1 Ivrea mantle wedge, arc of the Western Alps,
2 and kinematic evolution of the Alps-Apennines
3 orogenic system
4 Stefan M. Schmid1, Eduard Kissling1, Tobias Diehl1, Douwe J.J vanHinsbergen2,
5 Giancarlo Molli3
6 1Institute of Geophysics, ETH Zürich, Sonneggstrasse 5, 8092 Zürich, Switzerland
7 2Department of Earth Sciences, Utrecht University, Heidelberglaan 2, 3584 CS, Utrecht, The
8 Netherlands
9 3Dipartimento di Scienze della Terra, Università di Pisa, Pisa, Italy
10 Manuscript submitted to Swiss Journal of Geosciences, version October 2016
11 Abstract
12 The construction of five crustal-scale profiles across the Western Alps and the Ivrea
13 mantle wedge integrates up-to-date geological and geophysical information and
14 reveals important along strike changes in the overall structure of the crust of the
15 Western Alpine arc. Tectonic analysis of the profiles, together with a review of the
16 existing literature allows for proposing the following multistage evolution of the arc
17 of the Western Alps: (1) Exhumation of the mantle beneath the Ivrea Zone to shallow
18 crustal depths during Mesozoic is a prerequisite for the formation of a strong Ivrea
19 mantle wedge whose strength exceeds that of surrounding mostly quartz-bearing
20 units, and consequently allows for indentation of the Ivrea mantle wedge and
21 eastward back-thrusting of the western Alps during Alpine orogeny. (2) A first early
22 stage (pre-35 Ma) of the West-Alpine orogenic evolution is characterized by top-
1 23 NNW thrusting in sinistral transpression causing at least some 260km displacement
24 of internal Western Alps and E-W-striking Alps farther east, together with the Adria
25 micro-plate, towards N to NNW with respect to stable Europe. (3) The second stage
26 (35-25 Ma), further accentuating the arc, is associated with top-WNW thrusting in the
27 external zones of the central portion of the arc and is related to the lateral indentation
28 of the Ivrea mantle slice towards WNW by some 100-150km. (4) The final stage of
29 arc formation (25-0 Ma) is associated with orogeny in the Apennines leading to
30 oroclinal bending in the southernmost Western Alps in connection with the 50°
31 counterclockwise rotation of the Corsica-Sardinia block and the Ligurian Alps.
32 Analysis of existing literature data on the Alps-Apennines transition zone reveals that
33 substantial parts of the Northern Apennines formerly suffered Alpine-type shortening
34 associated with an E-dipping Alpine subduction zone and were backthrusted to the
35 NE during Apenninic orogeny that commences in the Oligocene.
36
37
38 A. Introduction
39 The Western Alps represent a prime example of a very arcuate mountain belt
40 curving around from an E-W strike with dominant top-N nappe stacking in the
41 Eastern Alps to a N-S strike in the northern Western Alps and finally into an E-W
42 strike with top-S nappe stacking in the Ligurian Alps straddling the Mediterranean
43 coast (Fig. 1). The question after the mechanisms of arc formation is an old and much
44 debated issue in many orogens (Marshak 1988) involving, in the case of the Alps,
45 elements of oroclinal bending of an initially straight orogen (Carey 1955, van der
46 Voo 2004), induced by rigid body rotation of the hinterland (Vialon et al. 1989)
47 and/or by W-directed indentation of the Adria microplate (Laubscher 1971; Schmid
2 48 & Kissling 2000). Transpressional movements associated with large top NNW
49 displacements of portions of the inner zone of the Western Alps that occur at a small
50 angle to the present-day N-S strike of the Western Alps were also invoked for being
51 responsible for explaining the curvature from the Eastern Alps into the N-S striking
52 Western Alps (Ricou & Siddans 1986). There have also been attempts to explain
53 diverging kinematic indicators along the arc of the Western Alps ranging from top-N
54 Western Switzerland to top-SW in southern France without taking resort to a multi-
55 stage evolution allowing for strike-slip movements and/or large rotations (Platt et al.
56 1989; their group 3 kinematic indicators formed in the 40-15 Ma time interval).
57 However, as pointed out by Butler (1986; his Fig. 2), strongly divergent thrust
58 transport directions over large distances inevitably result in excessive orogen-parallel
59 stretching.
60 In the Western Alps many previous workers (e.g. Laubscher 1971; Butler 1983)
61 recognized the crucial role of the Ivrea mantle wedge residing at relatively shallow
62 crustal levels and locally reaching the surface. Novel geophysical methods such as
63 teleseismic tomography aiming at the structure of the earth’s mantle (Lippitsch et al.
64 2003) and crustal tomography based on local earthquake data (Diehl et al. 2009),
65 together with older information based on controlled source seismology (Blundell et
66 al. 1992), led to substantial recent improvements in revealing the deep structure of the
67 arc of the Western Alps. Th new geophysical data set of unprecedented quality calls
68 for integration with a wealth of geological data collected over many decades to
69 elucidate the formation of the arc of the Western Alps that also demands to analyze
70 the Alps-Apennines transition area located in the Ligurian Alps and the Northern
71 Apennines (Molli et al. 2010).
72 This contribution integrates the new results of crustal tomography and
3 73 improved knowledge on Moho depth (Spada et al. 2013) with a geological
74 interpretation of five crustal-scale transects across the Western Alps. Thereby we
75 focus on the role of the Ivrea mantle wedge. Another important aim of our
76 geological-geophysical interpretations is to put new constraints on the question as to
77 how much continental crust was subducted in the Alps and what the variations in pre-
78 orogenic crustal thickness might have been (see discussions by, e.g., Ménard &
79 Molnar 1991; Butler et al. 2013; Mohn et al. 2014).
80 Two transects are revised versions of previously published profiles (Schmid et
81 al. (2004) while three of them were newly constructed. Special emphasis is given to
82 improved mapping of the outlines of the Ivrea mantle wedge based on geophysical
83 data. This is then followed by an analysis of the complex multi-stage evolution of the
84 arc of the Western Alps and an attempt to map the associated displacement history of
85 geological units of the Western Alps-Northern Apennines system using plate
86 reconstruction methodology (GPlates free software (Boyden et al., 2011) embedded
87 in the global plate kinematic reconstruction of Seton et al. 2012). Finally, we provide
88 a review-type discussion of the Alps-Apennines transition area and discuss
89 controversial issues concerning questions of multi-stage vs. single subduction
90 systems in the Apennines.
91
92
93 B. Crustal-scale profiles across the Western Alps and the Ivrea mantle wedge
94 integrating geological and geophysical information
95 To constrain the geometry of the Ivrea mantle wedge, and as a basis for
96 discussing its role during the formation of the arc of the Western Alps, we
97 constructed a series of crustal scale profiles along the arc of the Western Alps (Fig. 1).
4 98 These profiles integrate geological data with recent geophysical data focusing on the
99 deep structure. The construction of the profiles is based on a series of assumptions,
100 the most important one being that advocated by Schmid & Kissling (2000), namely
101 that the European lower crust is decoupled from the upper crust and goes into
102 subduction together with the underlying mantle for considerable distance below the
103 Ivrea mantle wedge. This invokes a basal detachment of the external thrust sheets at
104 the interface between upper and lower crust at a depth of 20-30km. While this
105 assumption is compatible with the velocity structure (see below) and recent
106 geophysical findings of Zhao et al. (2015) it is by no means generally accepted (see
107 discussion in Butler 2013). However, there are considerable uncertainties about the
108 exact pre-orogenic thickness of the European and Briançonnais continental crust in
109 more internal parts of the profiles and this is one reason why we restrain from area
110 balancing over the entire length of our transects. The main reason is that transport
111 directions changed with time during the formation of the arc of the Western Alps,
112 making 2D area balance along the transects impossible.
113 The geological literature used for constructing the individual profiles varies
114 and hence will be cited when presenting the profiles. However, we used one and the
115 same set of geophysical data for constraining the deep structure in all the profiles.
116 Firstly, we used the results of a tomographic high-resolution 3-D P-wave model of
117 the Alpine crust obtained by Diehl et al. (2009). This 3-D velocity model was used
118 for constructing parts of the sections of Fig. 2 that present contours of equal P-wave
119 velocity superimposed with a geological interpretation. When comparing the velocity
120 contours with the geological interpretation it has to be kept in mind, however, that
121 these contours have a low spatial resolution since they are based on assigning a given
122 velocity to a sizable cell volume of 20x20 km (horizontal) and 15km (vertical).
5 123 Secondly, we complemented these tomographic data by geophysical data better
124 constraining Moho-depth. Where available we took Moho depths from Spada et al.
125 (2013; their Fig. 11; see also Wagner et al. 2012 regarding the methodology) who
126 combined results of controlled-source seismology with local receiver functions Moho
127 depth maps and local earthquake tomography-derived Moho maps.
128 Comparison between Moho depth compiled by Spada et al. (2013),
129 characterized by a P-wave velocity jump, and the velocity model of Diehl et al.
130 (2009) that defines Moho depth in terms a high velocity gradient in the range of 6.5–
131 8.0 km s−1, noticeable over large parts of the profiles, shows that on average the
132 Moho as defined by Spada et al. (2013) best corresponds with the velocity contour of
133 7.25 km s−1 in areas characterized by a subhorizontal or only weakly inclined Moho
134 (Diehl et al. 2009). In order to define the strongly inclined boundaries of the Ivrea
135 mantle wedge at shallower depth <30km, not well constrained by the compilation of
136 Spada et al. (2013), we followed the 7.0 km s−1 contour in all profiles.
137 In the profiles ECORS-CROP, NFP20-West and Ticino rather thin slices of the
138 Ivrea wedge and/or the lower crust of the Adria microplate, geologically speaking the
139 Ivrea Zone (Handy & Zingg 1991), reach the surface. In these profiles the outlines of
140 the Ivrea mantle wedge are too thin to be resolved by the P-wave velocity model of
141 Diehl et al. (2009) given its limited spatial resolution. The geometry of the very thin
142 and strongly dismembered Ivrea mantle wedge and Adriatic lower crust is only
143 tentatively drawn, being only constrained by the surface exposures.
144 Another problem one faces when determining the shape of the Ivrea mantle
145 wedge is that the velocities >7.25 km s−1 expected for the Ivrea mantle rocks that are
146 part of the Adriatic microplate and the adjacent European mantle become
147 indistinguishable from velocities expected for potential remnants of the subduction
6 148 channel separating European and Ivrea (=Adriatic) mantle marking the present-day
149 plate boundary at great depth. At a depth of 30-45 km and temperatures around
150 400°C- 500°C, we expect to reach blueschist facies conditions, the exact depth
151 depending on temperature (Oberhänsli et al. 2004). Based on thermal modeling
152 (Bousquet et al. 1997) the expectation of temperatures between 450° and 500°C at a
153 depth of 30km is realistic. The P-wave velocity of blueschist rocks at 400°C and 1
154 GPa is around 7.25 km s−1 on average according to Fujimoto et al. (2010). From a
155 depth of around 50km onward, again depending on temperature, we expect eclogite
156 facies conditions associated with even higher velocities around 8.0-8.5 km s−1
157 (Worthington et al. 2013). Hence, the contours of velocities >7.25 km s-1 are expected
158 to cross over from the European to the Ivrea mantle rocks traversing the subduction
159 channel formed by high-P metamorphic rocks. Such crossover is actually seen in all
160 the profiles of Fig 2, also for velocity contours at 7.0, 6.5 and occasionally at 6.5km
161 s−1. This is because velocities expected for the subducted European lower crust
162 (around 6.5 km s−1; Ye et al. 1995) are similar to those expected for the stack of
163 nappes that is northerly adjacent to the Ivrea mantle wedge, nappes that are only
164 partially overprinted by high-P metamorphism (e.g. Henry et al. 1993). All this,
165 together with the limited resolution of the 3-D velocity model, puts severe limits to a
166 straightforward correlation of velocities with geological-tectonic entities, particularly
167 at greater depths.
168
169 1. Argentera profile (Fig. 2a)
170 The external SW part of this profile is characterized by minor shortening of
171 around 21km in Mesozoic sediments (Chaînes Subalpines) of the European foreland
172 (Lickorish & Ford 1998), associated with rather modest late-stage up-thrusting of the
7 173 Argentera external massif (Malaroda et al. 1970) during the Neogene. In this external
174 part of the profile the European Moho coincides with a strong velocity gradient
175 between 6.5 and 7.5 s-1 in the velocity model of Diehl et al. (2009).
176 The still unfolded Chaînes Subalpines were thrust by sedimentary units
177 attributed to the Subbriançonnais, together with non-metamorphic Helminthoid flysch
178 detached from the Piedmont-Liguria paleogeographical domain above ophiolite-
179 bearing mélanges (Autapie and Parpaillon nappes; Merle & Brun 1984). Final
180 emplacement of these higher nappes occurred during the Late Eocene (Kerckhove
181 1969), i.e. before thrusting along the late Alpine (Early Oligocene) out-of-sequence
182 thrust called “Penninic Front” and Neogene shortening in the Chaînes Subalpines.
183 Since the Autapie and Parpaillon nappes lack metamorphic overprint they are
184 considered as units derived from a part of the Piedmont-Liguria Ocean that always
185 remained in an upper plate position. They form relics of the front of the overriding
186 oceanic upper plate (Piedmont-Liguria Upper Plate units of Fig. 2a) that started to
187 thrust the oceanic lower plate (Piedmont-Liguria Lower Plate units of Fig. 2a) from
188 the late Cretaceous onwards, reaching the front of the Alps at the end of the Eocene,
189 together with the underlying Subbriançonnais units.
190 There is a rather abrupt lateral change of velocities at a given depth across the
191 boundary between external and internal Alps. This change coincides with the
192 downward projection of the southern end of a very steeply NNE dipping late Alpine
193 out-of-sequence thrust well documented in the Pelvoux section and known as the
194 Penninic Front (Ceriani et al. 2001; Lardeaux et al. 2006; Loprieno et al. 2011; see
195 description of the Pelvoux section). The term “Penninic Front”, synonymous with the
196 term “Briançonnais Frontal Thrust” coined by Tricart (1986) who first recognized the
197 importance of this Early Oligocene out-of-sequence thrust in marking the boundary
8 198 between internal and external parts of the Western Alps, is a misnomer in that earlier
199 emplaced Penninic units are occasionally also found in front of this out-of-sequence
200 thrust (Ubaye-Embrunais of profile in Fig. 2; Préalpes Romandes of Fig. 2d; see also
201 Fig. 1). We use the term Penninic Front nevertheless because it became deeply
202 entrenched in the literature by now. This top WMW Penninic Frontal Thrust turns
203 into the sinistral Acceglio-Cuneo Line (Fig. 1) and is characterized by a sinistrally
204 transpressive component as it swings around into the WNW-ESE strike of the
205 Argentera-Cuneo line whose western end is traversed by the Argentera profile. The
206 sinistral Acceglio-Cuneo Line allows for the WNW-directed indentation of the Ivrea
207 mantle wedge along its southern boundary located near Cuneo (Trullenque 2005;
208 Ceriani et al. 2001, see their Fig. 1). Laubscher (1971) first proposed the existence of
209 a sinistral strike slip zone (his “Ligurian shear zone”) allowing for westward
210 indentation of the Ivrea mantle wedge for kinematic reasons (see also Laubscher
211 1991). Based on the interpretation of fission track data (Fügenschuh & Schmid 2003)
212 and radiometric dating of mylonites (Simon-Labric et al. 2009) the thrust marking the
213 Penninic Front (Roselend thrust of Loprieno et al. 2011) was active during the Early
214 Oligocene (34-30 Ma). By downward extrapolation the steep dip of this crustal scale
215 transpressive fault projects (Malaroda et al. 1970) into parallelism with the steeply
216 overturned 6.5 s-1 velocity contour that marks the SW edge of a high-velocity area
217 with > 6.5 s-1 located within the high-pressure Briançonnais units at a depth of
218 >10km. This high velocity area is interpreted to consist of the high-pressure
219 Briançonnais units that include the Dora Maira “massif” exposed at the surface at the
220 eastern margin of the Internal Alps (Henry et al. 1993). We are aware that not all
221 exposed parts of the Dora Maira massif constituting the high pressure Briançonnais
222 show blueschist facies or eclogitic parageneses (Bousquet et al. 2012a). However, the
9 223 high velocities at depth require that most of the volume taken up by the Briançonnais
224 at depth must have undergone high-pressure metamorphism.
225 The shallow crustal parts of the metamorphic core of the Alps in this profile are
226 characterized by large-scale back-folds that formed between 35-30 Ma (Bucher et al.
227 2004) and which affected pre-existing nappe contacts between: (1) the high-pressure
228 Briançonnais units, (2) overlying so-called Acceglio and/or Pre-Piémontais units
229 (Lemoine & Tricart 1986; Caby 1996), representing the former transition of the
230 Briançonnais micro-continent into the Piedmont-Liguria Ocean, (3) Piedmont-Liguria
231 high pressure units and, finally, (4) non-metamorphic Piedmont-Liguria units that are
232 only preserved along strike in the area of the Chenaillet near Briançon (Manatschal et
233 al. 2011). We attribute the Chenaillet ophiolites to that part of the Piedmont-Liguria
234 Unit that occupied an upper plate position in the internal Western Alps (Schwartz et
235 al. 2007) where, in our interpretation, the Austroalpine nappe system is totally
236 lacking due to a westward lateral transition of the active margin of the Late
237 Cretaceous to Paleogene Alpine subduction zone from the northern edge of the Adria
238 continent into an intra-oceanic location (Molli, 2008; Handy et al. 2010). The
239 geometry of the shallow crustal parts of the Internal Alps depicted in this profile
240 closely follows the geometries shown by the works of Henry et al. (1993), Michard et
241 al. (2004), Lardeaux et al. (2006), and Ford et al. (2006), the surface-near parts of the
242 Penninic Front exposing the Subbriançonnais and non-metamorphic Briançonnais
243 units was drawn after a series of detailed profiles given by Gidon (1972).
244 The 7.0 s-1 velocity contour that runs parallel and near the Moho depicted by
245 controlled-source seismology in the Adria microplate rises SW-ward to a depth of
246 some 15 km in the NE part of the profile beneath the HP Briançonnais. It is taken as
247 marking the outline of the top of the Ivrea mantle wedge. The inferred overall
10 248 geometry demonstrates that crust and mantle of the Adriatic microplate became
249 wedged into the Alpine edifice and underplated shallow crustal layers in the internal
250 parts of the Alps. The dip of the Periadriatic line is constrained to a depth of some 30
251 km by connecting the southern prolongation of the Periadriatic Line (Fig. 1),
252 delimiting the eastern edge of the root of the Piedmont-Liguria units near the eastern
253 margin of the Dora Maira massif at the surface, with the western front of the Ivrea
254 mantle wedge paralleling the 7.0 s-1 contour. The exact geometry shown for the
255 downward narrowing deepest parts of internal Alps, i.e. the subduction channel
256 separating the European and Adriatic (micro) plates, is basically unconstrained by
257 geophysical data and the details shown are merely conceptual. The downward
258 bending of the 7.0 and 7.5 s-1 velocity contours across the suspected subduction
259 channel is unable to depict the geometry of the subduction channel, due to the high
260 velocities expected for the rocks within the subduction channel undergoing high-
261 pressure metamorphism, velocities expected to be similar to those of the European
262 lower crust, and at greater depth, similar to those of the adjacent mantle rocks.
263 The Oligocene age of formation of the steeply dipping root of the Piedmont-
264 Liguria Ocean, associated with rock uplift associated with exhumation (Fox et al.
265 2016) of these high-pressure units along the front of the adjacent Adria microplate
266 can be inferred from the cooling history of the Dora Maira high pressure unit
267 evidencing rapid cooling from 550°C to some 260°C within an extremely short time
268 interval between 32 and 30 Ma according to the pressure-time path constrained by
269 petrological data in combination with cooling ages derived from isotopic dating and
270 fission track analysis (Gebauer et al. 1997; Rubatto & Herrmann 2001; Berger &
271 Bousquet 2008). The earlier stages of this exhumation are possibly related to
272 exhumation within the subduction channel, triggered by the buoyancy of continental
11 273 rocks at mantle depth (Chemenda et al. 1995). These time constraints indicate that
274 both the formation of the steeply hinterland-dipping Penninic Front and the wedging
275 of the Adria mantle and crust associated with backthrusting of the internal Alps over
276 the Tertiary Piedmont Basin are contemporaneous. Both appear to be associated with
277 the WNW-directed indentation of the Ivrea mantle during the earliest Oligocene (34-
278 30Ma).
279 The shallow NE crustal parts of the profile are covered by the Oligo-Miocene
280 of the Tertiary Piedmont Basin and younger Plio-Pleistocene sediments. The
281 construction of the profile is based on subsurface data acquired by AGIP (Pieri &
282 Cropi 1981, Cassano et al. 1986, Schumacher & Laubscher 1996) that only vaguely
283 constrain the geometry depicted. The 5.5 to 7.0 velocity contours, as well as the
284 Moho depth indicated by controlled-source seismology, indicate that within the Adria
285 microplate Moho, lower crust and basement-sediment interface are E-dipping.
286 In the immediate vicinity of the Periadriatic line the thickness of the Tertiary
287 Piedmont Basin along the profile is considerable. Carrapa & Garcia-Castellanos
288 (2005) showed that the subsidence of the Tertiary Piedmont Basin is controlled by
289 far-field compression and thrust loading rather than by extensional tectonics. Hence,
290 the early stages of thrust loading are probably due to the Early Oligocene backthrust
291 of the metamorphic units of the Western Alps onto the Adria microplate.
292 However, at the eastern end of the profile the Tertiary Piedmont Basin overlies
293 “Apenninic” units (Fig. 2a). Compressional tectonics continued into the Miocene
294 when the Ligurian Alps became involved in N-directed thrusting, together with
295 Apenninic units exposed in the northern Monferrato hill (Fig. 1) west of Torino
296 (Piana 2000), first during the so-called Paleo-Apenninic phase (Schumacher &
297 Laubscher 1996) sealed by a Mid-Miocene unconformity. Later, Neo-Apenninic
12 298 thrusting formed the morphologically visible thrust front north of the Torino and
299 Monferrato hills. It is this N-directed Apenninic thrusting over the Adria microplate
300 that involved parts of the former Alps referred to as “Apenninic” units in the profile
301 of Fig. 2a (see later discussion on the Alps-Apennines transition). Differential N-
302 directed thrusting of joint Apenninic and Ligurian Alpine units calls for a sinistral
303 strike slip zone that separates the joint Apennines-Ligurian Alps unit moving to the
304 north from those parts of the Piedmont-Liguria units that are immediately adjacent to
305 the Dora Maira backthrust, such as depicted near the eastern end of the section (Fig.
306 1a). The exact 3D geometry of this complicated Early Miocene to recent scenario still
307 remains largely unresolved and is known as the problem of the “Ligurian knot”
308 (Laubscher et al. 1992).
309
310 2. Pelvoux profile (Fig. 2b)
311 The western part of this profile crosses the Vercors massif (Chaînes
312 Subalpines) and the more internal Belledonne–Grandes Rousses–Oisans external
313 massifs and their Mesozoic cover (Philippe et al. 1998; Dumont et al. 2008) (Fig. 1).
314 Profile balancing by Bellahsen et al. (2012), taking into account the effects of normal
315 faulting preceding Alpine orogeny (e.g. de Graciansky et al. 1989) related to the
316 opening of Alpine Tethys, the total amount of Miocene shortening in these external
317 parts of the Alps (Dauphinois) amounts to some 28km.
318 In the external parts of the profile the P-wave velocity contours within the
319 European lithosphere again exhibit a higher velocity gradient between 6.5 and 7.5 s-1
320 near the Moho. However, the 6.0 s-1 contour rises to an unusually shallow depth of
321 <10km twice, a feature that is considered robust and points to the presence of high-
322 velocity material at shallow depth. The more external amongst the two anomalies is
13 323 found underneath the Belledonne and Grandes Rousses massifs, an area that is known
324 for the occurrence of the pre-Alpine Chamrousse ophiolitic complex (Guillot &
325 Ménot 2009), associated with ophiolites and eclogites being part of a former Variscan
326 suture zone. This suture zone evolved into a major SW-NE striking and ca. 1500 km
327 long dextral wrenching zone in Late Carboniferous times, marking the SE margin of
328 the European Variscides with Gondwana. This broad mega shear zone is known as
329 the External Massifs Shear Zone (Guillot & Ménot 2009) or East Variscan Shear
330 Zone (Padovano et al. 2012). The interpretation of this anomaly in terms of high-
331 velocity material constituting this steeply inclined mega shear zone is admittedly
332 speculative. The origin of the second more internal anomaly is unknown but possibly
333 also related to velocity inhomogeneities related to pre-Alpine structures.
334 At the surface the Penninic Front is exposed as a spectacular thrust zone behind
335 the Pelvoux massif and in the footwall of the Massif du Combeynot, dipping with ca.
336 35° to the E and associated with top-WNW out-of-sequence thrusting (Butler 1992;
337 Ceriani et al. 2001; Ceriani & Schmid 2004, Trullenque et al. 2006). The linear
338 extrapolation of this dip observed at the surface to depth is unconstrained by currently
339 available geophysical data along this profile. Fig. 2b opts for an extrapolation of the
340 Penninic Front to great depth such as to join the subduction channel at the base of the
341 Ivrea mantle wedge, consistent with what was observed along the Argentera profile
342 described above. The 6.5 and 7.0 s-1 contours that followed the lower crust of the
343 European lithosphere in the external Western Alps start to rise to shallower levels at
344 an angle of around 45° after having crossed the Penninic Front, indicating that
345 westward they enter high-pressure Briançonnais units.
346 The shallow crustal parts of the metamorphic core of the Alps are drawn
347 according to data published in Barféty et al (1995), Caby (1996), Ganne et al. (2006),
14 348 and Lardeaux et al. (2006). In the Briançonnais units sedimentation continued until
349 the Midle (internal Briançonnais) to Late Eocene (external Briançonnais) (Jaillard
350 1999) and hence predates thrusting at the Early Oligocene Penninic Front. They very
351 much resemble those found along the Argentera profile. Again, almost the entire
352 section across the inner Western Alps, except the very frontal part, is characterized by
353 late-stage backfolding of pre-existing nappe contacts. The geometry of the top of the
354 Ivrea mantle wedge is constrained by the 7.0 s-1 velocity contour and controlled
355 seismology data. Exact location and width of the subduction channel remain
356 unknown, but an inverse velocity gradient of the 7.5 s-1 contour associated with a
357 down-bending of the same contour line points to the existence of a subduction
358 channel below the Ivrea mantle wedge. A similar, although not identical geometry of
359 a geological transect located near the Pelvoux profile was proposed by Zhao et al.
360 (2015), who also postulated subduction of European continental lower crust all the
361 way below the Ivrea mantle wedge, based on their interpretation of seismic imaging
362 by P-wave receiver function analysis. In comparison, their model proposes a much
363 shallower depth for the top of the Ivrea mantle wedge (at around 10km) and a much
364 wider (about 30km) subduction channel below the Ivrea mantle wedge, features that
365 are not compatible with the 3-D velocity model of Diehl et al. (2009) and the Moho
366 depth as compiled by Spada et al. (2013).
367 The eastern parts of the profile are very similar to what was described for the
368 Argentera section. They cross the northern parts of the Dora Maira massif and exhibit
369 the same overturned root of the ophiolitic series of the Piedmont Lower Plate unit,
370 juxtaposed with the front of the Adria microplate including the Ivrea mantle wedge
371 along a backthrust onto the Tertiary Piedmont Basin, formed in the earliest
372 Oligocene, i.e. contemporaneous with the formation of the Penninic Front. At its
15 373 easternmost end the profile also reaches the Apenninic units thrust northward and
374 perpendicular to the trace of the profile towards the Torino Hills in the Late Miocene.
375
376 3. ECORS-CROP profile (Fig. 2c)
377 This transect follows the trace of the geophysical campaign launched in 1985
378 by the French-Italian Étude Continentale et Océanique par Reflexion et Refraction
379 Sismique - Progetto Strategico Crosta Profonda (ECORS-CROP) geophysical
380 campaign in the Western Alps involving high-resolution deep seismic sounding
381 Roure et al. (1996). The geological interpretation of this profile is a modified version
382 of the profile previously published by Schmid & Kissling (2000).
383 The external Alps exhibit 27km shortening, an amount very similar to that
384 along the Pelvoux profile, calculated on the basis an area balance within an originally
385 17km thick upper crust, assuming decoupling at the interface with the lower
386 European crust. This shortening is of Miocene age and postdates thrusting at the
387 Penninic Front (Fügenschuh & Schmid 2003). Note that Butler (1983) postulated
388 very substantially higher amounts of shortening (>50km) since his line balance
389 assumes detachment along a floor thrust located only 1km below base Triassic. The
390 6.5 km s−1 velocity contour beneath the external massifs, in the case of this profile the
391 Belledonne Massif, indicates the presence of a steep tabular high-velocity body in the
392 lower part of the upper crust that we again (see description Pelvoux profile)
393 tentatively attribute to the East Variscan Shear Zone of Late Variscan age (Padovano
394 et al. 2012). The velocities within the rest of the European upper crust are, apart from
395 this anomaly, < 6.5 km s−1 all the way back to the very internal parts of the Alps
396 below the Gran Paradiso massif, built up the high pressure Briançonnais unit. This
397 new finding about the P-wave velocity in the upper crust of the external part of the
16 398 Western Alps (Diehl et al. 2009) is incompatible with the presence of an
399 allochthonous wedge of lower crust underneath the external Western Alps as
400 postulated by Schmid & Kissling (2000) on the basis of gravity arguments (Bayer et
401 al. 1989) and W-dipping seismic reflectors (Thouvenot et al. 1996) at the front of that
402 putative wedge. Below the Gran Paradiso massif the 6.5 km s−1 contour steeply raises
403 to a depth of only some 10 km, across the Penninic Front and into the area of the
404 high-pressure part of the Briançonnais units, indicating a high-pressure metamorphic
405 overprint also at great depth.
406 At the surface the profile crosses the type locality of the top-WNW Penninic
407 Front (“Roselend Thrust” of Loprieno et al. 2011). This thrust, together with a 10 km
408 wide belt of units attributed to the Valaisan paleogeographic domain, containing,
409 amongst other lithologies, black schists interlayered with mafic sills and low-T
410 eclogites (Bousquet et al. 2002, 2012a; Loprieno et al. 2011) forms a planar band of
411 strong reflectors that can be traced down to a depth of some 15 km along the
412 ECORS-CROP reflection seismic profile (Thouvenot et al. 1996). Following Schmid
413 & Kissling (2000), and in analogy with the Pelvoux profile discussed above, we
414 kinematically link the Penninic Front with the base of the Adria mantle wedge.
415 The near-surface geometry of the upper crustal structures within the
416 Briançonnais units, namely the non-metamorphic Zone Houllière and the High-
417 Pressure Briançonnais units (Ruitor, Zona Interna and Gran Paradiso nappes), as well
418 as the geometries of the Piedmont-Liguria Units (Zermatt-Saas and Combin Zones)
419 are drawn following Schmid & Kissling (2000) and Bucher et al. (2003, 2004). They
420 are characterized by large-scale post-nappe backthrusts and backfolds formed
421 between 35 and 31 Ma, i.e. contemporaneous with thrusting along the Penninic Front.
422 High-pressure metamorphism related to subduction was dated at 48-47 Ma in the
17 423 Briançonnais units, overprinted by greenschist facies conditions related to nappe
424 stacking at around 43-39 Ma (Villa et al. 2014). Pressures associated with the high-
425 pressure event increase eastward and reach eclogite facies conditions in parts of the
426 Gran Paradiso internal massif, characterized by P-wave velocities around 6.0 to
427 almost 7.0 km s−1 at greater depth according to the velocity model of Diehl et al.
428 (2009).
429 Compared to the Briançonnais units high-pressure metamorphism within the
430 more internal Zermatt-Saas-Fee unit (Piedmont-Liguria Lower Plate of Fig. 2c) and
431 within small continental slices at its top indicates an earlier onset of a long and
432 protracted high P overprint (62-42 Ma) according to age dating (e.g. Dal Piaz et al.
433 2001; Skora et al. 2015; Weber et al. 2015; Fassmer et al. 2016). The most internal
434 parts of the ECORS-CROP profile cross the Sesia Zone, which together with the
435 Dent Blanche Klippe, was classically considered as an “Austroalpine” unit. This is
436 because the protoliths forming the Sesia Zone are undoubtedly derived from the
437 Adria microplate in a paleogeographical sense (Compagnoni et al. 1977). However,
438 in a plate tectonic context, the Sesia Zone that reached eclogite facies conditions with
439 up to 2 GPa pressures in the period 85-60 Ma (Regis et al. 2014) and consisting of
440 thin crustal sheets (Giuntoli & Engi 2016) is interpreted as having been accreted as a
441 part of the former Lower Plate to the Upper Plate Units (i.e., the Canavese, Ivrea
442 Zone and Southern Alps) during Late Cretaceous to early Cenozoic orogeny in the
443 Western Alps, associated with the subduction of the Piedmont-Liguria ocean (Babist
444 et al. 2006, Handy et al. 2010). The situation is different in the Austroalpine nappes
445 proper where high-pressure metamorphism is significantly older and related to
446 Cretaceous intra-continental subduction that evolved within the future Austroalpine
447 units that formed the upper plate during Cenozoic orogeny (Schmid et al. 2004). In
18 448 summary, the Sesia Zone is regarded as a lower plate unit following Froitzheim et al.
449 (1996; their Fig. 2) and Pleuger et al. (2007; their Fig. 14) who proposed that the
450 Sesia Zone represents a continental fragment floating within the Piedmont-Liguria
451 oceanic domain that rifted and drifted away from Adria during the opening of Alpine
452 Tethys in the Mesozoic. Two important geological observations support this finding.
453 Firstly, the southern edge of the mostly continental Sesia high-pressure unit,
454 immediately adjacent to the Insubric Line, contains a high-pressure slice containing
455 omphacite-bearing mantle peridotites, which are testimony of a former oceanic tract
456 SE of the Sesia Zone that underwent high-pressure metamorphism (Pognante 1989,
457 Barnes et al. 2014). Secondly, the non-metamorphic Canavese Zone, immediately
458 south of the Insubric Line and north of the Ivrea Zone contains, amongst other
459 lithologies, outcrops of exhumed mantle rocks overlain by Jurassic age radiolarites
460 (Beltrando et al. 2015a), which points to the existence of an oceanic tract between the
461 largely continental Sesia high-pressure unit and the Ivrea Zone. It is only the
462 Canavese Zone, together with the Ivrea Zone and the Southern Alps that form the
463 upper plate of the Alpine orogen in the internal parts of this transect. As demonstrated
464 in the pioneering work by Elter et al. (1966) the Canavese Zone has to be considered
465 the root of the highest nappe exposed in the Préalpes Romandes, the ophiolite-bearing
466 nappe des Gets (see profile of Fig. 2d; unit labeled as Piedmont-Liguria Upper Plate).
467 Both these units represent partly oceanic series that once formed the presently largely
468 eroded upper plate of the Alpine edifice.
469 In contrast to the more southerly-located profiles, very thin slices of the Ivrea
470 mantle wedge, accompanied by lower crustal lithologies of the Ivrea Zone, reach the
471 surface of the earth, slices that are too thin to have a notable effect on the velocity
472 contours. The main part of the Ivrea mantle wedge is confined to a depth >10km as
19 473 indicated by the 7.0 km s−1 contour line. Along this profile the 7.25 km s−1 contour
474 line reveals an inverted velocity gradient, allowing for localizing the base of the Ivrea
475 mantle wedge above the subduction channel.
476 The Adria lower and upper crustal units are affected by backthrusts
477 representing the easternmost extension of the Southern Alps. Details are drawn
478 according to the interpretation given by Roure et al. (1990). This thrusting, detected
479 in the subsurface of the Po plain in the course of the ECORS-CROP reflection
480 seismic campaign, is of pre-Burdigalian (pre-20.5 Ma) age and affects part of the
481 Gonfolite Lombarda sediments, providing an age range for thrusting of Late
482 Oligocene to Early Miocene. Note that this profile does not reach the front of the N-
483 directed Apenninic thrusts, which stays south of the SE end of this section.
484
485 4. NFP-20 West profile (Fig. 2d)
486 This transect follows the trace of a Swiss geophysical campaign (NFP 20 “Deep
487 Structure of Switzerland” whose results were published in Pfiffner et al. (1997).
488 Earlier geological interpretation of the deep structure revealed by controlled source
489 reflection and refraction seismology such as those by Escher et al. (1997) and Schmid
490 et al. (2004) have been revised in the light of the 3D velocity model of Diehl et al.
491 (2009).
492 Along this profile the effects of the foreland propagation of Alpine shortening
493 induced the decollement of the entire Molasse basin along Triassic evaporite horizons,
494 which led to the formation of the Jura Fold-and-Thrust Belt. The Alpine thrust front
495 jumped by some 100km to the north after about 12 Ma ago; this decollement is
496 associated with some 26km of shortening and roots in the external massifs (Aiguille
497 Rouge and M. Blanc massifs) of the Alps (Burkhard & Sommaruga 1998). However,
20 498 the total amount of shortening affecting the upper European crust, which along this
499 transect also includes the lowermost Penninic nappes behind the Penninic Front,
500 amounts to some 50km according to an area balance along this profile, again
501 assuming decoupling at the interface between upper and lower crust. This is almost
502 twice as much as deduced for the ECORS-CROP profile but, in this case, also
503 includes all the shortening that occurred since 35 Ma ago. The deep structure of the
504 External Alps along this profile is characterized very marked undulations of the
505 velocity contours between 7.5 and 6.0 km s−1. Some of these undulations may again
506 have been caused by velocity gradients related to pre-Alpine structures. However,
507 one of these undulations, an antiformal structure located underneath the external
508 massifs, appears to also affect the Moho; the closely spaced 7.5, 7.0 and 6.5 km s−1
509 velocity contours all consistently indicate an antiformal structure affecting the Moho
510 and lower crust. A local up doming of the Moho by about 5km correlates well with
511 earlier observations in reflectivity pattern along the NFP20West near-vertical
512 reflection profile (Pfiffner et al. 1997) that at the time were not further interpreted. As
513 this Moho geometry has now been consistently observed by two independent seismic
514 data sets, we assume this antiformal Moho-structure to be real and possibly
515 representing a pre-Alpine feature.
516 For constructing the near-surface structures in the most external parts of the
517 profile we mainly used Sommaruga (1997) for the Jura Mountains and Sommaruga et
518 al. (2012) for the Molasse basin. In the case of the External Massifs, the Helvetic
519 nappes and the Préalpes Romandes we based our interpretation of data provided by
520 Burkhard (1988), Marchant (1993), Steck et al. (1997), Escher et al. (1997) and
521 Burkhard & Sommaruga (1998).
522 The Penninic Front along this profile coincides with the south dipping Rhone-
21 523 Simplon line, a wide dextral shear zone exhibiting initially ductile deformation
524 starting to be active at around 35 Ma (Steck & Hunziker 1994; Schmid & Kissling
525 2000), which graded into brittle deformation from the Miocene onwards (Cardello &
526 Mancktelow 2015). The Simplon normal fault located immediately east of this
527 transect (Mancktelow 1992; Grasemann & Mancktelow 1993) represents a tensile
528 bridge that accommodates orogen-parallel extension during the backthrusting over
529 the Ivrea mantle wedge towards ESE and associated top WNW thrusting at the
530 Penninic Front of the Western Alps (Keller et al. 2006). The downward projection of
531 the Rhone-Simplon line to some 15km follows that proposed by Escher et al. (1997).
532 At a >15km depth any interpretation of the geometry of this fault and adjacent units
533 is only loosely constrained by migrated reflection seismic data (Marchant 1993) and
534 strike-parallel lateral projection of near surface structures. Our interpretation differs
535 from that of Escher et al. (1997) at a depth > 35km because we conceive the Rhone-
536 Simplon line as delimiting the northern border that acted as a lateral ramp for that
537 part of the Alpine upper crust that moved towards WNW together with the Ivrea
538 mantle wedge during its indentation.
539 The internal units of the Alps SE of the Rhone-Simplon line are characterized
540 by spectacular backfolds within the Briançonnais units whose geometry down to a
541 depth of about 20km is constrained by reflection seismology (Marchant 1993; Steck
542 et al. 1997) and orogen-parallel axial projections of surface structures (Escher et al.
543 1997; Keller & Schmid 2001; Kramer 2002). At a depth >20km the P-wave velocity
544 contours of Diehl et al. (2009) served as a guideline for constraining the geometry.
545 The 6.0 km s−1contour indicates a large-scale velocity inversion that follows the SE
546 limb of the Monte Rosa backfold down to a depth of some 30km. There the 6.0 km
547 s−1 contour swings back into a region characterized by a normal velocity gradient and
22 548 runs parallel to the 6.5 and 7.0 km s−1 velocity contours that are typical for the
549 European lower crust. As a consequence we limit the area characterized by such
550 backfolding to the uppermost 35km of the crust.
551 Farther to the SE the 6.0 and 6.5 km s−1 contours rise to a shallow depth of only
552 some 10 km and 15km, respectively. This rise occurs within the Sesia Unit that is
553 known to exhibit eclogite facies conditions in its most internal parts only (Bousquet
554 et al. 2012a). Because this eclogite facies part of the Sesia Zone is immediately
555 adjacent to the Ivrea mantle rocks, the limit between the Adria mantle wedge and the
556 under-plated Sesia zone is ill constrained. The near-surface structures of the Southern
557 Alps are drawn after Cassano et al. (1986).
558
559 5. Ticino profile (Fig. 2e)
560 The Ticino profile runs parallel to the reflection seismic lines C1, C2 and C3 of
561 NFP 20 project (Pfiffner & Heitzmann 1997); the southernmost parts of this profile
562 are located near lines S4 and S5 of the same project (Schumacher 1997). At depth,
563 the profile crosses the NE-most part of the Ivrea mantle wedge, very near its
564 termination (see Fig. 3). Note also that this profile crosses a pronounced along-strike
565 axial culmination (the core of the Lepontine dome; Berger et al. 2005; Berger &
566 Mercolli 2006; Steck et al. 2013) at a location where the distance between external
567 massifs and Periadriatic line is near a minimum (Fig. 1), going hand in hand with a
568 maximum amount of shortening and exhumation. Since this transect is located east of
569 the Rhone-Simplon line, dextral transpression allowing for top WNW thrusting along
570 the Penninic Front is taken up by the Periadriatic line only.
571 According to an area balance, again assuming decoupling at the interface
572 between upper and lower crust, the total amount of shortening affecting the upper
23 573 European crust (including the Penninic nappes below the Valaisan and Briançonnais
574 units) since around 35 Ma increases by as much as some 115 km when compared to
575 the 50km determined along the NFP20-West traverse. Rosenberg & Kissling (2013)
576 measured some 71 km shortening along a profile located between our NFP20-West
577 and Ticino profiles (their Simplon profile). This indicates a steady increase in the
578 amount of post-collisional (post 35 Ma) shortening within the accreted European
579 lower plate towards northeast into the Ticino transect.
580 In the external part of the profile the 6.0 km s−1 contour rises to shallow depth
581 underneath the frontal part of the external massifs, hinting at the existence of a pre-
582 Alpine structure (External Massifs Shear Zone of Guillot & Ménot 2009?) also along
583 this profile. The near-surface features in the foreland of the Molasse Basin are drawn
584 according to Sommaruga et al. (2012), the geometry of the external Aar and Gastern
585 massifs is modified after Müller (1938) and Pfiffner (2011), the structures in the
586 Helvetic nappes, overlying Penninic nappes and the Gotthard nappe (formerly
587 considered a “massif”) are drawn after Menkveld (1995) and Pfiffner et al. (2010).
588 The internal parts of the transect, depicting the nappe pile within and above the
589 Lepontine dome, are characterized by a large-scale backfold associated with the
590 Northern Steep Belt located within the southern border of the Subpenninic Gotthard
591 nappe (Schmid et al. 2004; Berger et al. 2005), separated by a flat-lying nappe pile
592 from the Southern Steep Belt as defined by Milnes (1974), who was first in realizing
593 the importance of post-nappe recumbent mega-folds affecting a previously existing
594 nappe stack consisting of Briançonnais, Penninic and Subpenninic units. The
595 formation of the Southern Steep Belt, the southern limb of a second more southerly
596 located backfold, is closely associated with backthrusting of the Alpine nappe stack
597 over the Ivrea mantle wedge and overlying crust in the area of the Southern Alps and
24 598 along the Periadriatic Line that initiated between 35 and 30 Ma in a dextrally
599 transpressive scenario (Schmid et al.1987 & 1989; Steck & Hunziker 1994). Note
600 that Miocene thrusting in the Southern Alps is younger, i.e. post-20Ma (Schönborn
601 1992). Based on the offset of the Alpine grade of metamorphism across the
602 Periadriatic Line, the vertical component of dextral transpression amounts to some
603 20km along the trace of this profile. The amount of the simultaneously occurring
604 horizontal displacement across the transpressive Periadriatic Line remains unknown
605 due to the lack of markers; estimates vary between 100km (Schmid & Kissling 2000)
606 and 240km (Handy et al. 2010). Details concerning the shallower parts of this central
607 portion of the profile are drawn based on structural information by Milnes (1974),
608 Huber et al. (1989), Leu (1987), Berger et al. (2005), Steck et al. (2013) and own
609 observations. Reflection seismic data along line C1 and C2 (Pfiffner & Heitzmann
610 1997) were used for placing nappe contacts at the base of the Antigorio nappe and the
611 Gotthard nappe, respectively, and for estimating the depth of the top of the European
612 lower crust.
613 At a depth >20km the NW edge of the Ivrea mantle wedge is defined by the 7.0
614 km s−1 contour. At a depth of <20km the Periadriatic fault, whose dip is evaluated on
615 the basis of surface observations and a series of reflection seismic profiles (Schmid &
616 Kissling 2000; their Fig. 5) delimits the NW edge of the Ivrea mantle wedge,
617 including the Ivrea Zone that represents the adjacent Adriatic lower crust. The Ivrea
618 Zone, including some ultramafic bodies at its northern rim, has been rotated along a
619 near-horizontal axis into a vertical orientation due to Alpine deformation under low-
620 grade metamorphic conditions (Handy 1987; Schmid et al. 1987; Zingg et al. 1990).
621 The Ivrea Zone is well known for representing a map-view section across lower
622 continental crust (Fountain 1976). It is likely that the topmost parts of the Ivrea Zone
25 623 and mantle became highly tectonized under semi-brittle conditions, as is
624 schematically shown in Fig. 2e, allowing for this rotation.
625 The 6.5 km s−1 contour rises to a shallow depth of around 10 km above the
626 Ivrea mantle wedge also in this section, crossing from the European lower crust
627 across the subduction channel over the presumed top of the Ivrea mantle wedge. The
628 northern limit of the Ivrea mantle wedge and the adjacent Ivrea Zone are constrained
629 by downward extrapolation of the Periadriatic Line to a depth of 20km; at greater
630 depth the Periadriatic line delimits the northern limit of the Ivrea mantle wedge.
631 The profile across the upper crust of the Southern Alps is partly based on the
632 geological interpretation of reflection seismic data gathered during the NFP-20 and
633 CROP campaigns (Schumacher 1997; Schumacher et al. 1997; lines C3-south, S4, S5,
634 S6, S7& CROP 88-1) and partly on the subsurface information provided by Cassano
635 et al. (1986). Furthermore, we based our interpretation also on the work of Bertotti
636 (1991) and Schönborn (1992). The total amount of shortening within the Southalpine
637 retro-belt amounts to 74km according to section balancing in the area crossed by this
638 transect (Schumacher et al. 1997). A large, although not well constrained part
639 (between 24km and 43km along a section east of Lake Como according to Schönborn
640 1992) of this shortening pre-dates the Late Eocene. This leaves between 31 and 50
641 km of shortening that occurred in the Middle to Late Miocene. The lower bound of
642 this shortening estimate is more realistic in view of the overall geometry of the
643 transect depicted in Fig. 2e. The reason for the upward excursion of the 6.5 km s−1
644 contour inside the Adriatic upper crust remains unknown; possibly it is related to
645 heterogeneities formed during Variscan tectonics.
646
647 6. Horizontal section (Fig. 3)
26 648
649 To discuss the overall shape of the Ivrea mantle wedge and lateral variations in
650 its geometry the five profiles described above are complemented with a horizontal
651 section across the 3-D P-wave tomographic model of Diehl et al. (2009) at a depth of
652 20km (Figure 3). This horizontal section shows (1) a curved area of positive P-wave
653 anomalies >10% nicely mimicking the well-known positive Bouguer anomaly
654 associated with the Ivrea mantle wedge (Kissling 1984; Serva 2005), (2) an extension
655 of the Ivrea mantle wedge across the Periadriatic Fault visible north of Locarno (in
656 reality the wedge underlies the N-dipping Periadriatic Fault, a backthrust in its
657 hangingwall see Fig. 2e), and, (3) a possible dextral offset of the central axis of the
658 Ivrea anomaly between profiles ECORS-CROP and NFP-20 West that leaves no trace
659 at the surface of the earth when consulting the geological maps of the area.
660 Interestingly, the central axis of the high-velocity area is located east of the
661 Periadriatic Fault in the north and gradually changes over to lying west of the
662 Periadriatic line in the south. This indicates increasing amounts of underthrusting of
663 the mantle wedge beneath the Internal Western Alps towards south, culminating in
664 what is seen in the Argentera section (Fig. 2a): a Periadriatic Fault that acted as a
665 huge backthrust and severe modification of the suture between Europe and Adria
666 (lower plate Piedmont-Liguria ophiolites) in the form of a huge backfold with the
667 high-pressure Briançonnais Units in its core (Fig. 2a).
668
669 C. Multi-stage evolution of the arc of the Western Alps
670 The following discussion on the evolution of the arc of the Western Alps
671 integrates the findings described above with a review of existing knowledge on the
672 multi-stage evolution of the Western Alps concerning kinematic evolution and time
27 673 constraints derived from field data. In view of the considerable difficulties in exactly
674 dating metamorphic events using isotopic ages, particularly in the case of HP
675 metamorphism (see review by Berger & Bousquet (2008), we also heavily rely on
676 biostratigraphical constraints. We also present tectonic reconstructions using Gplates
677 reconstruction software (Boyden et al. 2011) that have already been applied to parts
678 of the Western Mediterranean and eastern realm adjacent of the Alps-Apennines
679 orogenic system (van Hinsbergen & Schmid 2012; van Hinsbergen et al 2014a;
680 Advokaat et al. 2014).
681
682 1. Exhumation of Ivrea Zone and underlying lithospheric mantle in the
683 Mesozoic
684
685 The Ivrea Zone is well known for exposing a map-view section across lower
686 continental crust of the Southern Alps that formerly was part of the Adriatic
687 microcontinent (e.g. Fountain 1976, Handy 1987, Handy & Zingg 1991). It includes a
688 series of small peridotite lenses at its northern rim, due to later tilting in the context of
689 Alpine orogeny. Restoration of the pre-Alpine setting via backrotation of Ivrea Zone
690 and adjacent parts of the Southern Alps adjacent to the Periadriatic Line reveals that
691 the Ivrea Zone was part of the distal passive continental margin that formed in
692 connection with Tethyan rifting culminating at around 180-170 Ma ago (e.g. Bertotti
693 et al. 1993). The Ivrea Zone and underlying Ivrea mantle wedge were located near the
694 western margin of the distal continental margin of the Adria microplate that had a
695 strike (Weissert & Bernoulli 1985) almost parallel to the present-day NNE-SSW
696 strike of the Ivrea mantle wedge (Fig. 3). Based on a combination of P-T estimates of
697 pre-Alpine parageneses and radiometric dating it was inferred that crustal thickness
28 698 was locally reduced to 10km or less in the area of the Ivrea Zone and adjacent upper
699 crustal units (Handy et al. 1999). Hence the present-day slope of the Moho, rising
700 from some 40km to about 20km revealed by the Moho-map of Spada et al. (2013),
701 and, rising to some 15km in Figs. 2c&d, is interpreted as inherited from passive
702 margin formation in Mesozoic times. The Canavese area, coinciding with the ocean-
703 continent boundary, from which the Sesia extensional allochthon separated during
704 rifting (Froitzheim et al. 1996; Beltrando et al. 2015a,b), is only some 50km away
705 from the Ivrea Zone.
706 The shallow depth of the top of the Ivrea mantle wedge inherited from
707 Mesozoic rifting has important consequences regarding rheology and strength of the
708 Ivrea mantle wedge during Alpine orogeny. The brittle to crystal plastic transition of
709 peridotite is estimated to be at around 600° according to most workers (e.g.
710 Druiventak et al. 2011). Temperatures in the depth range of the Ivrea mantle wedge
711 (15-40km) are between 400° and 650°C in the central area of the Alpine orogen
712 according to the steady state kinematic model of Bousquet et al. (1997). Hence, most
713 of the Ivrea mantle wedge is expected to be controlled by frictional strength around
714 1000 MPa at 15 km depth in a compressional regime under dry conditions, reaching
715 excessive values at greater depth (Byerlee 1978). At its front the Ivrea mantle wedge
716 is surrounded by upper crustal nappes (Fig. 2) whose rheology is likely to be
717 controlled by viscous creep of large quartz-bearing volumes of rock. Extrapolation of
718 flow stress data to geological strain rates, combined with paleopiezometry using
719 recrystallized grain size, predict flow stresses between of 10-100 MPa within the
720 same 400° and 650°C temperature range (Stipp et al. 2002). This indicates that the
721 strength of the Ivrea mantle wedge exceeds that of quartz-bearing units at its front by
722 at least one order of magnitude. This suggests that indentation of the Ivrea mantle
29 723 wedge into the Alpine nappe stack probably played an important role in the formation
724 of the arc of the Western Alps, as will be further discussed below.
725
726 2. Top-NNW thrusting during sinistral transpression (Stage 1, pre-35 Ma)
727
728 The orientation of stretching lineations in combination with associated sense of
729 shear criteria have been applied to deduce tectonic transport directions in highly
730 sheared rocks from the Alps (e.g. Malavieille et al. 1984; Platt et al. 1989). Along the
731 arc of the Western Alps the orientation of the lineations is dominantly transverse to
732 the strike of the orogen. The inferred tectonic transport smoothly changes from top
733 NW in the area of Monte Rosa to top WSW in the area of the Dora Maira area
734 (Malavieille et al. 1984). This apparently simple arrangement, if taken to indicate
735 sense of movement, leads to obvious difficulties in geometrical and plate kinematic
736 terms (Platt et al. 1989). The major difficulties in interpreting kinematic data based
737 on stretching lineation analysis is that (1) stretching lineations may indicate the
738 orientations of finite strain (e.g. Ramsay 1980) rather than shearing during a
739 particular stage of a complex orogenic evolution and (2) that orientations of older
740 lineations may have been reoriented by younger deformations (e.g. Walcott and
741 White, 1998; Keller & Schmid 2001; van Hinsbergen and Schmid, 2012). Hence,
742 only careful local studies addressing the polyphase structural evolution of particular
743 areas can reveal a more complex kinematic evolution. The results of such studies are
744 compiled in Fig. 4.
745 A first generation of lineations indicating top north transport is best
746 documented in areas of the Western Alps located behind, i.e. east, of the Penninic
747 Front (areas 1, 2 and 3 in Fig. 4 and references therein). In these areas the orientations
30 748 of this first generation of inferred transport directions is related to pre-35Ma nappe
749 stacking. According to detailed analyses of structural superposition this first
750 generation of lineations pre-dates thrusting along the Penninic Front initiating at
751 around 35 Ma and is hence is of Eocene or even earlier age (e.g. Bucher et al. 2003;
752 Ceriani & Schmid 2004; Loprieno et al. 2011). However the orientations of these first
753 generation lineations scatter considerably between top-NW and top-NE due to later
754 re-orientation. In the more internal parts of the internal Western Alps the preservation
755 potential of the orientation of such early stretching lineations is low. In most places
756 they are not discernable any more because of intense shearing during subsequent
757 deformation during stage 2 (see below). In the more internal units of the Aosta
758 Valley, for example, they are seen to scatter too much for inferring transport
759 directions of this earliest phase (Bucher et al 2004). Only in some parts of the
760 eclogitic lower plate Piedmont-Liguria units is a former top NNW eclogite facies
761 lineation preserved (area 5 in Fig. 4; Philippot 1988, 1990). Interestingly, also parts
762 of the External Western Alps (Pelvoux massif and its cover) are known to have been
763 affected by stage 1 to N to NW-directed deformation of pre-Priabonian age (Dumont
764 et al. 2011).
765 Large-scale structural considerations (Schmid & Kissling 2000; Handy et al.
766 2010) and plate kinematic constraints (Fig. 5) indicate that thrusting and plate
767 convergence were top N to NNW before 35 Ma. This implies a scenario of sinistral
768 transpression in the Western Alps during the early stages of orogeny when the upper
769 plate tectonic units below which the Alpine nappe stack formed (Adria lithosphere;
770 oceanic in case of the Western Alps, continental in case of the Eastern Alps) moved
771 towards N to NNW, presumably past that part of the European plate presently located
772 in SE France, as first proposed by Ricou & Siddans (1986). Pre-35Ma plate
31 773 convergence was associated with nappe stacking and early stages of exhumation of
774 parts of nappes buried to high-P (e.g. Schmid et al. 1996, their Fig. 8). A point
775 located at the northern tip of Adria (Monti Lessini) moved some 260km to the NNW
776 in respect to stable Europe between 55 and 35 Ma go (1.3 cm/y) according to the
777 reconstruction shown in Fig. 5d-f, which assumes that Adria was part of Africa at this
778 time, and used the Europe-Africa plate circuit of Seton et al. (2012). This is more
779 than the 195km (0.98 cm/y) of shortening calculated by Schmid & Kissling (2000)
780 for the same time interval but considerably less than the estimate of Handy et al.
781 (2010) who inferred 465km shortening for the interval 35-67Ma (1.45cm/y), both
782 these estimates being solely based on interpretations of the geological structure and
783 metamorphic history of the Alps. According to Handy et al. (2010) top-N thrusting
784 started as early as some 84Ma ago, consistent with the plate circuit estimates for the
785 onset of convergence based on Seton et al. (2012), and first led to high-pressure
786 metamorphism in the Sesia Zone (Manzotti et al. 2014), predating high-pressure
787 metamorphism within the lower plate units derived from the Piedmont-Liguria Ocean
788 that started at around 60Ma ago (Berger & Bousquet 2008).
789 In summary, peak pressures within the Western Alps were reached in a scenario
790 of sinistral transpression along an original approximately N-S striking segment of the
791 western Alps (Fig. 5d&e) during this first stage. Most of the exhumation of the
792 eclogites to moderate crustal levels very probably also took place during this first
793 stage of the evolution of the arc of the Western Alps.
794
795 3. Top-WNW thrusting and WNW-directed indentation of the Ivrea mantle slice
796 (Stage 2, 35-25 Ma)
797
32 798 A second population of stretching lineations is associated with deformation that
799 post-dates stage 1 top-NNW thrusting according to structural overprinting criteria.
800 The top-WNW lineations formed during this stage can reliably be attributed directly
801 to thrusting along the Penninic Front in the case of locations 1,3,4 and 7 in Fig. 4
802 (references cited therein). There the inferred thrusting directions vary between
803 azimuth 260° and 345° with a mean at around azimuth 300°. The measurements at
804 locations 2 and 5 were taken in more internal positions with respect to the Penninic
805 Front while those at location 6 are from the footwall of the Penninic Front (lineations
806 in the “schistes à bloc formation” immediately below the base of the Embrunais-
807 Ubaye nappe stack; Trullenque 2005). Given the similar spread in orientations
808 shearing during stage 2 top-WNW thrusting is likely also for locations 2, 5 and 6.
809 Several independent sets of data provide dating constraints. 40Ar-39Ar dating of
810 phengites that formed syn-kinematically with shearing of basement rocks of the
811 Pelvoux massif in shear zones along the Penninic Front or in its immediate footwall
812 vary between 35 and 30 Ma with one younger age at 27Ma (Simon-Labric et al.
813 2009; Dumont et al. 2012). This dating confirms earlier evidence for the onset of top
814 WNW thrusting at around 32 Ma based on fission track dating (Fügenschuh &
815 Schmid 2003). Since the Penninic Front is associated with dextral shearing along the
816 Rhone-Simplon line (Fig. 1) and dextral shearing in the Insubric mylonites (part of
817 PL of Fig. 1 around Locarno) the dating of such dextral shearing provides further
818 timing constraints. The “ductile Simplon shear zone” of Steck et al. (2015), predating
819 late stages of normal faulting at the Simplon normal fault (Mancktelow 1992), was
820 active between 34 and 25 Ma ago as inferred from the ages of the mantle-derived
821 andesite and tonalite (32–29 Ma) and synkinematic crustal aplite and pegmatite
822 intrusions in the southern steep belt between Domodossola and Locarno (Steck et al.
33 823 2013). Dextral shearing in the mylonite belt of the Insubric line, a segment of the
824 Periadriatic Line (PL in Fig.1) initiated at around 32 Ma ago based on radiometric
825 age data from the eastern margin of the Bergell pluton (32-30 Ma; von Blanckenburg,
826 1992). The deep-seated Bergell intrusion including its subsequent fast exhumation is
827 closely related to a combination of backthrusting and dextral strike-slip movements
828 accommodated by the greenschist mylonites of the Insubric line (Schmid et al., 1989;
829 Berger et al., 1996). The top-WNW senses of shear in the “schistes à bloc” formation
830 (Kerckhove 1969) at location 6 of Fig. 4, affected by top-WNW shearing during a
831 first phase of deformation (Trullenque 2005), provides a stratigraphic age constraint;
832 this chaotic mélange overlies the Priabonian-age flysch de Champsaur (Bürgisser &
833 Ford 1998), and hence, its age of formation is considered to be latest Priabonian to
834 earliest Oligocene (around 34 Ma). A similar age constraint is provided for the
835 sinistral senses of shear observed at the western termination of the Argentera-Cuneo
836 Line at location 7 of Fig. 4 that is kinematically connected with the Penninic Front
837 (Trullenque 2005); these mylonites are found in immediate tectonic contact with the
838 grès d’Annot of Late Eocene to Early Oligocene age that terminate the evolution of
839 the Eocene Northern Alpine foreland basin in the southern Subalpine chains (Ford &
840 Lickorish 2004).
841 The end of stage 2 shearing is not equally well constrained but certainly pre-
842 dates the activity of the left-lateral Giudicarie transpressive belt starting to offset the
843 Periadriatic line at around 23-21 Ma ago (Scharf et al. 2013). An end of stage 2 at
844 around 25 Ma ago is compatible with the onset of the Paleo-Apenninic phase
845 (Schumacher & Laubscher 1996) at about this time. Stage 2 certainly pre-dates the
846 onset of rapid counterclockwise rotation of the Corsica-Sardinia block at 21 Ma ago
847 (Gattacecca et al. 2007) that we link to stage 3 oroclinal bending discussed in the next
34 848 chapter.
849 At the large scale thrusting at the Penninic Front is interpreted to be
850 kinematically connected to dextral strike slip along the Rhone Simplon and
851 Periadriatic Lines on the one hand, and to sinistral strike slip along the Argentera-
852 Cuneo Line on the other hand. The orogen-perpendicular lineations in the Western
853 Alps compiled by Malavieille et al. (1984) are probably related to stage 2 top-WNW
854 shearing and associated backfolding that is pervasive within entire volume of rocks
855 between Penninic Front and the Periadriatic Line between Ivrea and Cuneo; this
856 shearing formed new lineations and/or reoriented older ones (Fig. 1). It is proposed
857 that this pervasive shearing, linked to WNW-directed indentation of the Ivrea mantle
858 wedge, first proposed by Laubscher (1971), also took place during stage 2 (35 and 25
859 Ma). The amount of top WNW displacement along the Penninic Front is not well
860 constrained. Schmid & Kissling (2000) estimated WNW-directed displacement of the
861 Ivrea Zone to amount to around 100km, (Handy et al. 2010) proposed some 200km.
862 The reconstruction in Fig. 5c yields 93km and 150km for the SW and NE ends of the
863 Ivrea mantle wedge, respectively.
864 The reconstruction shown in Fig. 5c depicts displacement vectors that diverge
865 between Western Alps and Eastern Alps and hence are bound to induce orogen-
866 parallel extension in the Alps that is known to also have started at around 35 Ma ago
867 (Pleuger et al. 2008; Beltrando et al. 2010). The backfolding and backthrusting,
868 which is so typical for the Western Alps (see profiles of Fig. 2 and accompanying
869 text) and which affected a pre-existing nappe pile, also occurred during stage 2 (e.g.
870 Bucher et al. 2003). Hence, it is also associated with the indentation of the Ivrea
871 mantle wedge that underthrusted much of the internal zones of the Alps (see Fig 2a-
872 c). The spectacular backthrust over the Cenozoic cover of the Tertiary Piedmont
35 873 Basin shown at the eastern end the profiles shown in Fig. 2 a & b also formed during
874 stage 2.
875
876 4. Oroclinal bending and top SW thrusting in the southern part of the Western
877 Alps (Stage 3, 25-0 Ma)
878
879 The external Alps south of the Pelvoux massif are well known for late-stage
880 top-SW thrusting (e.g. Fry 1989, Lickorish & Ford 1998, Bürgisser & Ford 1998) that
881 is not seen farther to the north along the Western Alpine arc. By studying
882 overprinting relationships south of the Pelvoux massif (area at location 4 in Fig. 4)
883 Trullenque (2005) discovered that this top-SW thrusting actually post-dates stage 2
884 top WNW thrusting. Lickorish & Ford (1998) have shown that 10.5 km out of a total
885 of around 21 km shortening, associated with minor thrusting rooting in the basement
886 of the Argentera Massif, occurred in Late Miocene to Pliocene times. Ritz (1992),
887 based on a paleostress analysis in the Digne, Castellane and Nice thrust systems,
888 found that during Middle Miocene to recent times the σ1 trajectories varied from NE-
889 SW in the Digne area, gradually turning into a N-S orientation in the Castellane arc
890 and eventually to a NW-SE direction in the case of the Nice thrust sheet. This same
891 radial arrangement was found for the p-axes derived from stress inversion of
892 earthquake focal mechanism data by Delacou et al. (2004, their Fig. 5). This radial
893 pattern is very likely related to rotation around a pivot located east of the Argentera
894 massif in the area of the “Ligurian knot” (Laubscher et al. 1992), i.e. in an area where
895 the 3 Moho’s beneath Alps, Apennines and northern Tyrrhenian Sea meet (Spada et
896 al. 2013).
897 Three largely independent lines of evidence support the idea that oroclinal
36 898 bending in the southernmost Western Alps is to be seen in connection with the
899 structural observations mentioned above. Firstly, Collombet et al. (2002) reported
900 counterclockwise rotations about a subvertical axis in respect to stable Europe that
901 increase from 68° in the Ubaye region (area between profiles 1 and 2 of Figure 1) to
902 117° in the Ligurian Alps SE of Cuneo (Fig. 1). The data were collected in
903 sedimentary rocks of the Briançonnais tectonic unit. The age constrains for
904 magnetization are rather loose; the authors propose that the rocks were remagnetized
905 after the backthrusting phase (our stage 2) and therefore to be of Late Oligocene age.
906 Secondly, all the major tectonic elements of the internal Alps derived from the
907 Briançonnais and Piedmont-Liguria paleogeographical domains can be followed into
908 western Liguria and all the way to Genova and slightly beyond (Fig. 1). In the
909 Ligurian Alps all the pre-Oligocene Alpine structures, sealed by the Tertiary
910 Piedmont Basin, are top-S in present-day coordinates. Since the counterclockwise
911 rotation of stable Adria only amounts to 10±10° (van Hinsbergen et al., 2014b) or 20°
912 (Marton et al. 2011) at least a part of the 49° differential rotation of Liguria in respect
913 to the Ubaye region (117°-68°= 49°), namely 39±10°, is probably related to oroclinal
914 bending during stage 3. It is likely that oroclinal bending during stage 3 was partially
915 driven by eastward trench retreat of the Apennines slab. However, the rotation in the
916 Ubaye region itself, located north of the Argentera-Cuneo Line is probably also due
917 to stage 2 indentation of Adria. Thirdly, a 50° counterclockwise rotation phase of the
918 Corsica-Sardinia block occurred between 20.5 and 16 Ma ago (Gattacceca et al.
919 2007) around a pivot near the N tip of Corsica, as Corsica and Sardinia drifted away
920 from their pre-Miocene location south of southern France (Advokaat et al. 2014); this
921 virtually demands substantial counterclockwise rotation in Liguria associated with
922 oroclinal bending. Note that such counterclockwise rotation is directly demonstrated
37 923 by the ~45-50° anticlockwise rotation of mostly Oligocene sediments of the Tertiary
924 Piedmont Basin sealing pre-Oligocene Alpine structures evidenced by paleomagnetic
925 data (Maffione et al. 2008).
926 Oroclinal bending in the southernmost Western Alps during stage 3 is evidently
927 associated with orogeny in the Apennines. Slab rollback of the Adria plate towards N
928 in the northernmost Apennines is the underlying driving force for both oroclinal
929 bending and the rotation of Corsica and Sardinia. The discussion of the Alps-
930 Apennines transition, discussed in the next chapter, will shed more light to the
931 Miocene history of the Ligurian Alps that cannot be understood without discussing
932 orogeny in the Apennines and its spatial and temporal relation to orogeny in the Alps.
933
934 D. The Alps-Apennines transition in space and time
935 1. Tertiary Piedmont Basin and Epiligurian basins
936
937 Late Eocene to Neogene sedimentary basins found in the interior of the
938 Western Alpine arc, the Tertiary Piedmont Basin, and similar basins found on top of
939 the External Ligurides of the Northern Apennines, referred to as “Epiligurian” basins,
940 play a crucial role for understanding the spatial and temporal transitions between the
941 west- and north-verging Alps and the east- and north-verging Apennines (Fig. 6).
942 The southern limit of the outcropping pre-Pliocene parts of the Tertiary
943 Piedmont Basin (Fig. 1) transgresses the high-pressure lower plate units of the
944 Ligurian Alps (i.e. the high pressure Briançonnais and Piedmont-Liguria units), as
945 well as adjacent upper plate units of the Alps derived from the Piedmont-Liguria
946 Ocean found immediately east of Genova (i.e. the Antola nappe). As first pointed by
947 Elter & Pertusati (1973) the Antola nappe, predominantly consisting of Helminthoid
38 948 Flysch, is best considered as a former part of the Alps; they wrote: “La falda del
949 Monte Antola e le altre formazioni liguri comprese nel triangolo Genova-Val
950 Staffora-Levanto sono da considerarsi appartenenti alle Alpi in quanta sono saldate a
951 queste dai terreni trasgressivi del Bacino terziario Piemontese fin dall'Eocene”).
952 While earlier works suggested that this basin initiated by subsidence in a scenario of
953 crustal stretching (Mutti et al. 1995) there seems to be some consensus in more recent
954 publications that the Oligocene to Lower Miocene (Aquitanian) sediments of this
955 basin were laid down in a compressive to transpressive setting (e.g. Carrapa &
956 Garcia-Castellanos 2005; Mosca et al. 2010).
957 As shown in the profiles of Fig. 2 a & b the westernmost parts of the Tertiary
958 Piedmont Basin were installed onto the upper crust of the Adria plate, which in turn
959 overlies the Ivrea mantle wedge. It formed the foredeep of the backthrusted and -
960 folded Western Alps during the Oligocene, which later became buried under post-
961 Messinian cover. Such syn-sedimentary backthrusting to the east to southeast is seen
962 in the seismic sections of lines 3 and 4 of Mosca et al. (2010, their Fig. 4). Thrusting
963 is sealed by the Burdigalian unconformity that is ubiquitous in seismic sections of the
964 subsurface of the Po plain (e.g. Schumacher & Laubscher 1996 and references
965 therein). Such pre-Burdigalian thrusting is also seen along the front of the Southern
966 Alps in western Lombardy where it affects sediments of Oligocene to Lower
967 Miocene age, the Gonfolite Lombarda Group sensu lato (e.g. Bernoulli et al. 1989,
968 1993) that are part of the Southalpine foredeep (Garzanti & Malusa 2008) and
969 laterally connect with the Tertiary Piedmont Basin (Mosca et al. 2010). Subsidence in
970 the Piemonte region continued in Burdigalian and younger times, when this basin
971 started to become involved in N-directed thrusting in connection with the Apenninic
972 orogeny. The Torino and northern Monferrato hills immediately east of Torino (Fig.
39 973 1) are also considered as a part of the Tertiary Piedmont Basin by most authors
974 although they would better be considered a part of the Epiligurian basins (see below).
975 They morphologically mark the northern front of the Apenninic orogen that formed
976 in Late Miocene to Pliocene times. These two northernmost parts of the Tertiary
977 Piedmont Basin transgressed an “Alpine-type” basement in case of the Torino hills
978 (Piana 2000) and a part of the External Ligurides in the case of the northern
979 Monferrato hills (Laubscher 1971; Mosca et al. 2010; Festa & Codegone 2013).
980 Together with the rest of the Tertiary Piedmont Basin that transgressed Alpine
981 structures they became thrusted northwards as wedge-top basins, much like the so
982 called Epiligurian basins of the External Ligurides, described below, along-strike
983 farther to the east and southeast.
984 The term Epiligurian basins is used for wedge-top basins overlying the External
985 Ligurides of the northern Apennines (Fig. 6) whose age-range and sedimentary
986 characteristics are similar to those of the Tertiary Piedmont Basin (Molli et al. 2010
987 and references therein). The sedimentary sequence starts with the deposition of the
988 deep marine Middle to Upper Eocene Monte Piano marls above a regional
989 unconformity that indicates some pre-Apenninic deformation of parts of the External
990 Ligurides that represent a fossil ocean-continent transition (Marroni et al. 2001)
991 within the Adria plate that was part of the upper plate of the Alpine orogen. The Early
992 Oligocene Ranzano formation was laid down above a second unconformity and
993 received detritus from the metamorphic units of the Central Alps (Lepontine dome)
994 that started to become exhumed north of the Periadriatic line (Garzanti & Malusa
995 2008). In the Early Oligocene the External Ligurides were part of the Adriatic
996 foredeep in respect to the Southern Alps that formed due to “retroshearing”
997 (Beaumont et al 1994) within the Alpine orogen (Schmid et al. 1996). Subsidence and
40 998 siliciclastic sedimentation in the Epiligurian basins continued during the Late
999 Oligocene when these basins and the underlying units became involved in the
1000 Apenninic orogeny during its earliest stages. From ca. 23 Ma onward the Epiligurian
1001 basins evolved into wedge-top basins as the External Ligurides became underplated
1002 by the accretion of Subligurian and Tuscanide units within the evolving Apenninic
1003 orogen (Molli et al. 2010 and references therein). Very recently Piazza et al. (2016)
1004 provided evidence for intense syn-depositional sinistral transpression during
1005 Oligocene to Early Miocene times sealed by a Burdigalian unconformity within one
1006 of the Epiligurian basins. This sinistral transpression is contemporaneous with
1007 sinistral strike slip along the E-W striking Villalvernia-Varzi line (Elter & Pertusati
1008 1973) that defines the boundary between the Antola nappe often considered as a part
1009 of the Alps in the south and the External Ligurides considered as a part of the
1010 Apennines in the north (stippled line in Fig. 1 north of Genova). This sinistral strike
1011 slip movement was active in Late Oligocene to Early Miocene times as well (Di
1012 Giulio & Galabati 1995; Schumacher & Laubscher 1996). Laubscher (1991)
1013 considered the Villalvernia-Varzi line as accommodating the WNW-directed
1014 indentation of the Ivrea mantle wedge previously discussed (stage 2, during the
1015 evolution of the arc of the Western Alps). Although we consider the direct link
1016 between the sinistral Argentera-Cuneo line and the Villalvernia-Varzi line proposed
1017 by Laubscher (1991) as unlikely (see Figs. 1 & 6) sinistral strike slip motion along
1018 the Villalvernia-Varzi line and sinistral transpression in the Epiligurian basins were
1019 probably kinematically connected with stage 2 WNW-directed indentation of the
1020 Ivrea mantle wedge during the Oligocene.
1021 In summary, the Tertiary Piedmont Basin seals very substantial pre-Late
1022 Eocene Alpine structures. The Epiligurian basins are of a similar type and age but
41 1023 they were deposited above an angular unconformity onto a unit that never underwent
1024 substantial Alpine deformation, namely the External Ligurides of the Northern
1025 Apennines (Fig. 6). Later, i.e. from about 23 Ma onwards, the latter represented
1026 wedge-top basins riding on the External Ligurides, which thrusted over more external
1027 units of the Northern Apennines towards NE (Molli 2008; Molli et al. 2010, and
1028 references therein). In essence, the stratigraphic base of these basins post-dates the
1029 pre-Late Eocene top-NNW thrusting during stage 1 sinistral transpression in the
1030 Western Alps described earlier while it pre-dates the onset of top-NE thrusting during
1031 the Apenninic orogeny, which is contemporaneous with stage 3 oroclinal bending and
1032 top SW thrusting in the southern part of the Western Alps. It follows from this that
1033 substantial parts of these basins probably represented syn-orogenic deposits during
1034 stage 2 top-WNW thrusting and indentation of the Ivrea mantle wedge into the
1035 Western Alps at 35-25 Ma ago.
1036
1037 2. The Alps-Apennines transition in map view
1038
1039 An interpretative map sketch of the Alps-Apennines transition is presented in
1040 Fig. 6. Considerable parts of this area are either buried by deposits of the Tertiary
1041 Piedmont Basin and very young Cenozoic cover of the Po plain or by the waters of
1042 the Ligurian and North Tyrrhenian Sea; hence this compilation in Fig. 6 remains
1043 speculative in many respects.
1044 The continuation of the undeformed or little deformed Alpine foreland has to be
1045 located in Variscan Corsica that became thrusted by Alpine allochthonous units until
1046 the Late Eocene (Malavieille et al. 1998; Molli 2008). Note, however, that the
1047 present-day NW-SE strike of the Late Eocene Alpine front across the Ligurian Sea
42 1048 and Corsica restores to a N-S strike after the retro-deformation of 47° post-Eocene
1049 rotation of the Ligurian Alps (Maffione et al. 2008; see Figs. 5d&e) and 50° rotation
1050 of Corsica-Sardinia (Gattacceca et al. 2007). The Alpine allochthons in Corsica
1051 consist of, from bottom to top, the high-pressure units of the Tenda “massif”
1052 representing formerly subducted and subsequently exhumed parts of Variscan
1053 Corsica that remains devoid of an Alpine overprinted in the west, the high-pressure
1054 units of the Piedmont-Liguria Ocean and finally overlying non-metamorphic
1055 ophiolitic and continental nappes (Balagne and Nebbio nappes). The two lower units
1056 clearly represent lower plate units of the Alps that have to be correlated with the
1057 Briançonnais and Piedmont-Liguria units of the Western Alps, respectively. Although
1058 the origin of the uppermost non-metamorphic units is controversial (see Marroni &
1059 Pandolfi 2003 and references therein) we follow Malavieille et al. (1998) who argue
1060 for an origin of these highest units from the ocean-continent transition at the eastern
1061 margin of the Piedmont-Liguria Ocean. This makes them comparable to the non-
1062 metamorphic upper plate units of the Western Alps mapped in Figs. 1 & 6 (Préalpes
1063 Romandes, Embrunais-Ubaye and Western Liguria Helminthoid flysch). Oligocene-
1064 Miocene extension and transtension localized in the Gulf of Lion and the Ligurian
1065 Sea strongly affected the entire Corsican belt (Fig. 6), reducing its crustal thickness to
1066 normal thickness. In map view the width of the Alpine belt in Corsica is drastically
1067 reduced in comparison with the Western Alpine orogenic belt. We interpret this
1068 reduction in width to be partly related to the switch in subduction polarity between
1069 Alps and Apennines that led to a new west-dipping “Apenninic” subduction zone that
1070 must have evolved offshore eastern Corsica since latest Eocene or earliest Oligocene
1071 times (Fig. 6; Doglioni et al. 1998; Carminati et al. 2004; Molli & Malavieille 2011).
1072 This new plate boundary was located east of the Corsica basin (Mauffret et al.
43 1073 1999; see also CROP-03 and M-12A seismic sections in Finetti et al. 2001 and Finetti
1074 2005). It evolved after cessation of Alpine, southeastward subduction in the
1075 Oligocene, and was followed by the two-stage opening of the Gulf of Lion back-arc
1076 basin. In a first, Oligocene stage, at around 30 Ma Corsica-Sardinia rifted off the
1077 Provence margin of southern France (Séranne, 1999). This was followed by the 21-16
1078 Ma Corsica-Sardinia rotation around a pole located somewhere between the northern
1079 tip of Corsica and the Ligurian coast (Advokaat et al. 2014). The new plate boundary
1080 probably had a ~N-S-direction, slightly oblique to the NNW-SSE present-day strike
1081 of the Alpine orogen in Corsica (Fig. 6). In respect to the Apenninic orogeny both
1082 lower and upper plate units of Alpine Corsica occupied an upper plate position,
1083 remaining largely unaffected by Apenninic shortening, while easterly and northerly
1084 adjacent parts of the former Alps were heavily reworked when backthrusted onto the
1085 Adria continental margin and subsequently stretched as the Adria plate below the
1086 North Apennines rapidly rolled back eastward (Grandjacquet & Haccard 1977;
1087 Doglioni et al 1998; Rosenbaum & Lister 2004; Spakman & Wortel 2004; Marroni et
1088 al. 2010).
1089 The northern margin of the lower plate metamorphic core of the Ligurian Alps
1090 (AFSZ in Fig. 1) is not exposed being buried beneath the Cenozoic cover of the
1091 Tertiary Piedmont Basin and younger cover (Fig.1). In Fig. 6 it is mapped according
1092 to subsurface seismic data (Mosca et al. 2010). Along a transect located west of
1093 Genova (profile A in Mosca et al. 2010) the metamorphic units of the Ligurian Alps
1094 were thrusted northward by about 50 km over the Adria continental autochthonous
1095 during the Apenninic orogeny in Miocene-Pliocene times. This implies that the
1096 western continuation of the older (pre-25Ma) Argentera-Cuneo Line, delimiting the
1097 southern boundary of the indenting Ivrea mantle wedge and overlying internal
44 1098 Western nappe stack during the stage 2 top-WNW thrusting in the arc of the Western
1099 Alps, must be cut off by a N-S striking sinistral shear zone (Figs. 1 & 6). This
1100 putative shear zone, which is only schematically drawn in Fig. 6, delimits the western
1101 edge of the Ligurian Alps affected by these 50 km of Apenninic top N thrusting from
1102 a narrow strip of the Southern Alps retro-belt in the west.
1103 Eastward, i.e. north of Genova, the northern limit of the metamorphic core of
1104 the Ligurian Alps bends into a N-S strike and is exposed at the surface forming the
1105 well-known Sestri-Voltaggio line (or zone; Cortesogno & Haccard 1984) that is often
1106 taken as representing the Alps-Apennines junction (e.g. Castellarin 2001). Although
1107 reworked in later Neogene times (Crispini et al. 2009) this line basically represented
1108 a Late Paleocene to Early Eocene normal fault (Hoogerduijn Strating 1994) sealed by
1109 the deposits of the Tertiary Piedmont Basin. This normal fault juxtaposes the high-
1110 pressure rocks of the Voltri Group in the west (1.5-2.5 GPa), attributed to the Alpine
1111 lower plate, with low metamorphic grade units to the east (Sestri-Voltaggio Zone,
1112 various Ligurian flysch units and finally the Antola nappe; see Fig. 8d in Molli et al.
1113 2010), whose grade of metamorphism progressively decreases eastward and up-
1114 section (see Fig. 2 in Capponi et al. 2009). It finally reaches the diagenetic zone in
1115 terms of illite crystallinity (Ellero et al. 2001) over a short horizontal distance of only
1116 some 5-7 km. Following Elter & Pertusati (1973) who considered the Antola nappe
1117 as an integral part of the former Alps that became backthrusted during Alpine
1118 orogeny, we attribute the units east of the Sestri-Voltaggio normal fault to the Alpine
1119 upper plate (Fig. 6).
1120 The unknown Alpine-type basement of the Torino hills buried under the
1121 deposits of the Tertiary Piedmont Basin and the Antola nappe were mapped together
1122 as parts of a strip of former Alpine upper plate units mostly buried under Cenozoic
45 1123 deposits in Fig. 6. During the Apenninic orogeny, together with the metamorphic core
1124 of the Ligurian Alps, they thrusted the External Liguride units of the Monferrato hills
1125 along the Rio Freddo Fault Zone (Piana 2000) striking NW-SE between the Torino
1126 and Monferrato hills. The Rio Freddo Fault Zone links up with the E-W-striking
1127 Villalvernia-Varzi line (Fig. 1), a Late Oligocene to Early Miocene thrust, active
1128 during the deposition of a part of the Tertiary Piedmont basin fill, with a strong
1129 sinistral strike slip component (Laubscher et al. 1992; Schumacher & Laubscher
1130 1996). This sinistral transpressional fault is responsible for laterally transporting the
1131 External Ligurides exposed in the Monferrato hills in front of the Antola nappe and
1132 the Torino hills Alpine basement. At its eastern termination the Villalvernia-Varzi
1133 line bends into a N-S-strike as it transforms into an Apenninic top NNE thrust of the
1134 Antola nappe over the external Ligurides in the north and over the Internal Ligurides
1135 further south (Fig. 6). The polyphase history of the Antola nappe, with early and
1136 Alpine top-NW thrusting and folding, followed by top-NNE Apenninic thrusting over
1137 the Ligurides, is well documented (e.g. Marroni et al. 1999; 2002a; Levi et al. 2006).
1138 Offshore the continuation of this thrust is only tentatively drawn in Fig. 6; southward
1139 it is interpreted to eventually abut the northern continuation of the west-dipping
1140 upper/lower plate boundary of the Apennines orogen.
1141 The next eastward unit of the Apennines, the Internal Ligurides, has also to be
1142 considered as having represented an integral part of the former Alps. It occupied an
1143 upper plate position before is became backthrusted to the NNE during the Apenninic
1144 orogeny. The internal Ligurides are characterized by the presence of classical
1145 ophiolites and an Upper Jurassic to Lower Cretaceous sedimentary cover (cherts,
1146 Calpionella limestone and Palombini shales) associated with Upper Cretaceous–
1147 Paleocene turbiditic sequences (Marroni & Pandolfi 1996) that represent the pelagic
46 1148 cover of the ophiolitic basement of the Piedmont-Liguria ocean (Decandia & Elter
1149 1972) whose stratigraphy is virtually indistinguishable from that of the ophiolites of
1150 the Alps sensu stricto (Bernoulli et al. 2003). Moreover, since the pioneering work of
1151 Grandjacquet & Haccard (1977) several authors evidenced classical structural
1152 overprinting patterns at all scales, clearly showing that the early tectonic history of
1153 the Internal Ligurides was characterized by a top W to NW Alpine tectonic
1154 fingerprint, associated with low-grade metamorphism reaching some 300°C
1155 (Pertusati & Horrenberger 1975; Van Wamel 1987; Marroni & Pandolfi 1996; Ellero
1156 et al. 2001).
1157 The External Ligurides are far-travelled and detached units that are confined to
1158 the northernmost Apennines between Torino and Bologna (Fig. 6). They overlie thin
1159 slices of Subligurian units and the Cervarola unit representing flysch deposits of the
1160 Adriatic foredeep outcropping in two small windows (Molli et al. 2010). The
1161 External Ligurides dominantly consist of detached Cretaceous-Paleocene flysch
1162 sequences (Helminthoid flysch) overlying basal mélange-type complexes (Marroni et
1163 al. 1998, 2001, 2002b). Two main subgroups of units can be recognized: those
1164 associated with ophiolites and with ophiolite derived debris (Molli 1996), and others
1165 without ophiolites and associated with fragments of Mesozoic sedimentary sequences
1166 and conglomerates with continental Adria affinity often remarkably similar to those
1167 found in the Alpine upper plate units of the Préalpes Romandes (Elter et al. 1966).
1168 Hence, the External Ligurides are regarded as representing the former ocean-
1169 continent transition within the Adria plate. Very commonly the stratigraphic base of
1170 the Oligocene part of the Epiligurian basins only exhibits a modest angular
1171 unconformity up to few tens of degrees (Piazza et al. 2016). Due to the syn-
1172 depositional sinistral transpression during Oligocene to Early Miocene times sealed
47 1173 by a Burdigalian unconformity mentioned earlier, the base of the Burdigalian and
1174 younger deposits of the Epiligurian succession locally overlies overturned older
1175 sequences of Late Cretaceous to Early Miocene successions (Piazza et al. 2016). In
1176 summary, the External Ligurides represent an ocean-continent transition area within
1177 the Adria upper plate of the Alps that was only locally overprinted by substantial late
1178 Alpine deformations of Oligocene to Early Miocene age, largely synchronous with
1179 stage 2 deformation in the Western Alps.
1180 A substantial Alpine overprint has not been documented in the southerly
1181 adjacent Tuscan Ligurides, representing a Piedmont-Liguria ophiolitic sequence
1182 identical to that observed in the Internal Ligurides or the Alps (Principi et al. 2004).
1183 Due to massive extension locally leading to core complex formation (Brunet et al.
1184 2000), Tuscany exposes deep portions of the Apenninic orogen that locally reached
1185 blueschist facies conditions with pressures up to some 1.3 GPa (Rossetti et al. 2002),
1186 mostly affecting the underlying Tuscan nappe that represents the lower plate unit
1187 during Apenninic orogeny. Locally, also the immediately overlying Tuscan Ligurides
1188 (Jolivet et al. 1998; Brunet et al. 2000; Rossetti et al. 2002) became involved in
1189 subduction during the Apenninic orogeny. Hence, it is likely that the Tuscan
1190 Ligurides only became involved in W-directed Apenninic subduction and suffered
1191 penetrative deformation within a W-dipping subduction channel. A significant
1192 ophiolitic complex that is part of the Tuscan Ligurides is exposed on Gorgona Island
1193 (Orti et al. 2002). Metasediments associated with these ophiolites suffered blueschist
1194 facies metamorphism reaching about 1.5 GPa at around 25 Ma (Rossetti et al. 2001).
1195 Gorgona island is located far offshore Tuscany (Fig. 6) and very close to the northern
1196 continuation of the new plate boundary that formed during the earliest stages of
1197 Apenninic orogeny associated with W-directed subduction and subsequent roll back
48 1198 of the mantle slab below the North Apennines (Spakman & Wortel 2004). W-directed
1199 subduction initiated east of the Corsica basin that represents a >8.5 km thick rift basin
1200 filled with undeformed Late Oligocene to Burdigalian sediments (Mauffret et al.
1201 1999; Finetti 2005) deposited as the Adriatic slab started to roll back, causing
1202 extension in the Corsica basin located in the Apenninic upper plate. The range of
1203 radiometric ages obtained for the high-pressure metamorphism in the Tuscan units is
1204 also significantly younger (26-20 Ma; Rossetti et al. 2002) when compared with the
1205 radiometric ages obtained in the ophiolitic units of Corsica where Cretaceous to latest
1206 Eocene ages (Vitale Brovarone & Herwartz 2013, and references therein) have been
1207 obtained. In summary, the Tuscan Ligurides are very likely to only have been
1208 affected by Apenninic orogeny, associated with W-directed subduction in latest
1209 Oligocene to Miocene times.
1210 The Mesozoic Adriatic passive margin stratigraphy that accreted to the
1211 overriding European plate (Corsica) and the Ligurides during the Apenninic orogeny
1212 is exposed in the Tuscan nappe. The Chattian-Aquitanian Macigno turbidites (Di
1213 Giulio 1999) represent the youngest sediments involved. This 2.5-3 km thick
1214 turbiditic sequence is testimony of the deposition of siliciclastics derived from the
1215 Alps, particularly from the Lepontine dome (Garzanti & Malusa 2008). This dome
1216 shed an immense volume of metamorphic basement estimated at 17-20 x 103 km3 (Di
1217 Giulio 1999) as the Lepontine dome was uplifted by retro-thrusting and exhumed by
1218 erosion. This led to deposition of the Macigno pile of turbidites onto the still
1219 undeformed Adriatic foredeep during retro-folding and thrusting of the Alps north of
1220 the Periadriatic Line, predating most of the backthrusting of the Southern Alps in
1221 Lombardy (Schmid et al. 1996). Related exhumation, erosion and re-deposition onto
1222 the still undeformed Adriatic continental margin are largely synchronous with stage 2
49 1223 top-WNW thrusting and WNW-directed indentation of the Ivrea mantle wedge (Stage
1224 2, 35-25 Ma).
1225 The more external Apennines thrust sheets involve flysch deposits of the
1226 foredeep only. The Cervarola thrust sheet is made up of Aquitanian to Early
1227 Langhian turbidites that still received detritus from the Alps. The more external
1228 Umbria-Marche thrust sheets involve the Marnoso Arenacea sequence of Serravallian
1229 to Tortonian age (Di Giulio 1999). The post-Messinian frontal imbricates are buried
1230 below younger deposits of the Po plain (Picotti & Pazzaglia 2008).
1231
1232 3. Lateral or temporal change in subduction polarity between Alps and
1233 Apennines? A discussion
1234
1235 Since the early days of Argand (1924) most authors (e.g. Mattauer et al. 1981;
1236 Durand-Delga 1984; Molli et al. 2006; Molli and Malavieille 2011; Handy et al.
1237 2010) regarded the pre-Oligocene Corsican nappe stack as a continuation of the
1238 Alpine orogen. Some authors (e.g. Principi and Treves 1984; Lahondère, 1996;
1239 Jolivet et al. 1998; Brunet et al., 2000), however, regarded Alpine Corsica as an
1240 integral part of the Apennines, both having formed in connection with a W- dipping
1241 subduction zone since Cretaceous times. They considered Variscan Corsica as a
1242 backstop of the Apennines orogenic wedge onto which the deepest part of the
1243 Apennines accretionary complex, namely Alpine Corsica, was backthrusted. This
1244 view essentially transfers the area of the Alps-Apennines transition more to the north
1245 into the area around Genova, which is rather unlikely in the light of the available data
1246 on Alps and northern Apennines summarized in the previous chapters.
1247 Debate also exists on whether the Alpine subduction zone with Eurasia in a
50 1248 lower plate position has continued beyond Corsica (e.g., Handy et al., 2010), or
1249 whether a lateral change in pre-Oligocene subduction polarity occurred in an area
1250 located somewhere between Corsica and Sardinia active since Late Cretaceous to
1251 Early Eocene times (e.g. Marchant & Stampfli 1997; Faccenna et al. 2003; Lacombe
1252 & Jolivet 2005; Lustrino et al. 2009; Argnani 2012; Advokaat et al., 2014; van
1253 Hinsbergen et al. 2014a). This transition is supposed to have occurred along a
1254 discrete trench-trench transform located SE of Corsica-Sardinia (e.g. Argnani 2012)
1255 or more diffusively, with partly overlapping trenches associated with opposite
1256 subduction polarity somewhere offshore east of Corsica (e.g. Lacombe & Jolivet
1257 2005). An often quoted argument in favor of this view is the existence of Late Eocene
1258 magmatism in Sardinia (Lustrino et al. 2009) within but a tiny fraction (Calabona
1259 subvolcanics) of the dominantly latest Oligocene and younger igneous magmatic
1260 province of Sardinia, starting around 28 Ma and peaking during the 22–18 Myr time
1261 range. This younger magmatic province undoubtedly has to be interpreted as related
1262 to W-directed Apenninic subduction and roll back. However, the Late Eocene
1263 magmatism at the Calabona locality was also interpreted to represent the effect of the
1264 northwest-directed subduction system that started to consume ancient oceanic
1265 lithosphere of the Tethys Ocean. This subduction would have initiated in a time span
1266 prior to the Late Eocene that is long enough to culminate in 100-150 km of
1267 subduction required to generate an arc (Lustrino et al. 2009). Given the very slow
1268 Africa-Europe convergence rates in the pre-Late Eocene (only 250 km of
1269 convergence occurred between Africa and Eurasia prior to 35 Ma at the position of
1270 Sardinia, and even less farther to the west, van Hinsbergen et al. 2014a), this would
1271 suggest that an Alpine-vergent thrusting in Sardinia and farther to the west would
1272 have been a back-thrust above a long-lived Apennines-type subduction. However
51 1273 later, Lustrino et al. (2013) argued, based on geochemical arguments, that the Eocene
1274 manifestations of the subduction related magmatism are due to the partial melting of
1275 Hercynian lower crust rather than to the dehydration processes of subducting Alpine
1276 Tethys oceanic crust. They pessimistically concluded that “whatever subduction dip,
1277 convergence rate and depth of partial melting are chosen, it is not possible to obtain a
1278 geologically sound model able to explain the spatial–temporal relationships of the
1279 western Mediterranean subduction-related igneous districts”.
1280 Since the pioneering work of Boccaletti et al. (1971) another group of authors
1281 (e.g. Guerrera et al. 1993; Doglioni et al. 1999; Michard 2006; Molli 2008; Handy et
1282 al. 2010) favored the idea of a temporal rather than lateral change in subduction
1283 polarity for the entire Apennines area (see discussion in Michard et al. 2006, their
1284 figure 9). The validity of this hypothesis is intimately linked to the question about
1285 kinematics and timing of thrusting, and, paleogeographic origin of the Calabrian-
1286 Peloritanian Belt that consists of high-grade Variscan basement nappes thrusted on
1287 top of Ligurian-type ophiolites (Schenk 1990; Vitale & Ciarcia 2013). It is also
1288 related to the question as to weather the Calabrian belt is part of an ALKAPECA
1289 block (Bouillin et al. 1986; Michard et al. 2006; Handy et al. 2010) rather than a part
1290 of Sardinia, i.e. Europe.
1291 The data and their interpretation presented in the previous chapters favor a
1292 temporal rather than lateral change in subduction polarity at least for the Northern
1293 Apennines and their transition into the Alps. However, a more detailed discussion of
1294 the southern Apennines of Calabria is outside the scope of this contribution. The
1295 controversy around a lateral vs. temporal change in subduction polarity in Calabria
1296 needs be solved by future investigations in Calabria.
1297 Although our data and interpretations support a temporal change in subduction
52 1298 polarity in the Northern Apennines realm and Corsica this interpretation faces a
1299 serious problem that is not yet solved. It is not easy to explain why this change
1300 occurred so shortly after the end of E-directed Alpine subduction in Corsica (about
1301 35-37 Ma; Vitale Bavarone & Herwartz 2013). The switch in subduction polarity
1302 around Corsica is not precisely known but probably occurred in Latest Eocene to
1303 Early Oligocene times. According to kinematic and paleomagnetic reconstructions
1304 (Séranne 1999; Gattacecca et al. 2007; Advokaat et al. 2014) fast roll back of the East
1305 Ligurian oceanic crust and rotation of Corsica-Sardinia commenced at around 30Ma,
1306 followed by rotation of Sardinia between 21-16 Ma. This only leaves some 5-7 Ma
1307 for slab break and change of subduction polarity and sufficient depth of subduction in
1308 order to trigger roll back of the W-dipping slab. Apenninic W-directed subduction
1309 and fast roll-back was certainly triggered by the negative buoyancy of large parts of
1310 the Piedmont-Liguria Ocean that still stayed open when Alpine orogeny in the
1311 Ligurian Alps ended at the end of the Eocene (Handy et al. 2010). This points to a
1312 change from plate convergence-driven Alpine E- to SE-directed subduction in
1313 Western Alps and Corsica toward W-directed subduction almost entirely driven by
1314 slab roll back in the Northern Apennines.
1315
1316 E. Summary and conclusions
1317 The geological interpretation of the P-wave velocity model of Diehl et al.
1318 (2009) allows for depicting unprecedented details concerning the outlines of this
1319 Adriatic mantle wedge and integrates its outline into the overall structure of the
1320 Western Alps. Important along-strike changes are visualized in five radially arranged
1321 crustal-scale transects across the Alps (Fig. 2) and the tectonic maps of Figs. 1 & 6).
1322 The interpretation of these findings reveals an important difference between Eastern
53 1323 and Western Alps. While the upper plate of the Eastern Alps is defined by the
1324 Austroalpine nappe system consisting of crustal imbricates of continental material,
1325 the upper plate of the Western Alps is formed by non-metamorphic klippen of the
1326 Piedmont-Ligurian Oceanic crust and lithosphere (Molli 2008).
1327 Based on this, and combined with a review of existing structural data, a 3-stage
1328 evolution is proposed for the formation of the arc of the Western Alps. A first stage
1329 (Late Cretaceous-35 Ma) is characterized by top-NNW thrusting in sinistral
1330 transpression causing at least some 260km displacement of the E-W-striking Central
1331 and Eastern Alps, together with the Adria micro-plate, towards N to NNW with
1332 respect to stable Europe. A second stage (35-25 Ma) is characterized by lateral
1333 indentation of Adria including the Ivrea mantle wedge at its leading edge towards the
1334 WNW. This culminated in some 100-150km of right-lateral strike-slip of the Ivrea
1335 mantle wedge, Western Alps and Adriatic micro-plate relative to the eastern Alps and
1336 Eurasia. This was accommodated in the Western Alpine nappe pile by top-WNW
1337 thrusting in the external zones of the central portion of the arc along the Penninic
1338 Frontal thrust combined with underthrusting of the Ivrea mantle wedge below the
1339 Western Alpine nappe pile. The third and final stage of arc formation (25-0Ma) was
1340 associated with orogeny in the Apennines and led to oroclinal bending in the
1341 southernmost Western Alps, simultaneously with the 50° counterclockwise rotation
1342 of the Corsica-Sardinia block. Particularly during the second stage of the evolution of
1343 the arc of the Western Alps the Ivrea mantle wedge, exhumed to shallow crustal
1344 depth already during Jurassic rifting at the passive continental margin of the Adria
1345 continent and representing a high strength wedge, played a crucial role by controlling
1346 the lateral ends of indentation near Cuneo and the Rhone-Simplon fault.
1347 The third stage of the evolution of the arc of the Western Alps calls for an
54 1348 analysis of the Alps-Apennines transition zone. A pre-requisite for understanding this
1349 transition is the realization that those parts of the Piemont-Liguria oceanic domain
1350 that form the upper plate in the Western Alps remained open until the Late Eocene. A
1351 combination of the gravitational potential of a southwards enlarging domain of old
1352 and oceanic lithosphere, the arrival of the Corsican continent in the Alpine trench
1353 resisting subduction, and ongoing Africa-Europe convergence induced a change in
1354 subduction polarity from Alpine S to SE-directed subduction towards NW-directed
1355 subduction and roll back of the Apennines slab, pulled by the negative buoyancy of
1356 old oceanic lithosphere large parts of which remained unaffected by Alpine orogeny.
1357
1358
1359 Acknowledgements
1360
1361 We thank Peter Nievergelt for his great help during the preparation of the
1362 figures. The careful and critical reviews of Rob Butler and Niko Froitzheim helped
1363 much in improving an earlier version of the manuscript. We thank our colleagues of
1364 the new research initiative AlpArray for many inspiring discussions, and we look
1365 forward to learn about the new insights and concepts that will certainly result from
1366 this multi-disciplinary program.
1367
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2126 Weber, S., Sandmann, S., Miladinova, I., Fonseca, R.O.C., Froitzheim, N., Münker, C.
2127 & Bucher, K. (2015). Dating the initiation of Piemont-Liguria Ocean
2128 subduction: Lu-Hf garnet chronometry of eclogites from the Theodul
2129 Glacier Unit (Zermatt-Saas Zone, Switzerland). Swiss Journal of Geosciences,
2130 108, 183-199.
2131 Weissert, H.J. & Bemoulli, D. (1985). A transform margin in the Mesozoic Tethys:
2132 Evidence from the Swiss Alps. Geologische Rundschau, 71, 665-679.
2133 Worthington, J.R., Hacker, B.R. & Zandt, G. (2013). Distinguishing eclogite from
2134 peridotite: EBSD-based calculations of seismic velocities. Geophysical Journal
2135 International, 193, 489-505. doi: 10.1093/gji/ggt004.
2136 Ye, S., Ansorge, J. Kissling, E. & Mueller, S. (1995). Crustal structure beneath the
2137 eastern Swiss Alps derived from seismic refraction data. Tectonophysics, 242,
2138 199-221.
2139 Zhao, L., Paul, A., Guillot, S., Solarino, S., Malusà, M.G., Zheng, T. , Aubert, C.,
2140 Salimbeni, S., Dumont, T., Schwartz, S., Zhu, R. & Wang, Q. (2015). First
2141 seismic evidence for continental subduction beneath the Western Alps.
2142 Geology, 43, 815-818.
2143 Zingg, A., Handy, M.R., Hunziker, J.C. & Schmid, S.M. (1990).
2144 Tectonometamorphic history of the Ivrea Zone and its relationship to the
86 2145 evolution of the Southern Alps. Tectonophysics, 182, 169-192.
2146
2147 Figure legends
2148
2149 Fig. 1:
2150 Tectonic map of the arc of the Western Alps after Schmid et al. (2008) and Bousquet
2151 et al. (2012b) with indication of the traces of the profiles depicted in Fig. 2. The
2152 selected faults indicated in red are: RSL: Rhone-Simplon Line; PL: Periadriatic Line;
2153 PF: Penninic Front; ACL: Argentera-Cuneo Line; AFSV: front of the metamorphic
2154 core of the Alps and Sestri-Voltaggio line; VVL: Villalvernia-Varzi line.
2155 Abbreviations in black are: LA: Ligurian Alps; TH: Torino hills; MH: Monferrato
2156 hills; GB: Giudicarie transpressive belt.
2157
2158 Fig. 2:
2159 Five selected geological-tectonic transects across the Western Alps. The sections are
2160 superimposed with and partly based on vertical cross sections across the final P-wave
2161 velocity model of Diehl et al. (2009) indicated by red dashed lines. The Moho (blue
2162 line) is taken from a combination of controlled source seismology, local earthquake
2163 tomography and receiver function analysis (Wagner et al. 2012; Spada et al. 2013).
2164 See text for data sources and further discussion and Fig. 1 for the traces of the
2165 profiles. a: Argentera transect; b: Pelvoux transect; c: ECORS-CROP transect; d:
2166 NFP 20-West transect; e: Ticino transect.
2167
2168 Fig. 3:
2169 Horizontal section across 3-D P-wave tomographic model of Diehl et al. (2009) at a
87 2170 depth of 20km. Yellow triangles indicate seismic stations used for tomographic
2171 inversion, crosses indicate inversion grid nodes. Note that the velocity structure is
2172 shown as percentage change relative to the 1-D initial reference model rather than in
2173 terms the P-wave velocities shown in Fig. 2. Bold black contours include well-
2174 resolved regions covering most of the area depicted. Compare Fig. 1 for major
2175 tectonic units outlined: RSL: Rhone-Simplon Line; PL: Periadriatic Line; PF:
2176 Penninic Front; ACL: Argentera-Cuneo Line. Comparison with Fig. 2 shows that the
2177 entire area with P-wave anomaly >10%, as well as a part of the area with P-wave
2178 anomaly >5% between north of Locarno and Cuneo is within the Ivrea mantle wedge.
2179
2180 Fig. 4:
2181 Three generations of stretching lineations and associated senses of shear around the
2182 arc of the Western Alps. Red arrow: stage 1 top N transport; blue arrows: stage 2 top
2183 WNW transport; green: stage 3 top SW transport; arrow pairs reflect the variation of
2184 the measured lineations; see text for further details. Numbered areas of measurements
2185 and literature sources are as follows: area 1, Valaisan and Penninic Front (Loprieno et
2186 al. 2011); area 2, Zone Houillère of the Briançonnais Unit (Bucher et al. 2004); area
2187 3, Subbriançonnais and Penninic Front (Ceriani et al. 2001; Ceriani & Schmid 2004);
2188 area 4, Penninic Front (top NW) and basal decollement of Cenozoic flysch of the
2189 Dauphinois (top SW) (Bürgisser & Ford 1998; Trullenque 2005; Trullenque et al.
2190 2006); area 5, eclogite facies lineations (top N) and greenschist facies lineations (top
2191 SW) in M. Viso ophiolites (Philippot 1988, 1990); area 6, Latest Eocene “schistes è
2192 blocs” at the base of the Embrunais-Ubaye nappes (Trullenque 2005); area 7,
2193 Penninic Front N of the Argentera massif (Trullenque 2005); area 7, thrust transport
2194 constraints from the Dauphinois of the Digne thrust system (Lickorish & Ford 1998,
88 2195 Ford et al. 2006).
2196
2197 Fig. 5:
2198 Kinematic restoration of the Alps-Apennines orogenic system using Gplates
2199 reconstruction software (e.g. van Hinsbergen et al. 2014a) holding Europe fixed; for
2200 easier orientation all time slices also show the present day position of the external
2201 massifs and Tauern window in the Alps, as well as the present-day position of
2202 northern Adria. Arrowheads indicate displacement from a given time slice to the next
2203 one, they are anchored at the position of a chosen a set of key locations at the chosen
2204 time slice. Red lines: coast lines and outlines of external massifs and Tauern Window
2205 in the Alps; blue lines: outlines of the Argentera, Pelvoux and Tambo-Suretta nappes;
2206 green area: horizontal section across the Ivrea mantle slice taken from Fig. 3. The
2207 present day position of numbered locations shown is as follows: 1: Sardinia; 2:
2208 Corsica; 3: Dora Maira massif (Briançonnais); 4: Gran Paradiso massif
2209 (Briançonnais); 5: Tambo-Suretta nappes (Briançonnais); 6: Oberammergau (front
2210 Austroalpine nappes); 7: Dimaro (eastern end Tonale Line); 8: Mauls (western end
2211 Pustertal Line; 9: Slovenske Konjice (eastern end Gailtal line); 10: Cuneo (S end of
2212 Ivrea wedge); 11: Brione (N end of Ivrea wedge); 12: Pavia (Adria microplate); 13:
2213 Monti Lessini (Adria microplate); 14: Lignano (Adria microplate); 15: W of Ancona
2214 (frontal thrust of Umbria-Marche Apennines); 16: S of Parma (frontal thrust of
2215 external Ligurides, Apennines); 17: offshore Baleares at 25 Ma ago (N front of
2216 oceanic Adria microplate = Piedmont-Liguria lower plate).
2217
2218 Fig. 6:
2219 Interpretative map sketch of the Alps-Apennines transition; see text for further
89 2220 discussion.
90 Cenozoic basins Tectonic map of the Western Alps Cenozoic in general 6° 7° 8° 9° 10° 11° Cenozoic of Piedmontese basin
Basel Lower Plate Units of the Alps European continental margin Zürich St. Gallen Dijon Cenozoic flysch Besançon (Dauphinois and Helvetic Nappes) 4 NFP 20 West 5 Ticino Innsbruck European Mesozoic undeformed 47° European Mesozoic deformed 47° Bern Allochthonous Mesozoic & Cenozoic (Helvetic Nappes)
Chur European basement (External Massifs)
European upper crust of the Subpenninic Units
High-pressure Subpenninic Units 3 Ecors CROP Bolzano Valaisan partly oceanic domain Sion RSL Valaisan Genève PL 46° Locarno Briançonnais micro-continent GB 46° Sub-Briançonnais
PF Briançonnais Aosta High-pressure Briançonnais
Pre-Piemontais Acceglio
Moûtiers Ivrea Sesia-Dent Blanche Extensional Allochthons Milano Verona Sesia, Dent Blanche, Margna, Sella Units Grenoble Piedmont-Liguria Ocean 45° 2 Pelvoux Torino TH Lower Plate Piedmont-Liguria MH 45° VVL Upper Plate Units of the Alps Briançon AFSV Upper Plate Piedmont-Liguria and distal continental margin of Adria
Adria Permo-Mesozoic Gap PF Bologna Adria upper crust Cuneo Argentera Genova Adria lower crust ACL LA Austroalpine Permo-Mesozoic 44° Digne Austroalpine basement 44° 0 20 100 km Cenozoic plutons 1 Argentera Apennines Nice Apenninic Units in general 6° 7° 8° 9° 10° 11° a. a. Argentera transect Piedmont-Liguria Upper Plate Oligocene & Miocene & Sub-Briançonnais Acceglio Piedmont-Liguria Lower Plate Miocene Pliocene Black line = Late Jurassic = Pre-Piemontais Canavese & Piedmont- Eclogites km European Mesozoic Jura Liguria Upper Plate Pliocene & 10 Dauphinois & Helvetic n. Pleistocene 10
0 Argentera 0 High-pressure Adria Mesozoic 5.5 Briançonnais Apenninic 10 Units 10 Adria upper crust 5.5 Sub-Briançonnais 6.0 European upper crust Briançonnais 20 6.0 Adria lower crust 20 6.5 6.5 6.5 7.0 30 7.0 30 7.0 European lower crust 7.5 Adria mantle 40 7.5 40
European mantle 50 8.0 50 8.0 km / s 60 60
8.0
b. b. Pelvoux transect European Mesozoic Jura Dauphinois & Helvetic n. Briançonnais Acceglio Piedmont-Liguria Sub-Briançonnais Adria Mesozoic km = Pre-Piemontais Piedmont-Liguria Upper Plate km Black band = Lower Plate Eclogites 10 Late Jurassic Miocene & Oligocene & 10 Pliocene Miocene
0 0 Belledonne Grandes Pelvoux 5.0 Rousses Apenninic Units 10 European upper crust High-pressure 5.5 10 East Variscan Briançonnais Adria upper crust 6.0 Shear Zone Valaisan 6.0 Adria lower crust 20 20 6.5 6.5 7.0 30 7.0 30 European lower crust 7.5 7.5 40 40 Adria mantle European mantle 50 50 8.0 km / s
60 60
c. c. ECORS-CROP transect Pliocene & Pleistocene European Mesozoic Jura Briançonnais Piedmont-Liguria Sesia Miocene Dauphinois & Helvetic n. Lower Plate D. Blanche km Black line = Zone Eclogites Paleogene km Valaisan 10 Oligocene & Late Jurassic 10 Miocene Canavese & Piedmont- Liguria Upper Plate
0 High-pressure 0 Briançonnais 5.0 Adria Mesozoic Belledonne 5.5 10 6.0 M. Blanc Adria upper crust 10 East Variscan 7.0 6.0 20 6.5 shear Zone Adria lower crust 20 6.5 European upper crust 7.0 6.5 30 30 7.5 7.5 7.5 40 Adria mantle 40 8.0 km / s European lower crust 7.25 50 European mantle 50
60 60 8.0 d. d. NFP 20-West transect
European Mesozoic Jura Canavese & Piedmont- Sesia Miocene Dauphinois & Helvetic n. Black line = Briançonnais Liguria Upper Plate D. Blanche Piedmont-Liguria Late Jurassic High-pressure Sesia Adria Mesozoic Black line = Oligocene & Zone Briançonnais Lower Plate D. Blanche km Valaisan Pliocene km Late Jurassic Miocene Zone 10 Valaisan & Pleistocene 10
0 0 Eclogites 5.5 Adria upper crust 6.0 M. Blanc 6.0 10 European upper crust 10 6.0 6.5 Adria lower crust 20 20 7.0 30 30 6.5 European lower crust 6.5 7.5 7.0 Adria mantle 7.0 40 European lower crust 40 7.5 European mantle 7.5 50 8.0 km / s 50 8.5 8.0 European mantle 60 60 8.0 km / s 7.5
e. e. Ticino transect High-pressure Sesia European Piedmont-Liguria D. Blanche Black band = Valaisan basement Zone km European Mesozoic Jura Lower Plate Pliocene & km Late Jurassic Briançonnais 10 Dauphinois & Helvetic n. Pleistocene 10 Eclogites Paleogene
Gotthard 0 5.5 Aar 0 Adria Mesozoic 6.0 10 10
6.0 20 Adria upper crust 6.5 20 6.5 European upper crust 30 Adria lower crust 30 7.0 European lower crust 7.25 7.0
40 7.5 7.5 40 Adria mantle 8.0 8.0 50 European mantle 50
60 8.5 km / s 60
Europe Moho documented P-wave velocity model by CSS, LET, and RF of Diehl et al. (2009) Cenozoic Cenozoic flysch (Dauphinois and Helvetic nappes) Piedmont-Liguria Ocean Adria Pliocene & Valaisan Piedmont-Liguria Paleogene European Mesozoic Jura and Miocene Pleistocene Dauphinois & Helvetic n. Upper Plate Sub-Briançonnais Canavese & Montgenevre Adria Mesozoic European upper crust nappe superieure High-pressure European Briançonnais Piedmont-Liguria Adria upper crust Apenninic Units Adula, Cima Lunga Lower Plate High-pressure Adria lower crust European lower crust Briançonnais Extensional Allochthons Acceglio European mantle = Pre-Piemontais Sesia, D. Blanche unit Adria mantle 6˚E 7˚E 8˚E 9˚E 10˚E
z = 0020 km, ref vp: 6.27 km/s
NFP-20RSL WEST Locarno
Genève Ticino PL 46˚N ECORS-CROP
PF Milano
Torino Pelvoux 45˚N Briançon
Cuneo Genova ACL
Argentera 44˚N
−20 −15 −10 −5 0510 15 20 P-wave velocity change relative to 1D-reference model (%) a. 6 b. Today 25 Ma arrow head 0 Ma 5 8 7 9 11 14 4 13 12 3 Piedmont - Liguria upper plate 16 0 100 200 km 10 15
Ligurides
2
remaining oceanic lithosphere of Adria
1 0 100 200 km 17
c. 35 Ma d. 45 Ma arrow head 25 Ma arrow head 35 Ma
0 100 200 km 0 100 200 km
Piedmont- Liguria upper plate
oceanic lithosphere of Adria ALKAPECAoceanic lithosphere of Adria
e. f. 6 55 Ma Eastern Alps and 1: Sardinia Periadriatic Line arrow head 45 Ma 2: Corsica 8 5 9 3: Dora Maira 11 7 4: Gran Paradiso 5: Tambo-Suretta Briançonnais 13 6: Oberammergau 4 14 6 12 fixed Europe 7: E end of Tonale Line 6 3 8: W end Pustertal Line 11 0 100 200 km Australpine nappes Mauls 16 10 9 9: E end Gailtal Line 10 8 5 Slov. Konjice 5 15 10: SW Ivrea Valais ocean 9 8 Continental Adria 11: NE Ivrea 7 4 12: Adria Plate Pavia 11 4 13: Adria Plate Ivrea mantle wedge 3 Monti Lessini 3 Briançonnais 14: Adria Plate 2 Ligniano 12 13 15: front Umbria- 2 10 13 Iberia 7 14 Marche Apennine 12 1 Corsica- 14 W of Ancona 2 1 Sardinia 16: external Ligurides N. Apennine Adria Plate S of Parma 16 17: N front Adria Plate 16 17 near Baleares 1 oceanic litho- 15 0 100 200 km 15 sphere of Adria 17 Piedmont - Liguria lower plate 17 Alps Innsbruck Lower Plate units of the Alps Luzern Bern undeformed or little deformed European foreland eroded part of Austroalpine Lower Plate units of the Alps nappes Chur Lower Plate units of the Alps strongly affected by Miocene-age oroclinal bending and Apennines orogeny Bolzano Sion Upper Plate units of the Alps Genève Ligurian ocean - Adria continent transition units (Canavese, “Nappes Supérieures” of Préalpes and Locarno Embrunais-Ubaye) and Chenaillet ophiolite Upper Plate units of the ocean-continent transition substantially affected by Miocene oroclinal bending and Apennines orogeny (Torino hills, Helminthoid flysch of 36 Antola nappe and Alpes Maritimes, “Nappes supérieures” 20° of Corsica) Milano Austroalpine nappes (Adria-derived continental upper Verona crustal nappe stack formed during Cretaceous orogeny 40 and displaced over the Alpine Lower Plate units in Torino the Eocene) 44 Briançon Southern Alps retro-belt derived from and thrust onto continental Adria 48 48 Undeformed continental Adria 52 Apennines 47° Bologna Oceanic upper Plate units of the Alps intensely deformed during Alpine orogeny and reworked Genova during Apennines orogeny 44 117° Internal Ligurides (Ligurian Ocean) 40 Oceanic and ocean-continent transition units little or not affected by Alpine orogeny 36 Nice Firenze Tuscan Ligurides (Ligurian Ocean)
Gorgona External Ligurides and Subligurian units island (Ligurian Ocean - Adria continent transition) Units derived from the Adria continent and the Apennines foredeep orogenic fronts of Alps and Apennines Tuscan nappe (distal Adria continental margin) topped by Chattian-Aquitanian Macigno turbidites upper plate of Apennines Cervarola thrust sheets made up of Aquitanian to Penninic Front Early Langhian turbidites contours of European and Adriatic Moho after Spada et al. (2013) Umbria-Marche thrust sheets made up of Langhian to Tortonian Marnoso Arenacia foredeep turbidites outlines of Ivrea mantle wedge at a 36 depth of 20 km (see Fig. 3) 50° post-Messinian frontal imbricates approximate former western limit of the 100 km eroded part of the Austroalpine nappes 10° E 42° N Tertiary Piedmont basin and Epiligurian basins