeformed conglomerate, indi- cating Middle Triassic tectonics. Locality Lasta, Valle di Livinallongo.

Fig. 216 - The Caotico Eterogeneo (Ce) occurs in channelized bodies interbedded in the turbiditic volcanoclastic sandstones of Ladinian age (V). This is an example in the Livinallongo Valley, in the southern slope of the Col di Lana, where the top is folded and thrusting toward the observer.

Fig. 218 - Prograding Cassian Dolomite (D) of the Gardenaccia Massif over the San Cassiano Fm (S), toward the Val Badia. R, Raibl Fm; DP, Dolomia Principale. Fig. 217 - Carbonatic olistholith (Marmolada Limestone) included in the Upper Ladinian volcanic Marmolada Conglomerate. Road from Arabba to Passo Pordoi. Matteo Minervini, Elena Capoferri, David Hearty, Grant Ellis and others for scale.

Fig. 219 - Early Carnian San Cassiano Fm at Passo Sella. Carbonatic and volcanoclastic turbidites interfingered with emipelagic shale and marls. This basinal formation is eteropic to the Cassian Dolomite in the background. Silvia Campobasso, Cecile Huet, Eleonor: Vitagliano, Daniele Fragola and others for scale.

Fig. 220 - Undolomitized olistholith of the Cassian Dolomite carbonate plaform included in a coarse grained (volcaniclastic rich) tur- biditic facies of the San Cassiano Fm, Passo Sella. Casey Moore for scale.

Fig. 221 - Interfingering between the clinoformed megabreccias of the Cassian Dolomite (D), representing the slope, and the basin of the San Cassiano Fm (S) in the green. R, Raibl Fm; DP, Dolomia Principale.

Fig. 222 - Stratigraphy of the western margin of the Sella Massif, central Dolomites. DP, Norian tidal flat Dolomia Principale; R, Upp Carnian Raibl Fm; Du, Carnian shallow water lagoon facies Diirrenstein Dolomite; D2, prograding megabreccias of the Carnian Upp Cassian Dolomite. S2, basinal sediments of the Carnian Upper San Cassiano Fm. In the background to the north toward Passo Garden D1, megabreccias of the Carnian Lower Cassian Dolomite, S1, the basinal sediments of the Lower San Cassiano Fm. In the Gardenacc massif, the Dolomia Principale overrides the Cretaceous Puez Marls (K). Picture by Ulrich Prinz taken from a paraglide.

Fig. 223 - Carnian-Norian transition in the Mt. Faloria, southeast of Cortina. DP, Norian tidal flat Dolomia Principale; R, Upper Carnian Raibl Fm, here reddish, tidal flat and including some evaporites; Du, Carnian shallow water lagoon facies Diirrenstein Dolomite; D, megabreccias of the Carnian Cassian Dolomite.

Fig. 224 - Intraformational sedimentary dyke in the Dolomia Principale (DP) near Sass Pordoi in the Sella Massif.

Fig. 225 - Structural Model of Italy (after Bic! et alii, 1989).

‘ig. 227 - GoogleEarth map of the central-eastern Alps which formed by the subduction of the European plate beneath the Adriatic plate. [he Dolomites are an element of the Southern Alps which constitutes the retrobelt of the Alpine orogen. The Venetian Alps have been shortened main- tormation occurred at the same time as the a > a, 1987). The Dolomites to the north, in post-Varisca: times, underwent a number of tectonic event: which may be summarized as follow: Permian and Triassic rifting phases broke th area into N-S trending basins with differen degrees of subsidence. A Middle Triassic trar stensive-transpressive event then deformed th region along a N70°-90° trend, generating {lc wer structures within the basement . Volcanc tectonic domal uplift and subsequent calder formation occurred at the same time as th

Fig. 228 - Moho depth in the central eastern Alps (after KUMMEROW et alii, 2004). Late Ladinian magmatism. Early Jurassic rif- ting also controlled the subsidence which was larger in the eastern side of the Dolomites. This long period of generalized subsidence was followed by a pre-Neogene (Late Cretaceous- Paleogene?) ENE-WSW compression that generated a WSW-verging belt. This shorte- ning amounts to at least 10 km. During the Neogene, the Dolomites as far north as the Insubric Lineament, were the innermost part of a S-verging thrust belt. The basement of the Dolomites was thrust southwards along the Valsugana Line onto the sedimentary cover of the Venetian Prealps for at least 10 km. This caused a regional uplift of 3-5 km. The Valsugana Line and its backthrusts on the nor- thern side of the central Dolomites generated a 60 km wide pop-up in the form of a synclino- rium within which the sedimentary cover adapted itself mainly by flexural slip often for- ming triangle zones. The shortening linked to this folding is about 5 km with Neogene thrusts faulting and folding the shallower WSW-vereging pre-existing thrust planes. On

Fig. 229 - Evidence for the subduction of the European plate beneath the Adriatic plate from the TRANSALP profile in the central eastern Alps. EM, European Moho; AM, Adriatic Moho (after KUMMEROW ef alii., 2004). the north-eastern side of the Dolomites, Neogene deformation is apparently more strictly controlled by the transpressive effects of the Insubric Lineament (Pusteria Line) and shortening of the sedimentary cover may be greater than in the central Dolomites. Minor deformation linked to the Giudicarie belt is present in the western Dolomites. The eastern Dolomites are a recess of the earlier Alpine deformation where the ESE-verging Giudicarie system represents the left-lateral transpressive oblique ramp, and the WSW-ver- ging Paleogene thrusts of the central eastern Dolomites are the right-lateral transpressive oblique ramp. The recess is located on the Permo-Mesozoic Trento Horst. The present structure of the Dolomites is thus

Fig. 230 - Seismicity distribution in the central-eastern Alps along the TRANSALP (after Kummerow et al., 2004). Note how seismici- ty is concentrated in areas of lower topography as suggested by CARMINATI et alii (2004). In a N-S section, the central Dolomites form a wide synclinorium (figs. 232, 233), genetically related to the Valsugana thrust in the south, and its N-verging backthrusts to the north (Funes and Passo delle Erbe Lines). Therefore, the Dolomites are within a pop-up-related syn- cline generated by the two conjugate thrusts. The Valsugana thrust to the south over the Dolomites are within a pop-up-related syn-

Fig. 231 - Interpretation of the TRANSALP profile after CASTELLARIN et alii (2006). stratigraphic horizons. The different tectonic phases of the Dolomites (figs. 239 - 268) formed mainly at low tempera- ture (diagenesis) and are essentially brittle tec- tonics. However, Ladinian magmatism, Mesozoic burial and the Alpine event produ- ced anchimetamorphic temperatures. Slow ductile behavior has been recognized for the Werfen Formation, and a temperature of about 200°, suggestive of anchimetamorphism conditions (Garzanti, 1985), has been propo- sed for Carnian sandstones in the western part of the Southern Alps.

Fig. 232 - Cross-section of the Southern Alps along the southern segment of the Transalp (white thick line in upper map), modified after DOGLIONI, 1987. Map after Transalp Working Group. UP-Mz, Upper Permian-Mesozoic; T, Tertiary.

Fig. 233 - View of the southwestern Dolomites from the Sella Massif. The whole view is the southern limb of the Dolomites synclino- rium, being the Lagorai monocline related to the ramp of the Valsugana thrust. ted examples, if an oceanic or thinned conti- nental lithosphere was occurring in the fore- land of the Hercynian retrobelt during Late Permian or Early Triassic, a W-directed sub- duction zone could have taken place, retrea- ting toward Slovenia, Hungary, being the relict of this hypothetic system now below the Pannonian basin. Therefore the Middle Triassic subsidence in the Dolomites is more compatible with a backarc setting of a W- directed subduction, and with the shoshonitic signature of the magmatic event that has been interpreted as subduction-related (PISA ef alii, 1979). Past plate motions in the Atlantic-

Fig. 235 - Simplified tectonic map of the Venetian Alps and location of the next section. 1) Hercynian crystalline basement and Permian ignimbrites; 2) Late Permian and Mesozoic sedimentary cover; 3) Tertiary sediments, Flysch and Molasse; 4) Triassic and Tertiary volcanics; 5) Quaternary. Red lines, thrusts. Fig. 234 - The Southern Alps underwent compressions at different times and in different areas. K: area of Late Cretaceous-Paleogene compression (green); P: Paleogene-Early Neogene compression (Dinarides); N: Paleogene-Neogene compression (Southern Alps). The rectangle indicates the area of the next figure which has been deformed only by the SSE-vergent Neogene Southalpine thrust belt and was located in the foreland of the Paleogene WSW-vergent Dinaric thrust belt. Note in the eastern side the overlap betwe- en Southern Alps and Dinarides.

Fig. 236 - Balanced cross-section across the Venetian Alps. See Fig. 235 for location. Horizontal scale = vertical scale. Legend: Bas, crystalline basement; HGr, Late Hercynian granite; Tr, Late Permian-Lower and Middle Triassic formations; DP, Late Triassic (Dolomia Principale); G, Liassic platform facies (Calcari Grigi) gradually passing southward to Liassic-Dogger basinal facies in the Venetian Plain (Soverzene Formation, Igne Formation, Vajont Limestone); R, Dogger-Malm basinal facies (Lower and Upper Rosso Ammonitico, Fonzaso Formation; thin black layer) B, Early Cretaceous (Biancone); S, Late Cretaceous (Scaglia Rossa); E, Paleogene (Possagno Marls, etc.); N, Late Oligocene-Neogene Molasse; Q, Quaternary. to Mesozoic faults, thickness and facies varia- tions. The thrusts are arranged in an imbricate fan geometry. A frontal triangle zone laterally ends at transfer faults. Earlier stages of the thrust belt were characterized by frontal trian- gle zones, which have later been involved and cut by the progression of the internal thrusts. The Southern Alps were part of a Mesozoic continental margin, according to stratigraphic analysis, i.e. facies and thicknesses changes (BERNOULLI ef alii 1979; WINTERER & BOSELLINI 1981). The area can be divided into three main structural sectors during Mesozoic. These are, from west to east: the Trento Platform, the Belluno Basin and the Friuli Platform. True Mesozoic normal faults (ie. Liassic) have been documented at the western border of the Trento Platform (CASTELLARIN 1972; DOGLIONI & BOSELLINI 1987) and pro- posed also at the eastern margin (WINTERER & BOSELLINI 1981; BOSELLINI e¢ az 1981; BOSELLINI & DOGLIONI 1986; MASETTI & BIANCHIN 1987, DOGLIONI & NERI 1988). Tt . WA an ee ee eewewnl Lawl lee, Pesrenel anna ea ler

Assunta Well at 4747 m where Late Triassic Fig. 237 - Cross-section of the Dolomites and eastern Venetian Prealps (after SCHONBORN, 1999)

Fig. 239 - Simplified tectonic map of the Dolomites. Fig. 238 - Schematic picture of the Venetian Alps front. The foothills are characterized by a triangle zone. The deep thrust plane (Bassano Line) reaches the surface (Maniago Line) through a transfer zone (Caneva Line). The Venetian Plain represents the foredeep of two opposite thrust belts. The geo- metry of the foreland basin changes from the triangle zone to the normal thrust to the east. In the first case the clastic sedi- ments are pushed southwards, tilted and eroded. In the second case they are simply thrusted.

Fig. 240 - E-W cross-section of the Dolomites, showing WSW-vergence of thrusting. Location in figure 239.

Fig. 241 - NNW-SSE cross-section of the Dolomites, showing the wide basement syncline generated by the pop-up. Location in fig- 239.

Fig. 242 - Croda del Vallon Bianco, northwest of Cortina. The thrust is WSW-verging and mainly Jurassic Calcari Grigi outcrop both in the hangingwall and in the footwall. The geometry of the hangingwall cut-off allows to reconstruct the undulated geometry of the fault plane. are thrust Dy the Valsugana Line (VENZO 1939) and Pliocene shales are folded along the frontal triangle zone. Moreover Messinian- Pliocene onlap geometries in the southern bor- der of the chain support a mainly Neogene age of the deformation. The extension of the unconformities within the molasse is a fun- ction of the thrust belt structure and reflects areas of stronger uplift. A problem is represen- ted by the style and timing of the orogenic evolution: has the chain southward regularly risen in a continuum creep since Late Oligocene time and do the unconformities record moments of sea level fall (low stand)? Or was the chain generated step by step so that the unconformities simply mark moments of tectonic crisis? In general, plates move with a regular velocity suggesting that in areas of deformation the tectonic evolution should fol- low an almost constant activity. If the tectonic evolution generated with a constant regularity and tectonic crisis had a wavelength too short or too long with respect to the eustatic sea level changes, then an interesting problem appears in dating the thrust belt: the timing of SS ROS oe Ree gay meng pe Satasity mae sete es To the north, the Belluno Line may have been a blind thrust generating a triangle zone during earlier stages of the deformation, later rising at the surface in the northern limb of the Belluno Syncline. This is supported by the steep attitude of the northern limb of the Belluno Syncline which is difficult to explain geometrically as a simple footwall syncline. We also note that triangle zones mainly occur in the Mesozoic Belluno Basin, rather than in the neighboring platforms. In fact the Belluno Syncline is developed in the deepest structural zone with thick basinal lithofacies. This earlier structural situation had an important influen- ce in the morphology and source areas of the hydrographic pattern during the Late Miocene. In the Venetian segment of the Alps the GE es ee er ee ee ey oe) fe

Fig. 243 - Upper Oligocene-Lower Miocene conglomerate (MP) at Mt. Parei - Col Bechei, to the northwest of Cortina, sealing an angu- lar unconformity with the underlying Liassic Calcari Grigi (CG) deformed by WSW-verging thrust and folds. This outcrop dates active pre-Late Oligocene WSW-verging shortening in the Dolomites. the tectonic crisis has been considered as the time missing at regional unconformities and the age of coarse-grained sediment supply (Le. the Messinian Conglomerate) onlapping the discontinuity. But if the unconformities recor- ded only moments of general lowstand (in this case global or confined to the salinity crisis in the Mediterranean area) then we could argue that the chain developed more gradually and that the sea level oscillations orchestrated the arrangement of the syntectonic sedimentary sequences. The different interpretations of the unconformities and conglomeratic supply in the molassic sequences allow different tectonic reconstructions. With the sea level change interpretation (VAIL ¢f alii 1977) the chain rises constantly during Late Oligocene-Neogene times, but if we assume the unconformities and conglomeratic supply to be tectonic-rela- tectonics already during Paleogene times.

Fig. 245 - The Col Bechei structure, northwestward of Cortina. The WSW-verging thrust (1) is sealed by the Upper Oligocene Mt. Parei conglomerate. The thrust and the related fault-propagation recumbent fold are cross-cut by the later S-verging alpine thrusts (2, and 3). Numbers refer to the kinematic sequence. The two almost orthogonal compressive phases generated an interference structure. P, Dolomia Principale (Norian); D, Dachstein Limestone (Rhaetian); G, Calcari Grigi (Liassic); R, Ammonitico Rosso (Dogger-Malm); M, Mt. Parei Conglomerate (Oligocene-Miocene). After DOGLIONI & SIORPAES (1990). Fig. 244 - Col Bechei western side. The overturned fold in the hangingwall of a WSW-verging thrust, has been later cross-cut by a S- verging thrust. See following figure. CG, Liassic Calcari Grigi.

swells. The present tectonic configuration is due to the inherited Mesozoic background. The structural evolution of the area followed boundary features as transfer zones at horst margins (i.e. at the Trento Platform and Friuli Platform margins) which have influenced the

Fig. 246 - Southern view of the Puez-Gardenaccia Massif, showing N-S trending listric normal faults, above Colfosco, in Val Badia. The faults dip to the east and cut the succession up to the Dolomia Principale. They are somewhere accompanied by reddish oxidations and by sedimentary dykes. The grass ledge corresponds to peritidal deposits of the Diirrenstein or Raibl Fm. Legend, DP, Norian Dolomia Principale; R, Upper Carnian Raibl Fm; C, Lower Carnian Cassian Dolomite; S, San Cassiano Fm. Compare with the section of fig. 251 where the above shown tensional structures are represented below Gardenaccia.

Fig. 247 - Example of a W-verging anticline inheriting a normal fault of Mesozoic(?) age at the top of the Col di Lana, central Dolomites. V, Upper Ladinian volcaniclastic sandstone; S, Lower Carnian San Cassiano Fm. thrust belt. Earlier Mesozoic features strongly influenced the evolution of the chain. Any kind of structural undulation along strike of the thrust belt is associated with pre-existing synse- dimentary faults, thickness and/or facies varia- tions in the sedimentary cover. The thrusts are arranged in an imbricate fan geometry and show a frontal triangle zone which was proba- bly present at earlier stages of the thrust belt in more internal zones. The variations along stri- ke of the deformation are reflected also in the Neogene and Quaternary molasse. The frontal triangle zone appears to be a growth fold rising from Late Oligocene to Quaternary times because clastic sedimentation on the southern limb of the anticline shows onlap geometries and reduced thicknesses. According to this pro- gressive evolution of the thrust belt, the uncon- formities within the molasse could record low- stands of eustatic cycles. The effects of four independent Neogene to The effects of four independent Neogene to make the Venetian Alps a classic example ol

Fig. 248 - Northeastern slope of the Sass de Putia, northern Dolomites. A N-S trending normal fault lowers the eastern side of the mas- sif. C, Marmolada Limestone; L, Livinallongo Fm; Co, Contrin Fm; W, Werfen Fm; G, Gardena Sandstone.

Fig. 249 - N-S trending normal fault likely of Mesozoic age outcropping near Cencenighe. Co, Contrin Fm; W, Werfen Fm. aU IS UMportant tO Gistieuish their Meany and kinematic overlap of the related thrust sheets. WSW-verging thrusts enter the Dolomites section of the Transalp profile. 3) The Carpathians subduction probably follo- wed the same geodynamic scenario of the Apennines, which started to develop along the retrobelt of the Alps where in the foreland to the east there was oceanic or thinned continen- tal lithosphere, like the W-directed subduction zones of the Barbados and the Sandwich arcs in the Atlantic ocean (DOGLIONI ¢éf alii, 1999). According to this interpretation, the Carpathians developed along the retrobelt of the northern Dinarides, such as the Balkans in the south, where oceanic or thinned continen- tal lithosphere of the Dacide basin was located to the east of the retrobelt. The consumption of this basin accompanied the E-ward retreat of the subduction to the present position in Vrancea. This implies that in the Pannonian basin there are elements of both Alps and Dinarides stretched and scattered.

Fig. 250 - Evidences of Middle Triassic rifting in the Adriatic Sea, comparable to what outcropping in the Southern Alps. Front page, compressed random 3D seismic profile crossing the North Adriatic Mid Triassic depocenters and ridges. Above, two-way time isopach of Northadriatic Anisian-Carnian section. The mapped interval extends from the Midtriassic Unconformity to the base of Dolomia Principale, thus including the Late Carnian evaporite seal. After FRANCIOSI & VIGNOLO (2002), EAGE. of the Alps and external Dinarides. The long- term subsidence of Venice and other coastal cities appears determined by the retreat of the subduction hinge of the Adriatic plate dipping underneath the Apennines. The flexure of the subduction hinge affects areas more than 250 km far to the northeast of the Apennines front, providing data on the viscoelastic beha- vior of the Adriatic plate continental litho- sphere. Slab rollback of such a continental lithosphere can only be ascribed to the 4) Since the Paleogene, the Venetian and Friuli plains underwent subsidence due to the load of the Alpine and Dinarides thrust sheets, for- ming two related and anastomosed foredeeps. However, the thinning of the Pliocene- Quaternary sediments (MERLINI ¢7 a/ii, 2002), and the decrease of the regional monocline dip moving away from the Apennines front (MaRIOTTI & DOGLIONI, 2000), they indicate an active subsidence all around this belt. The Apennines foreland subsidence affected most However, the thinning of the Pliocene-

Fig. 251 - Paleotectonic section of the central-northern Dolomites between the Trento Horst to the west, and the Belluno graben to the east. All Mesozoic formations thicken moving eastward, suggesting active normal faulting at the hinge between the structural high anc the low throughout the Triassic and Jurassic. Note how the extension and the differential subsidence is distributed in a number of nor- mal faults, each one with relatively small offset. K, Cretaceous; J, Jurassic; DP, Norian Dolomia Principale; CD, Carnian Cassiar Dolomite. Notice their regular spacing and the occurrence of the Calcari Grigi only to the east (to the west the Jurassic J is representec by the Ammonitico Rosso). The extensional fault system seems to have been active also during the Triassic, as suggested by the westward thinning of the Dolomia Principale (P) and of the underlying sedimentary section. Also the Raibl Fm changes its facies anc thins toward the west. K, Puez Marls Fm, Cretaceous. The WSW-vergent thrusts developed as ramps along inherited extensional zones. or a platform margin.

Fig. 252 - Angular unconformity between the Diirrenstein Dolomite (Du) and the overlying Raibl Fm (R) at the Rifugio Di Bona, beneath the Tofana Massif. DP, Dolomia Principale.

Fig. 253 - As in modern examples, where W-directed subduction zones developed along the retrobelt of an opposite subduction zone. the Middle Triassic of the Dolomites could have recorded the backarc rifting of such a setting. The Hercynian orogen was probably a collage of different subductions zones, but in general shows the characters of E-directed subductions zones, as the double vergen- ce. If a small oceanic embayment was located east of the Hercynian retrobelt during Late Permian or Early Triassic, then a W-direc- ted subduction zone could have developed, generating fast subsidence rate in the backarc basin (e.g., >600 m/Ma, Middle Triassic Southern Alps), and the shoshonitic magmatism. topography the monocline 1s at places even

Fig. 255 - Extensional Mesozoic architecture of the Southern Alps at the end of Lower Cretaceous time. 1) Permian - Carnian p.p.; 2) Carnian p.p. - Rhaetian; 3) Hettangian - Bajocian; 4) Bathonian -Aptian pp.; 5-6) Location where samples have been collected for orga- nic matter maturity analyses and for thermal modelling (after FANTONI & ScoTTI, 2003). The Upper Cretaceous crustal structure in only speculative and does not reflect the present day crustal thicknesses, which have been increased by the Alpine compressional tectonics. Fig. 254 - Interpreted W-E Early Cretaceous cross section from the Asiago plateau to the Cansiglio plateau, showing the coeval tensio- nal tectonics, which produced differential subsidence in the area. Note the basinward prograding carbonate platform in the eastern side (M.Cavallo Limestone) and the different thicknesses in the basinal sedimentation (Biancone). Neptunian dikes (vertical yellow seg- ments) characterize the zones of tensional tectonics (i.e. the Arsi¢é Lake, Fonzaso). The Neogene deformation has been strongly influen- ced by the pre-existing structural and stratigraphic geometries. See location and compare the general tectonics of the area with the inhe- rited Mesozoic structural background in fig. 149.

Fig. 256 - Recontruction of the Late Triassic-Jurassic rifting in the central-western Southern Alps (after BERTOTTI et alii, 1993). The Venetian Plateau represents the Trento Horst.

Fig. 258 - Main tectonic phases and related structures in the Dolomites after the Hercynian orogeny, apart the extensional Mesozoic tectonics (after DOGLIONI, 1987).

Fig. 260 - The earlier alpine deformation inherited the Trento horst determining a recess on the structural high, and a salient in the gra- ben to the east. The transfer zone consists of right-lateral transpression with thin-skinned WSW-verging thrusts. The later deeper thick- skinned deformation, involving the basement, generated the Dolomites pop-up, folding and tilting the earlier shallower thrusts. Fig. 259 - Overlap of two thrusts perpendicular to each other. A) pre-deformation stage; B) a thrust moves toward the reader with two lateral ramp cut-offs, strike view; C) later a thrust propagates with staircase trajectory, dip view. Assuming north to the left, this exam- ple could be applied to the interference pattern in the central-eastern Dolomites between older (Eocene?) WSW-verging thrusts and folds, later deformed by the deeper Miocene to Present SSE-verging thrusts and folds.

Fig. 261 - Main tectonic phases in the upper crust of the Dolomites. A, Permian-Triassic rifting; B, strike-slip tectonics (prevalent tran- stension with localized transpression) along a N70° trend; C, W-SW vergent thrusting of Paleogene age; D, Neogene SE-vergent thru- sting, pop-up formation of the Dolomites synclinorium.

Fig. 262 - The WSW verging thrust structures of pre-Late Oligocene age in the Dolomites can be interpreted as the prolongation of the Dinarides front into the Southern Alps, in which there is a vast area of overlap and interference structures.

Fig. 263 - Alternatively to the previous figure, the WSW verging thrusts in the Dolomites are here interpreted as the oblique right-late- ral transpressive ramp of the advancing Eocene alpine thrusting between the Trento recess and the Belluno salient. In this view the area of Alps-Dinarides overlap is more restricted to the east.

Fig. 264 - 3D view and interference between the Alpine and Dinaric subductions in northeastern Italy. The extension of the Pannonian, Carpathians subduction-related back-arc basin, crosscuts the Eastern Alps and the northern Dinarides. This indicates that, in the same area, independent geodynamic settings may concur and plate boundaries are passive features.

Fig. 265 - Seismic reflection profile (CROP M-18) of the northern Adriatic Sea, parallel to the coast, offshore Venice. Note the thinning of the Pleistocene sediments moving north-eastward, indicating coeval differential subsidence in the underlying rocks. The Pleistocene Pliocene boundary is marked. M, Messinian unconformity. Vertical scale in seconds, two way time.

Fig. 266 - Seismic and borehole data constrain the geometry and southwestward thickening of the Pleistocene sediments in the Po Basin. northern Adriatic Sea and Friuli plain.

Fig. 267 - Schematic profile across the Apennines, Po Basin and Venetian area, showing the curvature of the Adriatic plate in the fore- land of the Apennines, associated to the slab retreat. The 1mmi/yr subsidence rate of Venice is the distal effect of the Apennines subduc- tion hinge retreat (after CARMINATI et alii, 2003).

Fig. 268 - The subsidence related to the flexure induced by the Apennines subduction decreases from south to north. It affects all o the Po Plain basin and part of the Alps and of the Dinarides, where it contrasts the general uplift trend related to the orogenesis. Als the Dolomites are slightly affected by the Apennines subduction subsidence. The tectonic front of the Southern Alps corresponds tc the alpine foothills in the eastern side, whereas it is buried in the western Po Basin due to the closer effect of the Apennines slab retre- at. The lower panel represents an idealized cross-section between the northern Apennines and the western Southern Alps. The las ones are the south-vergent retrobelt of the right-lateral transpressive arm of the Alps, and formed between the Late Cretaceous anc the Late Miocene, plus recent reactivations. Their shallow foredeep and the thrust planes have been southward tilted by the Lat Miocene - Quaternary northward propagation of the basement monocline generated by the Apenninic subduction (after DOGLIONI 1993; CARMINATI et alii, 2003).

Fig. 269 - Location area of the field trip. Numbers indicate the days of the excursion.

Fig. 1.0 - Itinerary: Longarone, Prealpi Bellunesi, Feltre. Google Earth view. Dashed red lines are the main thrusts of the area.

Fig. 1.2 - Eastern flank of the Piave Valley, near Longarone. The fold hinge (h) observable in the Calcare del Vajont (Dogger) belongs to the northern flank of the anticline developed in the hangingwall of the Belluno Thrust. It should correspond at depth to a transi- tion from horizontal detachment to ramp of the thrust fault. The flank dipping to the north was the movement plane of the Vajont landslide, detached along the underlying Fonzaso Fm. Fig. 1.1 - Cross-section through the eastern Venetian Alps. The Belluno Thrust emplaced at the margin between the Belluno Basin sequence (in the hangingwall) and the Friuli Platform sequence (in the footwall). Legend: C, undifferentiated crystalline basement; T, Late Permian and Early-Middle Triassic formations; P, Late Triassic Dolomia Principale; J, undifferentiated Jurassic: platform facies (Calcari Grigi) in the southern part of the section, gradually passing northward to basinal facies (Soverzene Formation, Igne Formation, Vajont Limestone, Fonzaso Formation, Ammonitico Rosso); K, Cretaceous, platform facies (Calcare del Monte Cavallo, brick pat- tern) gradually passing northward to slope deposits and basinal facies (Calcare di Soccher, Biancone); S, Scaglia Rossa, Late Cretaceous - Paleocene; F, Eocene Flysch; M, Late Oligocene - Early-Middle Miocene Molasse; ME, Messinian conglomerates.

Fig. 1.4 - The eastward prolongation of a Valsugana Thrust splay outcropping in the southern cliff of the Mt. Borga. The thrust is ver- ging toward the observer and shows a lateral ramp geometry dipping to the east. In the hangingwall the overturned forelimb indicates fault-propagation fold kinematics. Va, Dogger Vajont Limestone; SR, Upper Cretaceous-Lower Tertiary Scaglia Rossa.

Fig. 1.3 - Western flank of the Piave Valley, above Longarone. Whitish Jurassic limestones (mostly Calcare del Vajont, Va) were thrust over the Scaglia Rossa (SR) along a sharp decollement plane. The thrust is a splay of the Valsugana Line. To the left, the thrust fault cuts downsection the Scaglia Rossa sediments, or it is folded by the deeper thrusts.

Fig. 1.6 - Frontal view of the Liassic growth fault in the basinal deposits of the Igne formation, in the Vajont canjon. Notice the ramifi- cation of the fault plane and the lateral and oblique ramp geometry of the fault. Extensional slickenfibers confirm that the hangingwall moved downward with respect to the footwall. Small conjugate fractures are confined to the hangingwall of the fault plane.

Fig. 1.8 - The landslide from Mt. Toc that generated the Vajont tragedy of the 9th October 1963, in which the water of the underlying artificial lake was squeezed out and devastated the city of Longarone with 1910 casualties. The decoupling surface is the basinal Middle Jurassic Fonzaso Fm, and the dip of the beds is due to the deeper S-verging ramp of the Belluno Thrust. Fig. 1.7 - Schematic three-dimensional view of the exten- sional fault of the previous figures. A major regional exten- sional structure can be envisaged in the Col delle Tosatte Line, probably an extensional Mesozoic fault dipping to the west, reactivated as a sinistral transpressive fault during the Neogene Southalpine compressional deformation.

Fig. 1.9 - The thrust outcropping at the top of Mt. Borga, has a lateral ramp cutting downsection eastward where it continues, east of the Vajont area, near Erto. Here it outcrops at lower elevation and generates a triangle zone as in this photo. The white part of the thrust is behind or below the ridge in the foreground. Ar, Ammonitico Rosso; Va, Vajont Lm.

Fig. 1.10 - Footwall syncline beneath a branch of the Valsugana thrust in Val Zemola. CV, Middle Jurassic Vajont Limestone; AR, Upper Jurassic Ammonitico Rosso.

Fig. 1.11 - A branch of the Valsugana thrust system affects the Norian Dolomia Principale (DP) at the toe of the Duranno Mt., at the top of the Val Zemola. Note the different cutoff angles of the hangingwall and the footrwall.

Fig. 1.12 - Fault-propagation fold in the hanginwall of the Belluno Thrust, west of Mas. CG, Liassic Calcari Grigi; DP, Norian Dolomia Principale. Not visible in the footwall the Tertiary Molasse and Flysch.

Fig. 1.15 - Southern view of the Southern Alps front with the active Montello anticline affecting Upper Miocene conglomera- tes and Pliocene shales.

Fig. 1.14 - The Eocene Flysch (F) covered unconformably by the Upper Oligocene Molasse (M) in the Belluno Syncline, in the footwall of the Belluno Thrust.

Fig. 2.0 - Itinerary: Feltre, Val Cismon, Passo Rolle, Passo Valles, Moena. Dashed red lines are the main thrusts.

Fig. 2.2 - The Tezze thrust cropping out above the village of Tezze, in Valsugana. The frontal overturned limb of the fold suggests an origin by fault-propagation. This mechanism, however, was coupled to fault-bend folding, due to the ramp geometry of the fault in the Calcari Grigi (G) and the flat geometry at the bottom of the Biancone (B). Ammonitico Rosso inferiore, Fonzaso Fm, Ammonitico Rosso superiore (R).

Fig. 2.3 - Cross-section of the Vette Feltrine, after D'ALBERTO et alii (1995).

Fig. 2.4 - The Belluno thrust as it crops out along the western flank of the Cismon valley. The vertical and somewhere overturned exter- nal flank of the hangingwall anticline indicates folding by fault-propagation. DP, Dolomia Principale (Norian); CG, Calcari Grigi (Lias); B, Biancone (Lower Cretaceous).

Fig. 2.6 - Valsugana thrust, north of Passo del Broccon. In the hangingwall the crystalline basement (H), here cropping out as an apo- physis of the late-Hercynian Cima d'Asta granite. The basement is thrust towards the south over the Liassic Calcari Grigi (CG).

Fig. 2.7 - Same thrust as before, but more to the east: the Valsugana thrust outcrops with the Cassian Dolomite (Carnian, D) overri- ding the Dolomia Principale (Norian, DP).

Fig. 2.8 - The Valsugana thrust in the southern cliff of the Mt. Comedon, Sass de Mura massif, Vette Feltrine. Note the triangle zone ir the footwall. DP, Dolomia Principale (Norian); C, Calcari Grigi (Lias); F, Fonzaso Fm (Dogger); Photo by Lucio D'Alberto.

Fig. 2.9 - The Valsugana thrust cuts through inherited Mesozoic structures: the Trento (or Atesina) Platform and the Carnico-Bellunes¢ Basin. This is evidenced by the thickness of the sedimentary cover that shows values around 2 km on the Trento horst and 4-5 km in the Carnico-Bellunese graben to the east (after BOSELLINI & DOGLIONI, 1986). The strike of the normal paleo-faults is always ca. N-S Unlabelled yellow, crystalline basement; g, late-Hercynian granitoid; G, Gardena Sandstone (Upper Permian); P, Bellerophon Fm (Uppet Permian); W, Werfen Fm (Scythian); Bricks, Anisian; L, Ladinian and Carnian carbonate platforms; Lb, Livinallongo Fm (Ladinian); V. volcanoclastic deposits (Late Ladinian); S, San Cassiano Fm (Carnian); Du, Diirrenstein Dolomite (Late Carnian); R, Raibl Fm (Late Carnian); DP, Dolomia Principale (Norian); J, Calcari Grigi (Liassic); D, Dachstein Limestone (Rhaetian); K, Cretaceous; T, Tertiary. One of the main Mesozoic extensional faults accommodating differential subsidence between the platform and basin realms is the Passo Rolle Line, later reactivated as a transfer (mainly strike-slip) fault between the Lagorai monocline and the Pale di San Martino syncline.

Fig. 2.10 - The Mesozoic structuration can be recognized also in geological maps that show a structural high of the basement in the han ging-wall of the Valsugana thrust between Trento and the Cismon Valley, developed along the Passo Rolle Line. To the east of this linea ment, the Valsugana Line cuts mainly through sedimentary cover, since this area was structurally depressed before the onset of com pression. A and B are the traces of the cross sections of the next figure. 1, crystalline basement; 2, late-Hercynian granitoids; 3, Permia ignimbrites; 4, Upper Permian and Lower Triassic; 5, Upper-Middle Triassic; 6, Jurassic-Cretaceous; 7, Tertiary.

Fig. 2.11 - Geological cross-sections through the Valsugana Line: notice that along section A the basement (b) is more elevated in th hangingwall with respect to section B, although no significative differences of shortening occur between the two sections. This featur is explained by the fact that along section A the Valsugana Line cuts the horst of the Trento platform (or Atesina Platform) whereas i section B it cuts the Bellunese graben. 1, Undifferentiated crystalline basin; 2, Gardena Sandstone; 3, Bellerophon Fm; 4, Werfen Fm 5, Anysian and Ladinian formations; 6, Carnian formations; 7, Dolomia Principale (Norian); 8, Jurassic limestones; 9, Cretaceous lime stones and marls. The traces of the sections are shown in fig. 2.10.

Fig. 2.12 - Cross-section across the Valsugana thrust, at the southern margin of the Dolomites, in the zone of Passo Cereda. The base- ment of the Dolomites (B) thrusts over and wedges within the sedimentary cover of the Venetian Prealps, developing a triangle struc- ture. The basement shows a fault-bend-fold. The southern flank of the fold shows the Dolomia Principale (DP) wedging between the basement and Ladinian carbonates (L), generating thus another triangle structure, with the substitution of the cover at its bottom (direct contact between basement and Dolomia Principale. P, Upper Permian; W, Scythian and Anisian; G, Jurassic limestones; C, Cretaceous limestones and marls.

Fig. 2.13 - The Pale di San Martino are located at the core of a regional undifferentiated basement (B) syncline. The sedimentary cover accommodated the folding by flexural slip and shear, thus developing south-verging thrusts in the southern limb and north-vergent thrusts in the northern limb of the syncline. UP, Gardena Sandstone and evaporitic Bellerophon Fm (Upper Permian) which acted as a main decollement level between cover and basement (see next figure). LT, Lower Triassic; MT, Middle-Upper Triassic.

Fig. 2.14 - North-vergent en-echelon folds in the evaporitic Upper Permian Bellerophon Fm (B). The occurrence of gypsum anc shales lowers the frictional resistance, thus enabling sliding. The outcrop is at the base of the Pale di San Martino massif, alon; the road to Passo Rolle from San Martino di Castrozza. Folds are confined within this formation made of shale, gypsum, marl limestone, and indicate a N-vergent decollement. Note the overlying unfolded Werfen Fm (W) and Marmolada Limestone (C). I testifies the flexural sliding (towards the north in this case, since the outcrop is positioned in the northern flank) of the sedimen tary cover within the basement syncline.

Fig. 2.18 - Positive flower structure cropping out near Passo Rolle (locality Malga Fosse). Notice the uplifted wedge between two high angle faults which branch at the bottom of the picture. The structure is oriented N5°-10°E and shows a sinistral shear. Note that moving up along the faults the undulated geometry determines either normal or reverse offset. A, Andraz horizon that, as reference level, marks the vertical throws. The outcrop shows various levels of the Werfen Fm (Scythian). At the bottom of the flower structure, the faults con- verge in the Bellerophon Fm (B) evaporites (Upper Permian). A dark Ladinian volcanic dyke (v) is subparallel to the western fault. Legend: M, Mazzin Member; A, Andraz horizon; Si, Siusi Member; Ca, Campil Member; V, volcanic dyke.

Fig. 2.21 - A Ladinian latit-basaltic dyke at Passo Rolle cross-cutting the Upper Permian Gardena Sandstone. The alpine deformation concentrated in the dyke where mesoscale NO-10° trending left-lateral strike slip faults can be observed. Fig. 2.20 - System of mesoscale thrust faults near Passo Rolle. The imbricated fan propagated from right to left. The kinematics are indicated by the dip of the fault planes: the fault to the right (1) is more inclined and reflects the transport along the ramp of the thrust in the middle (2). The distance among fault planes is constant since the lithologies (oolitic limestones, marly limestones and siltsto- nes of the Siusi Member of the Werfen Fm, Scythian, Lower Triassic) cross-cut by the faults do not change laterally. Micro- and meso-structures, slickenlines, small conjugate backthrusts, pressure solution and other features make of this small outcrop a minia- ture of a fold-and-thrust belt such as the Southern Alps.

Fig. 2.22 - Geological reconstruction of the Passo Rolle area in the Anisian. Notice that the area is located in the tensional hinge zone between the Trento Horst and the Belluno Graben. This is evidenced by the occurrence of the extensional fault system of the Passo Rolle Line, by the thinning of the Scythian and Anisian formations and, at places, by the lack of some levels of the series towards the north- west (DOGLIONI & NERI, 1988). The Anisian unconformities are sutured by late Anisian formations. The alpine compressional deforma- tion inherited this complex architecture: for example, the thrust fault of the Cimon della Pala becomes horizontal in the evaporitic level of the Cencenighe Member. The flower structure of Malga Fosse (Fig. 2.18) is probably related to sinistral transfer motions along the Passo Rolle Line and re-sheared a Mesozoic tensional structure. Moreover, the Upper Permian Bellerophon Fm is characterized by a blind thrust that generated an accommodation fold in the overlying cover, now visible between Crode Rosse and Malga Fosse.

Fig. 2.23 - The about N-S trending Cismon Valley, with the Passo Rolle at its head, is the seat of normal faults separating the Trento Horst and the Belluno Graben to the east. Since the Permo-Triassic thickness of the sedimentary cover in the Trento Horst is about half (2 km) of the cover to the east, this supports a long lasting history of this hinge, acting as one of the major hinge zones during the Tethyan rift. During the Alpine shortening, the hinge was reactivated as a left-lateral strike-slip transfer zone. In the hangingwall of the Valsugana thrust, to the west the basement dips steadily northward about 30°, parallel to the deep thrust plane. To the east, the basement and the cover are folded within the Pale di San Martino syncline.

Fig. 2.24 - South of the Pale di San Martino, the S-verging Valsugana thrust carries the fault-propagation fold of Norian Dolomia Principale (DP) over the Lower Cretaceous Scaglia Variegata, Biancone (B) and the Middle-Upper Jurassic Fonzaso Fm (F). Note the long overturned forelimb.

Fig. 2.25 - Valsugana thrust footwall syncline affecting Liassic Calcari Grigi, south of Agordo in the Mt. Cielo.

Fig. 3.0 - Itinerary: Moena, Passo S. Pellegrino, Avoscan, Arabba

Fig. 3.2 - Passo Feudo, along the Stava Line, west of Predazzo. A latitic-basaltic dyke (V) of upper Ladinian age cuts chevron faults developed in the dark limestones of the Upper Anisian Moena Fm (Mo), dating to the Middle Triassic the age of deformation.

Fig. 3.3 - Morphological map of the basement (top of the Permian porphyrites) in the central-western Dolomites. The contour lines (every 100 m) are referred to the mean sea level. Notice that the northern flank of the Cima Bocche anticline is truncated by a fault system that seems to continue, to the west, in the Trodena and Stava Lines (Monte Rocca structure, Fig. 3.4). The motion along the alignement is sinistral transpressive. The Upper Ladinian-Lower Carnian(?) plutonic bodies and the Predazzo caldera cut, and therefore date this N70°E alignement (Stava Line-northern flank of the Cima Bocche anticline) that deformed sedimen- tary rocks up to the upper Ladinian. Notice that volcanism also developed along the alignement. In the left part of the figure, the opposite vergence of the Stava and Trodena Lines can be observed. Such opposite vergence has been interpreted as the result of a flower structure. The thick dashed line is the trace of Fig. 3.4).

Fig. 3.5 - Cross section through the Stava Line, near Passo Feudo, west of Predazzo. The section is evidently not-retrodeformable in two dimensions. This suggests strike-slip kinematics for the Stava Line. Notice the subvertical geometry of the Livinallongo Fm (Liv) tectonically in contact with the Monte Agnello subhorizontal carbonate platform sediments (LP). LP, Sciliar Dolomite and Latemar Limestone (Ladinian); Liv, Livinallongo Fm (coeval Ladinian basinal facies); D, Ladinian volcanic dykes; VC, Upper Ladinian lavas and volcaniclastic deposits; A, Upper Anisian: Moena Fm north of the Stava Line, Contrin Fm and Serla Dolomite south of the line; W, Scythian Werfen Fm; B, Upper Permian Bellerophon Fm; V, Upper Permian Gardena Sandstone; P, Permian Ignimbrites. Numerous volcanic dykes are not represented.

Fig. 3.6 - Scheme illustrating that both extensional and compressional deformation affecting a platform-basin system can be retro- deformed, contrary to strike-slip motions, as seen for the Stava Line, near Passo Feudo (Fig. 3.5). Notice that in the strike-slip pane! the toe wedge of the carbonate platform is missing.

Fig. 3.7 - Reconstruction of the Monte Agnello and Latemar carbonate platforms before the strike-slip motion occurred along the Stava Line and of the volcanic event that cut the Stava tectonic alignement. North is to the left.

Fig. 3.9 - Detail beneath the Col Becher. The piercing "diapiric" structure has been decapitated by the upper thrust, which has later beer folded. In the center is also visible the Anisian transtensional fault sealed by the Richthofen Conglomerate. Legend, Ladinian: C Marmolada Limestone; L, Livinallongo Fm; Upper Anisian: Co, Contrin Fm; R, Richthofen Conglomerate; Scythian: Werfen Fm: Cn Cencenighe Mb; Vb, Val Badia Mb; Ca, Campil Mb; Og, Gastropod Oolite; Si, Siusi Mb; A, Andraz horizon; M, Mazzin Mb. Late Permian: B, Bellerophon Fm).

Fig. 3.11 - Detail of the Col Becher structure: the Anisian fault lowered the Cencenighe Mb of the Werfen Fm, now in contact with the Campil Mb (dark red) in the footwall. The Richtofen Conglomerate and the Upper Anisian Contrin Fm unconformably overlie the ero- sional surface that leveled the extensional structure. Notice, in the Cencenighe Mb of the hanging wall a drag fold associated to the nor- mal or transtensional fault. Above, the Col Becher thrust fault with reddish siltstones of the Werfen Fm overriding the Anisian dolomi- tes. C, Campill Mb; V, Val Badia Mb; E, Cencenighe Mb; R, Richthofen Conglomerate; CO, Contrin Fm. See also the map and the sec- tions in the following figures.

Fig. 3.16 - Geological map and cross section of the double vergent diapiric anticline of Val San Nicolo. B, Bellerophon Fm; M Mazzin Mb; A, Andraz horizon; S, Siusi Mb; O, Gastropod Oolite; C, Campil Mb; V, Val Badia Mb; E, Cencenighe Mb; R. Richthofen Conglomerate; CO, Moena Fm, Contrin Fm; F, latitic-basaltic dykes, diatrems and subvulcanic bodies; CE, Caotico ete- rogeneo; DS, diapiric structure.

Fig. 3.17 - Northern head of Val San Nicolo. To the right an Upper Ladinian latitic-andesitic dyke (Vd) cuts the folds of the diapiric structure in the Upper Permian Bellerophon Fm (B), constraining their Middle Triassic age (Rossi, 1977). Notice that the folds (right bottom) transfer the shortening to a thrust fault that terminates, leaving the pace to other folds (top left). The thrust fault generates at the hinge of an anticline and terminates at the hinge of an overlying syncline. The two tip lines of the small thrust transfer the shorte- ning into the folds. Where the shortening is accommodated by the thrust fault (due to relatively more competent lithologies; central part of the photo), folds are much less frequent. In the more shaly layers the folds show hinge thickening.

Fig. 3.18 - Synsedimentary syncline filled by nodular limestones of the Livinallongo Fm (L), south of the top of Col Ombert mountain south of the diapiric structure of Valle di San Nicolo. This structure testifies the Ladinian transpressional tectonics of the centra Dolomites. The Contrin Fm (C) is folded with synclinal geometry due to a pop-up structure produced by conjugate thrust faults.

Fig. 3.19 - Detail of the previous structure: the southern limb of the synsedimentary syncline, with onlap geometries within the Livinallongo Fm, characterized by massive calcarenites passing laterally into reddish nodular limestones. The picture was taken north of the Massiccio della Costabella, between Passo Pasche and Col Ombert. Alberto Boz for scale.

Fig. 3.20 - Detail of the northern flank of the Avoscan (locality Roi, see also the following picture) diapiric anticline. The Bellerophon Fm (Bf) in the “fiammazza” (evaporitic) facies is in tectonic contact with a slice of the same formation with a “badio- ta” (lagoon) facies (Bb). The main movement plane is characterized by both vertical and horizontal slickenlines. (A) Andraz horizon and (S) Siusi Member of the Scythian Werfen Fm.

Fig. 3.22 - Structural scheme of the Ladinian tectonics of the Dolomites. The following tectonic phases can be distinguished: 1) deve- lopment of the (numbered) diapiric anticlines along the N70°E-trending transpressional Stava-Cima Bocche alignement; 2) trunca- tion of the diapiric structures by the thrust faults of Costabella, characterized by opposite vergence, a feature typical of flower struc- tures, and by the en-echelon thrusts of Marmolada and Col Rodella; 3) development of extensional faults with radial pattern around the Predazzo and Monzoni magmatic centers; such faults possibly accommodated with a dome shape the magma upraise and the emplacement of the radial dykes.

Fig. 3.23 - Relationship between the transpressional deformation of the basement P (top of the Permian Ignimbrites) and the diapiric structures in the sedimentary cover to the north of the Cima Bocche anticline. Notice the following three different structural levels: (1) at their bottom the ignimbrites suffered brittle deformation that triggered the diapiric anticlines in the above layer (2), where the Bellerophon Fm (B) and the Werfen Fm (W) show a more ductile deformation; in the uppermost level (3) the Anisian and Ladinian car- bonates (C) display again brittle deformation and are thrust over the diapiric anticlines of the middle layer, truncating them at their top. In the structural highs of the diapiric anticlines, a Ladinian (Varos) erosional surface, unconformably overlain by a volcaniclastic con- glomerate (CE, Caotico eterogeneo) is at places observed. The monzonitic intrusion (IN) fossilized the deformation.

Fig. 3.24 - The Cima Bocche Anticline involves the basement and the overlying Permian Ignimbrites. Its northern limb is truncated by few subvertical N70-90° trending, possibly left-lateral strike-slip faults along the Passo San Pellegrino.

Fig. 3.25 - Schematic profile through the Sassolungo, Col Rodella and Buffaure massifs. The Cassian Dolomite (CD) developed on the morphological and structural high generated by a sinistral transpressional push-up of Middle Triassic age. In particular, the push-up nucleated beneath a giant olistolite (LO) included in the marine landslide deposits of the Caotico eterogeneo (CE). The section is not retrodeformable due to the strong out-of-section component of deformation. SC: San Cassiano Fm; V, Marmolada Conglomerate, vol- caniclastic deposits and lavas; A, Livinallongo and Contrin Fms.; W, Werfen Fm; B, Bellerophon Fm; G, Gardena Sandstone; I. Ignimbrites, Permian rhyolites and crystalline basement. The main detachments are located in the Bellerophon Fm. Notice the two tran- spressional systems to the north and to the south, respectively the westward prosecutions of the Passo Gardena and Passo Fedaia all- gnments. South of Val di Fassa (right in the drawing), the Buffaure massif represented a sort of small foredeep basin for the push-up structure developed between the two transpressional belts. Here the volcaniclastic deposits are much thicker (800-1000 m) with respect to the 100-200 m thick volcaniclastic sediments in the push-up zone (Sassolungo and Col Rodella Massifs).

Fig. 3.26 - The Crepe Rosse synsedimentary syncline (located near Passo Fedaia) developed right to the south of the Triassic sinistral transpressional alignement of the Fedaia Lake, that is here buried. Notice the onlaps, expecially on the right flank of the syncline, of both volcaniclastic deposits and the interfingering carbonatic megabreccias.

Fig. 3.27 - The southern cliff of the Marmolada massif is extremely complicated by SW-vergent thrusts. B, Bellerophon Fm; W, Werfen Fm; C, Contrin Fm; L, Livinallongo Fm; M, Marmolada Limestone; A, Caotico eterogeneo.

Fig. 3.28 - Northern slope of the Marmolada Ladinian carbonate platform (left), onlapped by Upper Ladinian volcanic and volcanicla- stic rocks of the Padon ridge (right). North to the right, looking west. In the valley the Fedaia lake. Cristina Bagolan and Virginio Doglioni for scale.

Fig. 3.31 - Idealized section through Europe and the Po Plain in the Upper Triassic. Notice the occurrence of a structural high under the Po Plain. Several positive (i.e., transpressional) flower structures (e.g., Dolomites) occur within an area prevalently cha- racterized by subsidence (transtension). Fig. 3.29 - Schematic geological map of the Southern Alps at the Ladinian. Notice the en-echelon distribution of basins and structural highs with respect to the border between the Northern Zone and the Southern Mobile Belt (sensu BRusca, 1981). The distribution suggests a sini- stral shear. In the Northern Zone both carbonate platforms and basins develo- ped (with higher sedimentary thicknesses in the basins), whereas in the Southern Mobile Belt the basement was cropping out (PreERI & Groppi, 1981; BRUSCA et alii, 1981; DE ZANCHE & MIETTO, 1984; GARZANTI, 1985).

Fig. 5.50 - Upper-Middle Iriassic geo- dynamic reconstruction (modified after BERNOULLI & LEMOINE, 1980, BRANDNER, 1984, MASSON & MILES, 1986). Notice the sinistral relative motion between Africa and Europe, coe- val to the Atlantic rift. The magmatic centers (black dots) are localized within rift zones and transform zones characte- rized by mainly transtensional kinema- tics (e.g., Terranova-Azores-Gibraltar; Pyrenees). X-X" is the trace of the sec- tion of the following figure and the rec- tangle indicates the approximate position of the previous figure.

Subject: Analysis of a thrust; Mesozoic archi- tecture.

Fig. 4.2 - View of the Sella Massif from the south. The basal part consists of Cassian Dolomite (D) of Middle Carnian age. The over- lying debris in the middle consists of rocks belonging to the Upper Carnian Raibl Fm (R). The upper cliff is made of Norian Dolomia Principale (DP). Du, Diirrenstein Dolomite, S, San Cassiano Fm; V, Volcanoclastic sandstones; Ce, Caotico Eterogeneo. The thrust fol- lowed the shape of the Carnian platform slope, it continued in ramp into the Dolomia Principale, and with a staircase trajectory arrived at the top of the Piz Boe with its classic summit klippe geometry. The top klippe overlies a condensed Jurassic-Cretaceous succession.

Fig. 4.3 - Val Lasties, looking to the west from the Sella Massif. Beyond the Sella Massif, the Sassolungo Massif can be recognized. Observing the right flank of the valley, the basal cliff is still constituted by upper Cassian Dolomite (CD), the gentle slope by Raibl Fm (R) and the upper cliff by Dolomia Principale (P). Note that the Raibl Fm thickens moving away from the platform: at the outer- most peak of the Sella Massif (Piz Ciavaces), the Raibl Fm is some 80 m thick, whereas it dies out at the core of the Sella Massif. The Sella plateau, like other Dolomitic massifs is characterized by large fossil erosional terraces that according to Rossi (1957) could be related to an Oligocene erosional surface.

Fig. 4.1 - The Sella thrust might represent the front of the WSW-vergent Dinaric fold and thrust belt, of Paleogene-Lower Neogene age. The front of the Dinarides cuts obliquely the eastern Southern Alps and was dissected by later (Neogene) southalpine _ thrusts. Alternatively, the WSW-vergent thrusts in the Dolomites, as discussed earlier, could repre- sent an earlier Alpine right-lateral undulation along the eastern margin of the Trento Horst, conjugate to the left-lateral transpression along the Giudicarie belt, emplaced along the western margin of the Trento Horst.

Fig. 4.4 - Geological map of the Sella Group. Notice the radial progradation of the Carnian carbonate platform (Cassian Dolomite). The Sella group is a perfectly preserved atoll (Leonardi, 1967). The isopachs of the Raibl Fm indicate a thickening of the formation away from the Carnian platform. This suggests a relationship between the geometry of the Raibl Fm deposits and the underlying plat- form-basin system. To the right, south of Campolongo Pass a SSW-vergent thrust fault sutured by basinal deposits of the Sar Cassiano Fm can be observed (DOGLIONI & GOLDHAMMER, 1988).

Fig. 4.5 - Northern slope of the Sella Massif, west of Corvara. The Cassian Dolomite (upper platform D2) progrades northward over the San Cassiano Fm (S2). A lower Cassian Dolomite (D1) and related basin (S1) are below. Du, Diirrenstein Dolomite, R, Raibl Fm.

Fig. 4.6 - Northeastern slope of the Sella Massif. The Cassian Dolomite (upper platform D2) progrades northeastward over the San Cassiano Fm (S2). A lower Cassian Dolomite (D1) and related basin (S1) are below. Du, Diirrenstein Dolomite, R, Raibl Fm.

Fig. 4.7 - Southwestern margin of the Sella massif. The Carnian Cassian Dolomite (C) with its lagoon Diirrenstein Dolomite (Du) pro- grades over the basinal equivalent San Cassiano Fm (S). The Raibl Fm (R) above thins toward the interior of the atoll, for the larger subsidence of the platform margin over the basinal sediments. DP, Norian Dolomia Principale.

Fig. 4.8 - Geological section for the Upper Triassic through the Sella Massif. Notice the wedge shape of the Raibl Fm above the Cassian Dolomite prograding onto the San Cassiano Fm; the Raibl deposits die out in correspondance with the termination, at depth, of the basi- nal San Cassiano Fm. This suggests that the compaction of the Cassian basinal deposits generated the subsidence that allowed the depo- sition of the Raibl Fm. Notice the Triassic tectonic features buried under the Carnian formations (after DOGLIONI & GOLDHAMMER, 1988).

Fig. 4.9 - Photo and geological E-W section of the western mar- gin of the Sella group, towards Passo Sella. The Raibl wedge (R) is evidenced below the the Dolomia Principale (DP). The carbo- nate platform of the Cassian Dolomite (D) progrades onto the basin of the San Cassiano Fm (S). Du, Diirrenstein Dolomite. The platform break surface (PB) and the downlap surface (DL) are convergent, in agreement with the shallowing upward of the San Cassiano Fm. In the panel to the right the letters indicate the sections where decompaction modelling was performed: notice that the decompaction in the basinal deposits of the San Cassiano Fm (shales, marls, calcarenites, volcaniclastic sandsto- nes) is able to explain the ca. 2° tilt of the of the top of the Cassian carbonate platform and the magnitude of subsidence that allowed the deposition of the Raibl Fm deposits. In other words, the syn-Raibl Fm subsidence was controlled by the com- paction of the basinal Cassian deposits (after DOGLIONI & GOLDHAMMER, 1988).

Fig. 4.10 - Southern view of the Piz Boé. The thrusts show staircase geometry with decollements at the top of the Dolomia Principale (P) and of the Dachstein Limestone (D), with connection ramps. Compare with Fig. 4.11 and Fig. 4.12 to appreciate how deformation varies along the strike of the thrust fault. Legend, DP, Dolomia Principale; Da, Dachstein Limestone (after DOGLIONI, 1990).

Fig. 4.11 - Northern view of the Piz Boé. Notice the various triangle structures and the ramp-flat geometry controlled by stratigraphy The basal surface of the thrust system is stratigraphically higher with respect to the southern section of Piz Boé (Fig. 4.10). This is explained by the occurrence of an intermediate lateral ramp (Fig. 4.12). Legend, DP, Dolomia Principale, Norian; Da, Dachsteir Limestone, Rhaetian; Ar, Ammonitico Rosso, M-U Jurassic; K, Puez Marls, Cretaceous (after DOGLIONI, 1990).

Fig. 4.12 - Western view of the Piz Boe (left front page), a natural section along the strike of the thrust system, that verges towards the observer. The structural link between the previous two sections can be observed. Notice the numerous branches and the main lateral ramp connecting the basal decollements of the northern and southern sections. There is, thus, a continuous structural variation along strike. It is interesting to notice, in the hangingwall of the lateral ramp, a fold whose axis is parallel to the transport direction. Legend, DP, Dolomia Principale; Da, Dachstein Limestone; Ar, Ammonitico Rosso; K, Puez Marls.

Fig. 4.14 - Eastern view of Piz Boe. Note the constant undulate trajectory of the thrust planes. At the bottom right, a deep scar in the Norian Dolomia Principale (P) is filled by the Early Cretaceous Puez Marls (K). R, Rosso Ammonitico (Middle-Late Jurassic); D, Dachstein Limestone (Rhaetian).

Fig. 4.15 - Detail of the left upper part of the previous drawing. The topmost thrust of the imbricate system of thrusts of the Sella Massif showing undulated geometry with ramps and flats. Da, Dachstein Limestone; Ar, Ammonitico Rosso.

Fig. 4.17 - Geological map of Piz Boé, at the top of the Sella Massif (after DOGLIONI, 1990).

Fig. 4.18 - Geological cross sections across Piz Boé. A is a longitudinal section showing the lateral ramp and the thrust branches that allow the continuous variation of the geometries along strike. The structure is very irregular although the shortening (about 1 km) remains quite constant. The numbers indicate the reconstructed kinematics of the thrust faults. The occurrence of deep scars fillec by Puez Marls explains the tectonic transport of Cretaceous sediments above older rocks (e.g., section C). The location of the cross sections is in the previous figure.

Fig. 4.19 - General view of the top of the Sella Massif, i.e., the Piz Boé with its klippe made of a stack of thrusts with small lateral continuity and a number of branch lines. Fig. 4.20 - Stratigraphic relationships in the upper part of the Sella Massif and location of the thrust faults. Sequence boundaries are frequently detachment planes.The thrust faults cut pre-existing Mesozoic extensional faults. Notice also the Cretaceous erosional surface. CD, Cassian Dolomite (Middle Carnian); DU, Diirrenstein Dolomite (Upper Carnian); P, Dolomia Principale (Norian); D, Dachstein Limestone (Rhetian); R, Ammonitico Rosso (Middle-Upper Jurassic); K, Puez Marls (Lower Cretaceous).

Fig. 4.22 - Eastern ramp of the Sella Thrust, where two klippen of Dolomia Principale (DP) crop out in the northeastern part of the mas- sif. View from the road between Corvara and the Campolongo Pass. Du, Diirrenstein Dolomite; D, Cassian Dolomite.

Fig. 4.23 - In the Vallon del Serla (Serla canyon), east of Piz Boe, an Upper Carnian N-S oriented graben, sealed by the Dolomia Principale, crops out undeformed in the footwall of the Sella thrust. P, Dolomia Principale (Norian) above the unconformity; DU, Diirrenstein Dolomite; CD, Cassian Dolomite.

Fig. 4.25 - Geological section of the eastern flank of the Sella Massif. The eastward continuation of the Sella thrust (ramp in the Dolomia Principale and flat at its top) can be observed. Towards the bottom of the section, the thrust plane is partly eroded and crops out again more to the east in the meadows of the Altipiano del Cherz-Pralongia. A blind thrust, however, occurs above the slope of the Carnian carbonate platform (Cassian Dolomite) and is associated to a fault propagation fold. The inclined stratigraphic plane was thus a hetero- geneity that controlled the distribution of the shortening. N-S trending synsedimentary extensional faults disrupted the continuity of the slope. They are upper Carnian (syn-Diirrenstein Dolomite) or younger (after DOGLIONI, 1992).

Fig. 4.24 - Eastern flank of the Sella Group and Campolongo Pass. The black profile shows the trace of the section of Fig. 4.25. Along this section the eastward continuation of the Sella thrust to lower stratigraphic levels can be observed. CD, Cassian Dolomite (Middle Carnian); S, San Cassiano Fm (Middle Carnian); Du, Diirrenstein Dolomite (Upper Carnian); R, Raibl Fm; P, Dolomia Principale (Norian). In the foreground, near Arabba, the Liviné Line (a thrust fault sealed by the San Cassiano Fm, 600-700 m thick to the left and only 150-200 m to the right) is indicated.

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