Open access peer-reviewed chapter

The Ampferer-Type Subduction: A Case of Missing Arc Magmatism

Written By

Mohamed A. Abu El-Rus, Ali A. Khudier, Sadeq Hamid and Hassan Abbas

Submitted: 24 November 2022 Reviewed: 08 December 2022 Published: 06 January 2023

DOI: 10.5772/intechopen.109406

From the Edited Volume

Updates in Volcanology - Linking Active Volcanism and the Geological Record

Edited by Károly Németh

Chapter metrics overview

199 Chapter Downloads

View Full Metrics

Abstract

Ampferer-type subduction is a term that refers to the foundering of hyper-extended continental or embryonic oceanic basins (i.e., ocean-continent transitions) at passive continental margins. The lithospheric mantle underlying these rift basins is mechanically weaker, less dense, and more fertile than the lithospheric mantle underlying bounded continents. Therefore, orogens resulting from the closure of a narrow, immature extensional system are essentially controlled by mechanical processes without significant thermal and lithologic changes. Self-consistent, spontaneous subduction initiation (SI) due to the density contrast between the lithosphere and the crust of ocean-continent transitions is unlikely to occur. Additional far-field external horizontal forces are generally required for the SI. When the lithosphere subducts, the upper crust or serpentinized mantle and sediments separate from the lower crust, which becomes accreted to the orogen, while the lower crust subducts into the asthenosphere. Subduction of the lower crust, which typically consists of dry lithologies, does not allow significant flux-melting within the mantle wedge, so arc magmatism does not occur. As a result of melting inhibition within the mantle wedge during Ampferer-type subduction zones, the mantle beneath the resulting orogenic belts is fertile and thus has a high potential for magma generation during a subsequent breakup (i.e., magma-rich collapse).

Keywords

  • A-type subduction
  • B-type subduction
  • orogeny
  • missing arc magmatism
  • subduction
  • subduction initiation (SI)
  • ocean-continent transitions (OCT)

1. Introduction

The term “subduction” was first used by Amstutz [1] to describe the original concept of the downward thrusting of oceanic lithosphere beneath a continental or oceanic upper plate bounded by a Wadati-Benioff zone of earthquake foci [2, 3]. This process is essential for maintaining Earth’s surface constant as new oceanic lithosphere (crust and upper mantle) is formed at mid-ocean ridges, and older lithosphere is destroyed at convergent plate boundaries. With application to the various mechanisms of lithospheric convergence, foundering, and recycling, the term subduction has become broader over time and has lost some of its original meaning. It now refers to a series of processes in which material from the Earth’s uppermost layer is submerged into the asthenospheric mantle, and its chemical constituents are recycled back into the Earth’s interior. Regardless of submerging Earth’s lithosphere process into the convective asthenosphere, two categories of subduction zones on the modern Earth (ca, 1 Ga, [4]) can be distinguished based on the association of magmatism: (a) The subduction zones associated with volcanism and (b) subduction zones missing arc-volcanicity.

The subduction zones associated with giant volcanism are referred to as Benioff-type (or B-type or Pacific-type) subduction, which is characterized by the spontaneous initiation of giant oceanic lithosphere foundering, previously formed at a mid-ocean ridge into the convective upper mantle beneath oceanic or continental upper plates [5, 6, 7, 8]. While at the ocean-ocean convergence boundary, the older and colder plate (i.e., the denser plate) often subducts the younger and warmer oceanic plate, at the ocean-continental convergence boundary, the oceanic plate subducts beneath the less dense continental plate. Although the density contrast between the oceanic lithosphere and the asthenosphere may be a possible driving force for the initiation of subduction, it is considered a second order of importance compared to convective currents in the asthenosphere that exert drag forces on the base of the lithosphere [9, 10]. However, the negative buoyancy of the sinking lithosphere (which is denser than the underlying asthenosphere) results in slab pulls, which are thought to be the dominant driving forces of plate motions [6, 11, 12, 13]. One or two planar zone(s) of seismicity in the downgoing slab (the Wadati-Benioff plane) [3, 14] can reach the mantle transition depth of ~660 km [15]. In addition, heating of the subducting crust releases a significant volume of water from the hydrated lithologies of the subducting material, leading to the fluid-flux melting within the overlying mantle wedge and the generation of hydrous, near-continuous arc magmatism [4, 8, 15, 16, 17, 18, 19]. The generated magma is lighter than the surrounding mantle material and rises through the mantle and overlying crust [20]. It creates a chain of volcanic islands on the ocean floor known as an island arc at ocean-ocean convergence margins, or it forms a mountain chain with many volcanoes known as a volcanic arc at ocean-continental convergence margins. The hydrous, calcareous magmatism with low FeO content “calc-alkaline” is predominant in the magmatic arcs [21, 22, 23, 24]. A deep trench usually forms parallel to the convergence boundary as the crust sinks downward. More volcanic material and sedimentary rocks accumulate around the island arcs, eventually thrust into an accretionary wedge and onto the continental plate [15]. As a result, high-pressure, low-temperature metamorphic facies series are common at these convergent boundaries [25, 26, 27].

The term Ampferer-type subduction was first used to describe the “missing” continental crust in the Alpine orogen (European Alps) [28]. The term was recently revived considering a new understanding of hyper-extended basins and passive margins by [7, 8, 29], who interpreted the mechanism of lithospheric recycling in the Pyrenees and Western and Central Alps. The Ampferer-type (or A-type) subduction occurs when a continent or large island collides with another continent by subducting hyperextended continental basins that contain minor oceanic crust formed at rift margins [7, 8, 20]. These hyper-thinned basins and their mechanically weak serpentinized mantle beneath serve as focal points for initiating convergence and down-thrusting of these basins beneath passive margins [7, 20, 30, 31, 32]. Only the dry root lithosphere is subducting, while hydrated lithologies from the descending plate (hydrated mantle rocks and sedimentary deposits) are sequentially accreted into a nascent orogenic wedge, resulting in amagmatic closure associated with tremendous deformation of preexisting continental rocks, forcing the material upward, creating high mountains [7, 32, 33, 34].

Since the advent of plate tectonic theory in the 1960s, which assumes that subduction of the oceanic lithosphere primarily controls rigid plate motions [35, 36], alternative concepts for the lithosphere’s foundering process have been neglected [7, 8, 20, 37]. Many convergent boundaries that exhibit Ampferer-type subduction features could have been considered Benioff-type subduction in the strict sense. Therefore, this chapter presents the well-known cases of Ampferer-type subduction zones and addresses the challenging questions of when, where, and how subduction initiation (SI) occurs at/around passive margins and the reasons for missing arc magmatism along these margins.

Advertisement

2. The Ampferer-type subduction: convergence and assembly of the previously rifted lithosphere

Ampferer-type subduction occurs only at passive continental margins that hyperextend into rift basins, often accompanied by exhumation of the (subcontinental) mantle and may reach the stage of (magma-poor) embryonic oceans [7]. Such an embryonic basin-margin system is likely to be more laterally homogeneous than ocean basins, which contain significant spreading ridges. Under these conditions, the mechanical weakness of the serpentinized mantle and the hyper-thinned continental lithosphere serve as focal points for the initiation of convergence and downward thrusting of these basins beneath passive continental boundaries [7, 30, 31, 32]. The main features for recognizing Ampferer-type subductions are (i) coherent structural units or nappes (flake tectonics) transported over long distances modified by moderate deformation and comprising the (ultra-) high-pressure continental and oceanic fragments and (ii) the lack of Penrose-type oceanic crustal fragments [7, 38]. The Ampferer-type subductions instead comprise fragments of basins flooded by exhumed subcontinental mantle, like the magma-poor Iberia-Newfoundland ocean-continent transition zones (OCTs) [39, 40, 41].

Advertisement

3. Architecture and embryonic oceans and hyperextend rift basins

Although the Wilson cycle is often used to describe the continental collision zones as formed by the closure of broad oceans floored with Penrose-type crust [42], there are numerous orogenic belts (e.g., Pyrenees, Alps, and European Variscides) having conjugate precollisional margins separated by immature extensional zones (i.e., narrow oceans or hyperextended continental rift basins) [7, 29, 34, 43, 44]. Immature extension systems are rift systems that begin when the continental crust has been stretched to complete embrittlement, typically at a crustal stretching factor of about 3–4 [45], and whose development stopped before or with the initiation of seafloor spreading [43, 44, 46, 47] (Figure 1). These rift systems involve thinning continental crust to the exhumation of the subcontinental lithospheric mantle at the floor of rift basins [43, 44, 46, 47]. In these basins, the entire remaining crust and the uppermost part of the subcontinental mantle are subject to intense hydrothermal circulation that forms sericite and illite in the crustal rocks [48, 49, 50] and serpentine, chlorite, and talc up to 4–6 km deep in the lithospheric mantle [51, 52, 53, 54]. The hyperextension of the continents can also lead to decompression melting of the underlying asthenospheric mantle [55], commonly without the formation of mid-ocean ridge basalt (MORB)-type melts, or they are only partially extracted and tend to stagnate in and fertilize the overlying lithospheric mantle [56]. This process tends to homogenize (on a large scale) the uppermost lithospheric mantle (~30 km) beneath the rift basins into plagioclase-bearing lherzolite [41, 56]. However, growing evidence suggests that fertilization of the lithospheric mantle beneath the rift basins is uneven and gradually increases from the proximal to the unstretched continental part to an ideal homogeneous plagioclase-bearing peridotite beneath the distal part of the rift basin [41].

Figure 1.

(a) Various stages of extension and corresponding physical properties of the lithospheric mantle at the center of the rift. (b) Architecture and lithology of a typical magma-poor rifted margin [44].

The fertilized mantle under immature oceanic basins, such as beneath the southern part of the Porcupine Basin [57] or at hyperextended rifted margins such as the Iberian margin [58], the Newfoundland margin [59], and the Flemish Cap [60], is also characterized by a reduction in seismic velocities of about 1% at pressures up to 1 GPa and by >2% at higher pressures [44]. Although serpentinization is likely the main cause of the velocity decreases up to 6 km below the seafloor, it may not be responsible for reducing the seismic velocity at greater depths as the serpentine becomes unstable [44]. However, the presence of plagioclase in the fertilized mantle strongly affects the rheology of the lithospheric mantle because plagioclase is weaker than pyroxene and olivine [44]. As a result, the fertilized mantle behaves semi-brittle (i.e., the plagioclase stability field) between 18 and 40 km below the embryonic oceans and hyperextended rift basin. This mantle domain, where deformation can be accommodated along localized anastomosing shear zones, is broader than in the lithospheric mantle beneath the bounded continental margins and could have significance as a stress guide for subduction initiation of rift basins beneath passive continental margins at greater depths up to 30 km [44].

The absence of significant subduction-related magmatism in orogens formed from Ampferer-type subduction zones is interpreted because of the narrow width of the closing ocean, which did not allow significant decompression and/or flux melting [61, 62]. Therefore, the mantle beneath orogens resulting from the closure of a hyperextended rift basin or narrow embryonic ocean is likely fertile because it is hydrated and enriched in mobile constituents derived from subducting sediments, oceanic crust, and dehydrating serpentinite but lacks significant flux melting [63]. This fertile mantle could provide fusible components for intense magmatism during subsequent collapse (magma-rich orogenic collapse [43, 44]. This is shown in the widespread mafic to acidic intrusions, for instance in the crust in the Variscan region, Basin and Range province, and Canadian Cordillera [64, 65, 66, 67]. Conversely, orogenic belts formed by the closure of a large ocean are associated with intense flux melting within the mantle wedge due to the release of substantial fluids from dehydration of the large, subducted slab and decompression melting of the hot asthenosphere that rises to compensate for the down dragged mantle wedge material by the slab [68, 69, 70]. These melting processes form island arcs at the sea-flower or volcanic arcs at the continental margin [71] and deplete the source mantle wedge [72]. Orogens formed by the closure of a large ocean may therefore be underlain by a relatively depleted mantle [43]. Therefore, the subsequent collapse of these orogenic belts is devoid of magma (magma-poor collapse). The Western Europe-North Atlantic region is an ideal example to examine how shortening or incomplete Wilson cycles may differ from classic Wilson cycles, as it consists of orogens resulting from the closure of both narrow oceans (< 500 km) or immature, hyperextended rift systems (the Variscides of Western Europe) and broad, mature oceans (the Scandinavian Caledonides) (Figure 2) [44, 73].

Figure 2.

Two end-members proposed for the Wilson cycle in the North Atlantic as it is resulting from the closure of both broad, mature oceans (the Scandinavian Caledonites), and narrow oceans (<500 km) or immature, hyperextended rift systems (the Variscides of Western Europe) [44].

Advertisement

4. Subduction zone initiation (SZI)

Two concepts are commonly proposed for subduction zone initiation (SZI): (i) spontaneous or vertically forced and (ii) induced or horizontally forced SZI [74, 75, 76]. Vertically forced SZI is caused by the contrast between the underlying convicting mantle and the cooling lithosphere above. The viability of this scenario is controversial [77], especially for SZI at passive margins [78]. The body force resulting from density differences is most likely insufficient to break up and initiate a subduction zone if the rift basin lithosphere is older than 20 My after the continental breakup [79]. This assumption is not consistent with the ages of descending, hyperextended continental basins, which are generally older than 20 My (e.g., ~34 My for Oligocene subduction of the New Caledonia Basin beneath the northern Norfolk Ridge, SE Pacific; [20] and ~ 60–65 My for initiation of subduction of Piemonte-Liguria Ocean beneath the Adriatic continental margin, [78]. Moreover, data from recent and ancient subduction zones [76, 80] suggest that vertically forced SZI was probably not the dominant scenario during the last 100 million years. Therefore, the horizontally forced SZI model is preferred to overcome the increasing strength of the cooling lithosphere [76, 78, 81]. The far-field external horizontal forces can be caused by the mid-ocean ridge, mantle plume, neighboring sinking slab, and large-scale mantle convection [82]. However, [83] argues that vertical forces can accelerate, propagate, and facilitate the development of self-sustaining subduction zones that are initially dominated by horizontal forces.

It is most likely that the initiation of subduction zones by horizontal compressive forces requires a process that mechanically weakens or softens the lithosphere to aid in the localization stresses that eventually lead to the breakup [20]. Several mechanisms have been proposed for softening the lithosphere during SZI including mineral reaction and transformation [84], fluid-induced [20, 85, 86], microstructural evolution and anisotropy [87], mineral grain damage and plunging [88, 89], and thermal softening [78, 90, 91]. Petrological thermomechanical models show that a temperature increase of only ca. 50°C is sufficient for successful SZI [78, 92]. Therefore, thermal softening is considered a potential mechanism to form a shear zone transecting the lithosphere [90, 93, 94, 95, 96, 97] and initiate subduction at a hyperextended continental [78]. Serpentinization within the lithosphere may also serve as a stress conductor during subduction initiation because it is weaker than peridotite, continental crust, and oceanic crust [98, 99, 100]. In addition, greater serpentinization would reduce the coherence of the rift basin lithosphere [7], which generally exhibits shallow slab detachment (SD) early after the SZI [78]. In contrast, the low-serpentinized lithosphere is coherent and requires more compressional forces for SZI [79, 101]. The strong lithosphere may be able to maintain a continuous subducting slab down to 660 km depth for more than 20 Myr after basin closure [78].

Advertisement

5. Missing arc magmatism

Well-documented examples of subduction initiation involving mature oceanic lithosphere (Benioff-type subduction zones, e.g., the Neotethys supra subduction zone and the Izu-Bonin-Mariana arc) are characterized by an initial phase of upper plate extension and tholeiitic to boninitic magmatism [102, 103]. Once initiated, partial eclogitization and densification of the subducting oceanic lithosphere results in a slab-pull mechanism that is a driving force for self-sustaining subduction and ocean closure [6]. Dehydration of oceanic crust drives flux melting of the overlying mantle wedge and lower crust of the overriding plate, resulting in predominantly “calc-alkaline” magmatism [104, 105, 106, 107, 108]. Therefore, the presence of ophiolites associated with calc-alkaline magmatism and low-temperature, high-pressure metamorphic rocks are interpreted as unequivocal evidence for paleosubduction zones [109]. In contrast, subduction zones with the immature oceanic lithosphere (Ampferer-type subduction zones, e.g., the European Variscides, Pyrenean Belt, and Alpine Mountains) are characterized by ophiolites, subduction-related metamorphism, accretionary prisms, and syn-orogenic clastic sediments but and lacking calc-alkaline magmatism [7, 8, 20, 29]. Although the absence of calc-alkaline magmatism within the Ampferer-type subduction zone is commonly attributed to the narrow width of the closing ocean (<500 km), which did not allow for significant decompression and/or flux melting and thus habitable arc magmatism [61, 62], a recent study of a case of Ampferer-type subduction zones beneath New Caledonia in the Oligocene revealed that other causes of missing arc magmatism are also relevant [20]. The Oligocene subduction zone beneath New Caledonia is an ideal Ampferian-type paradigm for investigating the reason for missing the arc magmatism associated with Ampferian-type subduction zones because its paleogeographic elements have not been jumbled together by collisional deformation or dismembered by strike-slip faulting. In addition, the well-documented evolutionary history of New Caledonia and its relative tectonic simplicity is appropriate for specifying the operational processes within the Oligocene Ampferer-type subduction zone [20].

The New Caledonia Island is a remnant of Gondwana continental crust that is part of the Norfolk Ridge (Figure 3), which began to drift away from the eastern Australian margin during the opening of the Tasman Sea in the Late Cretaceous [110, 111, 112]. Norfolk Ridge is a suitable location for commencing Ampferer-type subduction due to the hyperextension of the east border of the Australian continent, which curtailment in tiny slivers in this region [20, 112]. The magmatism associated with the Oligocene subduction zone is rare and represented by in situ minor eruptive basalt-andesite lava flows of La Conception (c. 29.12 My Ar/Ar age, [20]) and two small isolated Oligocene granodiorite massifs (c. 24 My [113]) that are not far from the La Conception lavas: the Saint Louis Massif, located about 2 km to the east, and the Koum/Borindi Massif, located ∼55 km to the northeast (Figure 3). Thus, the correlation of La Conception lava with Saint Louis and Koum/Borindi Massifs is of great importance for studying the cross-geochemical trend associated with the Ampferer subduction zone, which is critical for understanding the dynamics of the mantle wedge, including how it convects, where and how much melting occurs, how much melt extraction affects source heterogeneity, and how non-subduction heterogeneity affects arc lavas. However, a recent study of the Oligocene subduction zones beneath New Caledonia [20] revealed the following reasons that may result in missing the arc magmatism in association with Ampferer-type subduction zones:

Figure 3.

The generalized geology map of New Caledonia shows the locations of all eight terranes. Red stars show locations of Oligocene magmatism (La conception lava quarry, the Saint Louis Massif [SLM], and the Koum/Borindi Massif [KBM]) [20].

5.1 Cold subduction zone

The opening of the New Caledonian Basin occurred during the rifting of the eastern margin of the Australian-Gondwanan continent in the Late Cretaceous time [114, 115]. This implies that the crust of the New Caledonian Basin was about 34 Ma old (i.e., cold crust) [79, 80, 81, 82, 83, 84, 85, 86, 87, 88, 89, 90, 91, 92, 93, 94, 95, 96, 97, 98, 99, 100, 101, 102, 103, 104, 105, 106, 107, 108] at the time of its subduction beneath the northern Norfolk Ridge in the early Oligocene (ca. 32 Ma) [116]. The mantle wedge overlying the descending plate was also cold. Its potential temperature (Tp °C), estimated from the composition of the La Conception lava, ranges from 1268 to 1316°C [20]. This estimate is below the potential temperatures of the ambient mantle (i.e., ∼1400°C) [117] and below the 1350°C required for arc basalt magma formation within the mantle wedge [118]. In addition, trace element modeling of the La Conception lavas suggests that the melting source was 119–129 km within the mantle wedge [20]. These constraints are consistent with cold subduction zones, where large amounts of water can be transported in the gabbro and peridotite layers of the subducted slabs to depths greater than 200 km, where it hydrates and partially melting the mantle wedge, in contrast to warm subduction zones, where the entire subducted slab becomes anhydrous at shallow to intermediate depths and the mantle wedge only at shallow depths [119, 120].

5.2 The high sinking velocity of the subducted slab

Given that the emplacement of the La Conception lavas occurred around 29.12 Ma and that the Oligocene subduction zone began at approximately 33.7–35 Ma at a location about 50 km west of New Caledonia (as indicated by a + 100 mGal gravity anomaly, [121, 122], the sinking velocity of the subducted slab to depth ∼119–129 km within the mantle wedge where the La Conception lavas generated [20] ranges from 4.0 to 5.1 cm/year. Such a rapid sinking velocity, combined with the cold nature of the subducting slab, may not allow enough interaction between the sinking slab and the mantle wedge to produce significant volumes of arc magmas [123, 124].

5.3 The depletion inheritance of the mantle beneath the continental border

The depletion heritage of the mantle beneath the continental border considering that Nb and Ta are least effectively transferred from the subducted oceanic slab to the overlying mantle melting column, the ratio of elements to Nb (or Ta) would measure the non-conservative character of the elements (% sz), which expresses the extent of their displacement from the average MORB at a given Nb [125]. Calculations of % sz for the elements in La Conception show that Rb, Pb, Th, Ba, U, K, and Pb are strongly non-conservative with % sz > 80%; LREE, Sr., and P2O5 are moderately non-conservative with a % sz range of 80–40%; Sm and Nd are non-conservative to slightly non-conservative with % sz from 40% to the detection limit, whereas other elements exhibited negative % sz values (Figure 4) [20]. The negative % sz values suggest that the mantle beneath the Norfolk Ridge was relatively depleted before the onset of the Oligocene subduction zone, which may partially inhibit the arc magmatism.

Figure 4.

The percentage of the subduction component (%sz) of the La Conception lavas is estimated from the displacement of the average mid-oceanic-ridge basalt (MORB) at a given Nb [20].

5.4 The continental nature of the subducting slab

The nature of the New Caledonian Basin itself may also play a significant role in reducing arc volcanism because the New Caledonian Basin is a continental crust that is commonly dominated by dry granitic material [126, 127]. Furthermore, unlike the oceanic lithosphere, the lithospheric mantle underlying continental crust does not contain voluminous hydrous mineral phases [7, 29]. Therefore, the sinking slab of the New Caledonian Basin beneath the Norfolk Ridge did not provide the hydrous fluids necessary to lower the solidus of the mantle wedge and produce the arc volcanism associated with the Oligocene subduction zone [20].

5.5 Formation of a magmatic barrier at the base of the overriding plate

Seismic tomography shows evidence of a deep reflector at a depth of 60 km (∼18 kbar) beneath the west coast of New Caledonia, and this structure was found to dip northward by at least 30° [115]. This deep seismic discontinuity was attributed to a remnant of a sinking slab of a subduction zone that formed in response to the blocking of an older subduction east of the Norfolk Ridge in the early Oligocene (ca. 32 My) [116]. However, this depth estimate for the subducted slab is not consistent with the fractionated REE and high Sr/Y content of the La Conception lava [20], suggesting the presence of residual garnet in the melting source (i.e., melting within the garnet stability field >25 kbar) [128]. Therefore, we believe that the seismic mirror beneath New Caledonia is the solidification front of accumulated mantle-derived melts at the base of the overriding plate. The incipient melt slowly percolates upward (i.e., porous flow), mainly in a vertical direction due to the buoyancy of magma and the high permeability of partially molten systems [129, 130], and eventually, it enters the cooler lithosphere where it begins to solidify [131]. This would cause of the development of a low-permeability zone or permeability barrier [131, 132], under which the subsequent pluses of magma accumulate, forming a melt-rich zone [133, 134, 135].

Advertisement

6. Examples of Orogenic belts resulting from Ampherer-subduction zones

Ideal examples of orogens resulting from the closure of a hyperextended rift system or an immature ocean include the Alps [34], Pyrenees [136, 137], and European Variscides [138, 139]. These orogens often occurred without significant subduction-related magmatic activity, so the original rock units before the collision were relatively well preserved [34, 43, 44, 137, 140].

6.1 Closure of the Piemonte-Liguria Ocean and Alpine orogeny

The Alpine orogen, upon subduction initiation at ~85–100 Ma [30, 141, 142], shows no evidence of magmatic activity during subduction initiation combined with a ~ 50 Ma hiatus in magmatism, or “arc gap” [29]. The orogen contains fragments of the Piemont-Liguria Ocean, which was not a Penrose-type oceanic crust. The Piemont-Liguria Ocean was formed in the Jurassic period when the paleocontinents Laurasia (to the north, with Europe) and Gondwana (to the south, with Africa) started to move away from each other [143]. The Piemont-Liguria Ocean contained several continental blocks such as the Adriatic plate (also known as the Apulian Plate) and Sesia-Dent Blanche unit separated from Gondwana and Briançonnais block separated from Laurasia (Figure 5). The Gondwanan continental block Adriatic plate was moving northward into Laurasia. In the Cretaceous period, the Piemont-Liguria Ocean lay between Europe (and a smaller plate called the Iberian plate) in the northwest and the Adriatic plate in the southeast. The European and Adriatic passive margins were likely hyperextended, while the Piemonte-Liguria ocean basin was mainly floored by an exhumed mantle and hyperextended continental crust rather than mature oceanic crust (e.g., [7, 34, 78, 144]. At ca. 85–80 Ma, subduction at the Adriatic margin led to subduction of the Sesia-Dent Blanche unit [145], and progressed from southeast to northwest across the PL Ocean (ca. 50–45 My), the Briançonnais (ca. 45–42 My), and the Valais Basin (ca. 42–35 My) as indicated by the tectono-metamorphic evolution of those units (Figure 5 [142, 145, 146, 147, 148]. However, the onset of subduction initiation at the passive margins of Laurasia allowed the accretion of the hydrated portion of the subducting of the remaining Piemont-Liguria Ocean plate within an orogenic Alpine wedge (Figure 5) [29]. It is unclear whether the subducted slab is currently separated or remains attached to the European plate (Figure 5) [149].

Figure 5.

Simplified conceptual geodynamic sketch of the Alpine orogeny after Candioti et al. [78]. (a) Passive margin geometry and embryonic oceans after the rifting phase. (b) Horizontally forced SZI on the Adriatic passive margin and subduction of the Sesia-Dent Blanche unit. (c) Possible scenarios for the collisional phase of Alpine orogeny with a slab that is continually subducting (left) or a slab that is separated (right).

6.2 Iberian plate kinematics and Pyrenees orogeny

The Pyrenean orogeny in southwestern Europe is an ~E-W trending mountain belt, about 450 km long and 125 km wide (Figure 6) that formed in the Late Cretaceous to Paleogene in response to convergence between the hyperextension NE continental margins of the Iberian Microplate and the Southwest Eurasian Plate [150]. The Iberia microplate was part of Pangea in the Paleozoic [151] but separated in the Late Jurassic [152, 153, 154]. Based on magnetic lineation in the Atlantic Ocean and Bay of Biscay to the west, the separation of the Iberian Microplate was N-S directional rifting before it rotated counterclockwise early Aptian (e.g., Scissor-type opening) (Figure 7) [155, 156, 157, 158]. Analysis of the paleomagnetic record suggests a ~ 35° counterclockwise rotation of the Iberia microplate completed in the early Aptian (126 ~ 118 My) [150]. Such a high amount of rotation should be associated with subduction beneath the Eurasian margin [150, 157, 158] and the North Pyrenean fault zone represents the suture between Iberia and southwest Eurasia (Figure 6) [150]. Structural and deep seismic studies have shown that the orogen is asymmetric, with the Iberian continental lithosphere underthrust at least ~80 km beneath Europe [ 159, 160]. However, several tomographic studies show no evidence of a subducted slab anywhere beneath the Pyrenees [161, 162, 163], ruling out the opening of a broad oceanic basin prior to Late Cretaceous convergence [161]. Vissers et al. [150] explained the absence of a remnant subducted slab by the motion of Iberia/Eurasia relative to the mantle, where the Pyrenean region may have laterally displaced from a subducting slab remnant after slab break-off. Mantle tomography indicates that such a slab remnant may exist today between 1900 and 1500 km depth beneath southern Algeria [150].

Figure 6.

(a) Geological sketch map of the Pyrenees [150]. Blue line indicates ECORS seismic section shown in (b). White diamonds denote mantle peridotite bodies. NPZ, North Pyrenean Zone; NPF, North Pyrenean Fault; SPZ, Southern Pyrenean Zone; B, Boixols thrust. (b) ECORS crustal-scale cross-section Pyrenees [150].

Figure 7.

Scissor-type scenario [159], for the plate kinematic of the Late Mesozoic motion of the Iberian Peninsula to Europe, with the inferred position of the Iberian Peninsula at M0 times (left panel) and the onset of the Alpine collision (right panel). NA, North America; IB, Iberia; EUR, Europe; NGFZ, Newfoundland-Gibraltar Fracture Zone. Circles with crosses denote the poles of the overall reconstruction. Circles labeled M0-A33 denote the poles of the phase describing the opening of the Bay of Biscay.

6.3 Closure of the Paleozoic Rheic ocean and Variscan orogeny

The Variscan Belt is a segment of a mountain system that has existed all around the world because of a sequence of Paleozoic collisional orogenic events (e.g., the Appalachians in America, Mauritanides in Africa, the Caledonides in Scandinavia and Scotland, the Urals in Russia, the Tien Shan in Asia, and the Lachlan Fold Belt in Australia) that marked the amalgamation of the supercontinent Pangea [138, 139] (Figure 8). The Variscan belt extends across Europe, with the best exposure in central and western Europe and parts of Morocco and Algeria to the north of the West African Craton [138, 165]. It formed between 480 and 250 My after the collision of Gondwana to the south and Baltica-Laurentia to the north [164, 166, 167]. The Variscan belt is a much more complex paleogeographic orogenic belt whose elements were not only shuffled together by collisional deformation but were also dismembered by strike-slip faults and the formation of oroclines [167]. Reviewing the evolution of the main tectonic units and their paleogeography position within the Variscan is beyond the scope of our work, and the reader may refer to [138, 167, 168, 169] for further details. However, the presence of suture zones characterized by ophiolites, subduction-related metamorphism and magmatism, accretionary prisms, and foreland basins with syn-orogenic clastic sediments within the European Variscides confirms the existence of several microcontinents between Laurussia and Gondwana. These microcontinents include the blocks of Avalonia, North Armorica (Franconia + Thuringia, separated by the failed Vesser Rift), and South Armorica (south-western central Iberia, north-central Armorican Massif, Bohemia, all probably united, and Paleo-Adria) (Figure 9) [138, 139, 166, 167, 170, 171]. Several authors [138, 170171] have proposed that the Variscan microcontinents separated from the north-Gondwana margin by back-arc spreading and southward subduction. Although this model may apply to the Avalonia microcontinent and the opening of the Rheic Ocean, it cannot explain the existence of other microcontinents and intervening rift/drift zones during the Cambrian and Ordovician [167]. Alternatively, a system of mantle plume activity beneath the North Gondwanan margin has been hypothesized to result in the opening of the Rheic and other peri-Gondwanan oceans [172, 173, 174]. Only the Rheic (between Avalonia and the Armorican microcontinents) evolved into a sizeable ocean documented by biogeography and paleomagnetism [167]. The other basins (i.e., the Saxo-Thuringian Ocean between north and South Armorica and the Galicia-Moldanubian Ocean between South Armorica and Palaeo-Adria) did not grow beyond the narrow ocean stage and did not pass through the complete Wilson Cycles [167]. Closure of the narrow oceans and their subduction beneath microcontinents in order from south to north, between the middle Devonian and the Tournaisian gave birth to orogenic belts with HP-UHP metamorphism with missing arc magmatism (i.e., Ampferer-type subduction).

Figure 8.

Permian (at 280 Ma) assembly of the continents showing the Paleozoic belts. Yellow, 400 ± 250 Ma; orange, 450 ± 400 Ma [164].

Figure 9.

Tentative reconstruction of the Variscan and surrounding areas at successive intervals from middle Ordovician (465 Ma) to Lower Carboniferous (340 Ma) [166] continental microplates. Blue, island arcs; black squares, distribution of Callixylon (petrified trees) in the Late Devonian (375 Ma).

Advertisement

7. Conclusions

Orogenic and post-orogenic magmatism is primarily controlled by the size and maturity of the basin/oceanic crust involved in the subduction zones. Orogenic belts formed from the closure of narrow oceans or hyperextended rift basins (Ampferer-type subduction) are characterized by the absence of arc magmatism but post-orogenic magma-rich collapse. In contrast, orogenic belts formed from the closure of broad oceans floored with Penrose-type crust (Benioff-type subduction) are characterized by arc magmatism but post-orogenic magma-poor collapse.

Advertisement

Acknowledgments

The present work was carried out with logistical and financial support from the Geology Department of Assiut University.

Advertisement

Conflict of interest

All figures are used with permission.

References

  1. 1. Amstutz A. Structures Alpines: Subductions Successives Dans l’Ossola. Vol. 241. Paris: Comptes rendus de l’Acad’emie des sciences; 1955. pp. 967-969
  2. 2. Wadati K. On the activity of deep-focus earthquakes in the Japan islands and neighbourhoods. The Geophysical Magazine. 1935;8:305-325
  3. 3. Benioff H. Seismic evidence for the fault origin of oceanic deeps. Geological Society of America Bulletin. 1949;60:1837-1856. DOI: 10.1130/00167606(1949)60[1837:SEFTFO]2.0.CO;2
  4. 4. Stern RJ. The evolution of plate tectonics. Philosophical Transactions. Series A, Mathematical, Physical, and Engineering Sciences. 2018;376:20170406. DOI: 10.1098/rsta.2017.0406
  5. 5. Carlson RL, Hilde TWC, Uyeda S. The driving mechanism of plate tectonics: Relation to age of the lithosphere at trenches. Geophysical Research Letters. 1983;10:297-300. DOI: 10.1029/GL010i004p00297
  6. 6. Cloos M. Lithospheric buoyancy and collisional orogenesis: Subduction of oceanic plateaus, continental margins, island arcs, spreading ridges, and seamounts. Geological Society of America Bulletin. 1993;105:715-737. DOI: 10.1130/0016-7606(1993)105<0715:LBACOS>2.3.CO;2
  7. 7. McCarthy A, Tugend J, Mohn G, Candioti L, Chelle-Michou C, Arculus R, et al. A case of Ampferer-type subduction and consequences for the Alps and the Pyrenees. American Journal of Science. 2020;320:313-372. DOI: 10.2475/04.2020.01
  8. 8. Chelle-Michou C, McCarthy A, Moyen JF, Cawood PA, Capitanio FA. Make subductions diverse again. Earth Science Reviews. 2020;226:103966. DOI: 10.1016/j.earscirev.2022.103966
  9. 9. Doglioni C. A proposal for the kinematic modelling of W-dipping subductions-possible applications to the Tyrrhenian-Apennines system. Terra Nova. 1991;3:423-434. DOI: 10.1111/j.1365-3121.1991.tb00172.x
  10. 10. Ziegler PA. Plate tectonics, plate moving mechanisms and rifting. Tectonophysics. 1992;215:9-34. DOI: 10.1016/0040-1951(92)90072-E
  11. 11. Ueda K, Gerya T, Sobolev SV. Subduction initiation by thermal–chemical plumes: Numerical studies. Physics of the Earth and Planetary Interiors. 2008;171:296-312. DOI: 10.1016/j.pepi.2008.06.032
  12. 12. Ganguly J, Freed AM, Saxena SK. Density profiles of oceanic slabs and surrounding mantle: Integrated thermodynamic and thermal modeling, and implications for the fate of slabs at the 660 km discontinuity. Physics of the Earth and Planetary Interiors. 2009;172:257-267. DOI: 10.1016/j.pepi.2008.10.005
  13. 13. Boonma K, Kumar A, Garcia-Castellanos D, Jimenez-Munt I, Fernandez M. Lithospheric mantle buoyancy: The role of tectonic convergence and mantle composition. Scientific Reports. 2019;9:17953. DOI: 10.1038/s41598-019-54374-w
  14. 14. Hasegawa A, Umino N, Takagi A. Double-planed structure of the deep seismic zone in the northeastern Japan arc. Tectonophysics. 1978;47:43-58. DOI: 10.1016/0040-1951(78)90150-6
  15. 15. Webb P. Introduction to Oceanography. Mountain View, CA: Roger Williams University; 2021. p. 383 Available from: http://rwu.pressbooks.pub/webboceanography
  16. 16. Gill JB. Orogenic Andesites and Plate Tectonics. Berlin Heidelberg: Springer; 1981. p. 392. DOI: 10.1007/978-3-642-68012-0
  17. 17. Tatsumi Y, Sakuyama M, Fukuyama H, Kushiro I. Generation of arc basalt magmas and thermal structure of the mantle wedge in subduction zones. Journal of Geophysical Research - Solid Earth. 1983;88:5815-5235. DOI: 10.1029/JB088iB07p05815
  18. 18. Abu El-Rus M A: Geological studies on Abu Ghalaga area, Eastern Desert, Egypt [M.Sc. thesis]. Assiut: Assiut University; 1991: p. 255
  19. 19. Arculus RJ. Evolution of arc magmas and their volatiles: The state of the planet. Frontiers and Challenges in Geophysics, Geophysical Monograph Series. 2004;150:95-108. DOI: 10.1016/0024-4937(94)90060-4
  20. 20. Nicholson KN, Abu El-Rus MA. A case of Ampferer-type subduction beneath the New Caledonia arc: Evidence for inefficient subduction of hydrated lithologies into the upper mantle. Geological Society of America Bulletin. DOI: 10.1130/B36166.1
  21. 21. Ahmed AA, Abu El-Ru MA. Pre-orogenic magmatic phases of Baranis Area, southeastern Desert, Egypt. Bulletin of Faculty of Sciences, Assiut University. 2002;32:73-97
  22. 22. Arculus RJ. Use and abuse of the terms calcalkaline and calcalkalic. Journal of Petrology. 2003;44:929-935. DOI: 10.1093/petrology/44.5.929
  23. 23. Khudier AA, Abu El-Rus MA, El-Gaby S, El-Nady O. Geochemical and geochronological studies on the infrastructural rocks of Meatiq and Hafafit swells, Eastern Desert, Egypt. Egyptian Journal of Geology. 2006;50:190-214
  24. 24. Khudier AA, Abu El-Rus MA, El-Gaby S, El-Nady O, Bishara WW. Sr-Nd isotopes and geochemistry of the infrastructural rocks in the Meatiq and Hafafit core complexes, Eastern Desert, Egypt: Evidence for involvement of pre-Neoproterozoic crust in the growth of Arabian-Nubian Shield. Island Arc. 2008;17:90-108. DOI: 10.1111/j.1440-1738.2007.00599.x
  25. 25. Maekawa H, Shozui M, Ishii T, Fryer P, Pearce JA. Blueschist metamorphism in an active subduction zone. Nature. 1993;364:520-523. DOI: 10.1038/364520a0
  26. 26. El-Fadly MA, Hamid SA, El-Rus MA, Khudeir AA. An inquiry into the structural evolution of the Neoproterozoic Shait Granite Complex, south Eastern Desert, Egypt. Assiut University Journal of Geology. 2018;47:41-56
  27. 27. Hamid S, Abu El-Rus MA, El Kazzaz YA, Khudeir AA. Neoproterozoic tectono-metamorphic evolution of Hafafit Dome “A” and the abutting Shait ophiolitic mélange domain at Gabel Mudargag area, central Eastern Desert, Egypt. Assiut University Journal of Geology. 2020;49:1-25
  28. 28. Ampferer O, Hammer H. Geologischer Querschnitt durch die Ostalpen vom Allgäu zum Gardasee. Jahrbuch der Kaiserlich Königlichen Geologischen Reichsanstalt. 1911;61:631-710
  29. 29. McCarthy A, Chelle-Michou C, Müntener O, Arculus R, Blundy J. Subduction initiation without magmatism: The case of the missing Alpine magmatic arc. Geology. 2018;46:1059-1062. DOI: 10.1130/g45366.1
  30. 30. Zanchetta S, Garzanti E, Doglioni C, Zanchi A. The Alps in the cretaceous: A doubly vergent pre-collisional orogen. Terra Nova. 2012;24:351-356. DOI: 10.1111/j.1365-3121.2012.01071.x
  31. 31. Tugend J, Manatschal G, Kusznir NJ, Masini E, Mohn G, Thinon I. Formation and deformation of hyperextended rift systems: Insights from rift domain mapping in the Bay of Biscay-Pyrenees. Tectonics. 2014;33:1239-1276. DOI: 10.1002/2014tc003529
  32. 32. Auzemery A, Yamato P, Duretz T, Willingshofer E, Matenco L, Porkol’ab K. Influence of magma-poor versus magma-rich passive margins on subduction initiation. Gondwana Research. 2022;103:172-186. DOI: 10.1016/j.gr.2021.11.012
  33. 33. Ernst WG. Alpine and Pacific styles of Phanerozoic mountain building: Subduction-zone petrogenesis of continental crust. Terra Nova. 2005;17:165-188. DOI: 10.1111/j.1365-3121.2005.00604.x
  34. 34. Mohn G, Manatschal G, Beltrando M, Haupert I. The role of rift-inherited hyper-extension in Alpine-type orogens. Terra Nova. 2014;26:347-353. DOI: 10.1111/ter.12104
  35. 35. Wilson JT. Did the Atlantic close and then re-open? Nature. 1966;211:676-681. DOI: 10.1038/211676a0
  36. 36. Wilson JT. Static or mobile earth: The current scientific revolution. Proceedings of the American Philosophical Society. 1968;112:309-320
  37. 37. Trümpy R: Why plate tectonics was not invented in the Alps. International Journal of Earth Sciences 2001; 90:477-483. https://doi.org/ 10.1007/s005310000175
  38. 38. François C, Pubellier M, Robert C, Bulois C, Jamaludin SNF, Oberhänsli R, et al. Temporal and spatial evolution of orogens: A guide for geological mapping. Episodes. 2022;45:265-283. DOI: 10.18814/epiiugs/2021/021025
  39. 39. Manatschal G, Müntener O. A type sequence across an ancient magmapoor ocean–continent transition: The example of the western Alpine Tethys ophiolites. Tectonophysics. 2009;473:4-19. DOI: 10.1016/j.tecto.2008.07.021
  40. 40. Mohn G, Manatschal G, Masini E, Müntener O. Rift-related inheritance in orogens: A case study from the Austroalpine nappes in Central Alps (SE-Switzerland and N-Italy). International Journal of Earth Sciences. 2011;100:937-961. DOI: 10.1007/s00531-010-0630-2
  41. 41. Picazo S, Müntener O, Manatschal G, Bauville A, Karner G, Johnson C. Mapping the nature of mantle domains in Western and Central Europe based on clinopyroxene and spinel chemistry: Evidence for mantle modification during an extensional cycle. Lithos. 2016;266-267:233-263. DOI: 10.1016/j.lithos.2016.08.029
  42. 42. Dewey JF, Burke K. Hot spots and collisional break-up: Implications for collisional orogeny. Geology. 1974;2:57-60. DOI: 10.1130/00917613(1974)2<57: HSACBI>2.0.CO;2
  43. 43. Chenin P, Manatschal G, Picazo S, Müntener O, Karner GD, Johnson C, et al. Influence of the architecture of magma-poor hyperextended rifted margins on orogens produced by the closure of narrow versus wide oceans. Geosphere. 2017;13:559-576. DOI: 10.1130/GES01363.1
  44. 44. Chenin P, Picazo S, Jammes S, Manatschal G, Müntener O, Karner GD. Potential role of lithospheric mantle composition in the Wilson cycle: A North Atlantic perspective. In: Wilson R W, Wilson RW, editors. Fifty Years of the Wilson Cycle Concept in Plate Tectonics. London: Geological Society; 2018. p. 470. DOI: 10.1144/SP470.10
  45. 45. Perez-Gussinyé M, Reston TJ. Rheological evolution during extension at non-volcanic rifted margins: Onset of serpentinization and development of detachments leading to continental breakup. Journal of Geophysical Research. 2001;106:3961-3975. DOI: 10.1029/2000JB900325
  46. 46. Sutra E, Manatschal G, Mohn G, Unternehr P. Quantification and restoration of extensional deformation along the Western Iberia and Newfoundland rifted margins. Geochemistry, Geophysics, Geosystems. 2013;14:2575-2597. DOI: 10.1002/ggge.20135
  47. 47. Doré T, Lundin E. Hyperextended continental margins – knowns and unknowns. Geology. 2015;43:95-96. DOI: 10.1130/focus012015.1
  48. 48. Manatschal G. Fluid-and reaction-assisted low angle normal faulting: Evidence from rift-related brittle fault rocks in the Alps (Err Nappe, eastern Switzerland). Journal of Structural Geology. 1999;21:777-793. DOI: 10.1016/S0191-8141(99)00069-3
  49. 49. Manatschal G. New models for evolution of magma-poor rifted margins based on a review of data and concepts from West Iberia and the Alps. International Journal of Earth Sciences. 2004;93:432-466. DOI: 10.1007/S00531-004-0394-7
  50. 50. Pinto VHG, Manatschal G, Karpoff AM, Viana A. Tracing mantle-reacted fluids in magma-poor rifted margins: The example of Alpine Tethyan rifted margins. Geochemistry, Geophysics, Geosystems. 2015;16:3271-3308. DOI: 10.1002/2015GC005830
  51. 51. Hess H. Serpentines, orogeny, and epeirogeny. In: Poldervaart A, editor. Crust of the Earth: A Symposium, Geological Society of America Special Papers. Vol. 62. Boulder, USA: The Geological Society of America; 1955. pp. 391-408. DOI: 10.1130/SPE62-P391
  52. 52. Christensen NI. Composition and evolution of the oceanic crust. Marine Geology. 1970;8:139-154. DOI: 10.1016/0025-3227(70)90002-2
  53. 53. Früh-Green GL, Connolly JA, Plas A, Kelley DS, Grobéty B. Serpentinization of oceanic peridotites: Implications for geochemical cycles and biological activity. In: Wilcock WS, Delong EF, Kelley DS, Baross JA, Craig Cary S, editors. The Subseafloor Biosphere at Mid-Ocean Ridges. Vol. 144. Washington, DC: American Geophysical Union, Geophysical Monographs; 2004. pp. 119-136. DOI: 10.1029/144GM08
  54. 54. Picazo S, Cannat M, Delacour A, Escartín J, Roumé-Jon S, Silantyev S. Deformation associated with the denudation of mantle-derived rocks at the Mid-Atlantic Ridge 13-15°N: The role of magmatic injections and hydrothermal alteration. Geochemistry, Geophysics, Geosystems. 2012;13:Q04G09. DOI: 10.1029/2012GC004121
  55. 55. Latin D, White N. Generating melt during lithospheric extension: Pure shear vs. simple shear. Geology. 1990;18:327-331. DOI: 10.1130/0091-7613(1990)018<0327:GMDLEP>2.3.CO;2
  56. 56. Müntener O, Manatschal G, Desmurs L, Pettke T. Plagioclase peridotites in ocean–continent transitions: Refertilized mantle domains generated by melt stagnation in the shallow mantle lithosphere. Journal of Petrology. 2010;51:255-294. DOI: 10.1093/petrology/egp087
  57. 57. O’reilly B, Hauser F, Ravaut C, Shannon P, Readman P. Crustal thinning, mantle exhumation and serpentinization in the Porcupine Basin, offshore Ireland: Evidence from wide-angle seismic data. Journal of the Geological Society. 2006;163:775-787. DOI: 10.1144/0016-76492005-07
  58. 58. Afilhado A, Matias L, Shiobara H, Hirn A, Mendesvictor L, Shimamura H: From unthinned continent to ocean: The deep structure of the West Iberia passive continental margin at 38°N. Tectonophysics 2008; 458: 9-50. https://doi.org/10.1016/j.tecto.2008.03.002
  59. 59. Lau KWH, Louden KE, Funck T, Tucholke BE, Holbrook WS, Hopper JR, et al. Crustal structure across the Grand Banks–Newfoundland Basin continental margin – I. Results from a seismic refraction profile. Geophysical Journal International. 2006;167:127-156. DOI: 10.1111/j.1365-246X.2006.02988.x
  60. 60. Funck T, Hopper JR, Larsen HC, Louden KE, Tucholke BE, Holbrook WS. Crustal structure of the ocean-continent transition at Flemish Cap: Seismic refraction results. Journal of Geophysical Research. 2003;108:2531. DOI: 10.1029/2003JB002434
  61. 61. Boillot G, Capdevila R. The Pyrenees: Subduction and collision? Earth and Planetary Science Letters. 1977;35:151-160. DOI: 10.1016/0012-821X(77)90038-3
  62. 62. Sengör AMC. Orogenic architecture as a guide to size of ocean lost in collisional mountain belts. Bulletin of the Technical University of Istanbul. 1991;44:43-74
  63. 63. Anderson DL. Speculations on the nature and cause of mantle heterogeneity. Tectonophysics. 2006;416:7-22. DOI: 10.1016/j.tecto.2005.07.011
  64. 64. Costa S, Rey P. Lower crustal rejuvenation and growth during post-thickening collapse: Insights from a crustal cross section through a Variscan metamorphic core complex. Geology. 1995;23:905-908. DOI: 10.1130/0091-7613(1995)023<0905:LCRAGD>2.3.CO;2
  65. 65. Vanderhaeghe O, Burg J-P, Teyssier C. Exhumation of migmatites in two collapsed orogens: Canadian Cordillera and French Variscides. In: Ring U, Brandon MT, Lister GS, Willett SD, editors. Exhumation Processes: Normal Faulting, Ductile Flow and Erosion. Vol. 154. London: Geological Society; 1999. pp. 181-204. DOI: 10.1144/GSL.SP.1999.154.01.08
  66. 66. Rey P, Vanderhaeghe O, Teyssier C. Gravitational collapse of the continental crust: Definition, regimes and modes. Tectonophysics. 2001;342:435-449. DOI: 10.1016/S0040-1951(01)00174-3
  67. 67. Petri B, Mohn G, Štípská P, Schulmann K, Manatschal G. The Sondalo gabbro contact aureole (Campo unit, Eastern Alps): Implications for midcrustal mafic magma emplacement. Contributions to Mineralogy and Petrology. 2016;171:1-21. DOI: 10.1007/s00410-016-1263-7
  68. 68. Iwamori H. Transportation of H2O and melting in subduction zones. Earth and Planetary Science Letters. 1998;160:65-80. DOI: 10.1016/S0012-821X(98)00080-6
  69. 69. Sisson TW, Bronto S. Evidence for pressure release melting beneath magmatic arcs from basalt at Galunggung, Indonesia. Nature. 1998;391:883-886. DOI: 10.1038/36087
  70. 70. Jagoutz O, Müntener O, Schmidt MW, Burg J-P. The roles of flux- and decompression melting and their respective fractionation lines for continental crust formation: Evidence from the Kohistan arc. Earth and Planetary Science Letters. 2011;303:25-36. DOI: 10.1016/j.epsl.2010.12.017
  71. 71. Uyeda S. Subduction zones and back arc basins – a review. Geologische Rundschau. 1981;70:552-569. DOI: 10.1007/BF01822135
  72. 72. Martinez F, Taylor B. Mantle wedge control on back-arc crustal accretion. Nature. 2002;416:417-420. DOI: 10.1038/416417a
  73. 73. Chenin P, Manatschal G, Lavier LL, Erratt D. Assessing the impact of orogenic inheritance on the architecture, timing, and magmatic budget of the North Atlantic rift system: A mapping approach. Journal of the Geological Society. 2015;172:711-720. DOI: 10.1144/jgs2014-139
  74. 74. Stern R. Subduction initiation: Spontaneous and induced. Earth and Planetary Science Letters. 2004;226:275-292. DOI: 10.1016/j.epsl.2004.08.007
  75. 75. Stern RJ, Gerya T. Subduction initiation in nature and models: A review. Tectonophysics. 2018;746:173-198. DOI: 10.1016/j.tecto.2017.10.014
  76. 76. Crameri F, Magni V, Domeier M, Shephard GE, Chotalia K, Cooper G, et al. A transdisciplinary and community-driven database to unravel subduction zone initiation. Nature Communications. 2020;11:3750. DOI: 10.1038/s41467-020-17522-9
  77. 77. Mueller S, Phillips RJ. On the initiation of subduction. Journal of Geophysical Research. 1991;96:651-665. DOI: 10.1029/90JB02237
  78. 78. Candioti L, Duretz T, Schmalholz SM. Horizontal force required for subduction initiation at passive margins with constraints from slab detachment. Frontiers in Earth Science. 2022;10:785418. DOI: 10.3389/feart.2022.785418
  79. 79. Cloetingh S, Wortel R, Vlaar NJ. On the initiation of subduction zones. Pure and Applied Geophysics. 1989;129:7-25. DOI: 10.1007/BF00874622
  80. 80. Lallemand S, Arcay D. Subduction initiation from the earliest stages to self-sustained subduction: Insights from the analysis of 70 Cenozoic sites. Earth Science Reviews. 2021;221:103779. DOI: 10.1016/j.earscirev.2021.103779
  81. 81. McKenzie DP. The initiation of trenches: A finite amplitude instability. In: Talwani M, Pitman WC III, editors. Island Arcs, Deep Sea Trenches and Back-Arc Basins, Volume 1. Washington, DC: American Geophysical Union; 1977. pp. 57-61. DOI: 10.1029/me001p0057
  82. 82. Zhong X, Li Z-H. Subduction initiation at passive continental margins: A review based on numerical studies. Solid Earth Sciences. 2021;6:249-267. DOI: 10.1016/j.sesci.2021.06.001
  83. 83. Shuck B, Gulick SPS, Van Avendonk HJA, Gurnis M, Sutherland R, Stock J, et al. Stress transition from horizontal to vertical forces during subduction initiation. Nature Geoscience. 2022;15:149-155. DOI: 10.1038/s41561-021-00880-4
  84. 84. White SH, Knipe RJ. Transformation- and reaction-enhanced ductility in rocks. Journal of the Geological Society. 1978;135:513-516. DOI: 10.1144/gsjgs.135.5.0513
  85. 85. Regenauer-Lieb K, Yuen DA, Branlund J. The initiation of subduction: Criticality by addition of water? Science. 2001;294:578-580. DOI: 10.1126/science.1063891
  86. 86. Dymkova D, Gerya T. Porous fluid flow enables oceanic subduction initiation on Earth. Geophysical Research Letters. 2013;40:5671-5676. DOI: 10.1002/2013GL057798
  87. 87. Montési LG. Fabric development as the key for forming ductile shear zones and enabling plate tectonics. Journal of Structural Geology. 2013;50:254-266. DOI: 10.1016/j.jsg.2012.12.011
  88. 88. Bercovici D, Ricard Y. Mechanisms for the generation of plate tectonics by two-phase grain-damage and pinning. Physics of the Earth and Planetary Interiors. 2012;202-203:27-55. DOI: 10.1016/j.pepi.2012.05.003
  89. 89. Mulyukova E, Bercovici D. Collapse of passive margins by lithospheric damage and plunging grain size. Earth and Planetary Science Letters. 2018;484:341-352. DOI: 10.1016/j.epsl.2017.12.022
  90. 90. Yuen DA, Fleitout L, Schubert G, Froidevaux C. Shear deformation zones along major transform faults and subducting slabs. Geophysical Journal International. 1978;54:93-119. DOI: 10.1111/j.1365-246X.1978.tb06758.x
  91. 91. Regenauer-Lieb K, Yuen DA. Rapid conversion of elastic energy into plastic shear heating during incipient necking of the lithosphere. Geophysical Research Letters. 1998;25:2737-2740. DOI: 10.1029/98GL02056
  92. 92. Kiss D, Podladchikov Y, Duretz T, Schmalholz SM. Spontaneous generation of ductile shear zones by thermal softening: Localization criterion, 1D to 3D modelling and application to the lithosphere. Earth and Planetary Science Letters. 2019;519:284-296. DOI: 10.1016/j.epsl.2019.05.026ff
  93. 93. Thielmann M, Kaus BJP. Shear heating induced lithospheric-scale localization: Does it result in subduction? Earth and Planetary Science Letters. 2012;359-360:1-13. DOI: 10.1016/j.epsl.2012.10.002
  94. 94. Jaquet Y, Schmalholz SM. Spontaneous ductile crustal shear zone formation by thermal softening and related stress, temperature and strain rate evolution. Tectonophysics. 2018;746:384-397. DOI: 10.1016/j.tecto.2017.01.012
  95. 95. Kiss D, Candioti LG, Duretz T, Schmalholz SM. Thermal softening induced subduction initiation at a passive margin. Geophysical Journal International. 2020;220:2068-2073. DOI: 10.1093/gji/ggz572
  96. 96. Auzemery A, Willingshofer E, Yamato P, Duretz T, Beekman F. Kinematic boundary conditions favouring subduction initiation at passive margins over subduction at mid-oceanic ridges. Frontiers in Earth Science. 2021;9:765893. DOI: 10.3389/feart.2021.765893
  97. 97. Auzemery A, Willingshofer E, Yamato P, Duretz T, Sokoutis D. Strain localization mechanisms for subduction initiation at passive margins. Global and Planetary Change. 2020;195:103323. DOI: 10.1016/j.gloplacha.2020.103323
  98. 98. Ulmer P, Trommsdorff V. Serpentine stability to mantle depths and subduction-related magmatism. Science. 1995;268:858-861. DOI: 10.1126/science. 268.5212.858
  99. 99. Escartm J, Hirth G, Evans B. Strength of slightly serpentinized peridotites: Implications for the tectonics of oceanic lithosphere. Geology. 2001;29:1023-1026. DOI: 10.1130/0091-7613(2001)029<1023:SOSSPI>2.0.CO;2
  100. 100. Hilairet N, Reynard B, Wang Y, Daniel I, Merkel S, Nishiyama N, et al. High pressure creep of serpentine, interseismic deformation, and initiation of subduction. Science. 2007;318:1910-1913. DOI: 10.1126/science.1148494
  101. 101. Zhong X, Li Z-H. Forced subduction initiation at passive continental margins: Velocity-driven versus stress-driven. Geophysical Research Letters. 2019;46:11054-11064. DOI: 10.1029/2019GL084022
  102. 102. Shervais JW. Birth, death, and resurrection: The life cycle of suprasubduction zone ophiolites. Geochemistry, Geophysics, Geosystems. 2001;2:1010. DOI: 10.1029/2000GC000080
  103. 103. Maffione M, van Hinsbergen DJJ, de Gelder GINO, van der Goes FC, Morris A. Kinematics of late cretaceous subduction initiation in the neo-Tethys ocean reconstructed from ophiolites of Turkey, Cyprus, and Syria. Journal of Geophysical Research. 2017;122:3953-3976. DOI: 10.1002/2016JB013821
  104. 104. Plank T, Langmuir CH. An evaluation of the global variations in the major element chemistry of arc basalts. Earth and Planetary Science Letters. 1988;90:349-370. DOI: 10.1016/0012-821X(88)90135-5
  105. 105. Lee C-TA, Luffi P, Plank T, Dalton H, Leeman WP. Constraints on the depths and temperatures of basaltic magma generation on earth and other terrestrial planets using new thermobarometers for mafic magmas. Earth and Planetary Science Letters. 2009;279:20-33. DOI: 10.1016/j.epsl.2008.12.020
  106. 106. Grove TL, Till CB, Krawczynski MJ. The role of H2O in subduction zone magmatism. Annual Review of Earth and Planetary Sciences. 2012;40:413-439. DOI: 10.1146/annurev-earth-042711-105310
  107. 107. Perrin A, Goes S, Prytulak J, Davies DR, Wilson C, Kramer S. Reconciling mantle wedge thermal structure with arc lava thermobarometric determinations in oceanic subduction zones. Geochemistry, Geophysics, Geosystems. 2016;17:4105-4127. DOI: 10.1002/2016GC006527
  108. 108. Khudeir AA, Paquette JL, Nicholson K, Johansson Å, Rooney TO, El-Fadly MA, et al. On the cratonization of the Arabian-Nubian Shield: Constraints from granite gneisses in south eastern desert, Egypt. Geoscience Frontiers. 2021;12:101148. DOI: 10.1016/j.gsf.2021.101148
  109. 109. Stern RJ. Evidence from ophiolites, blueschists, and ultrahigh-pressure metamorphic terranes that the modern episode of subduction tectonics began in Neoproterozoic time. Geology. 2005;33:557-560. DOI: 10.1130/G21365.1
  110. 110. Crawford AJ, Meffre S, Symonds PA. 120 to 0 Ma tectonic evolution of the Southwest Pacific and analogous geological evolution of the 600 to 220 Ma Tasman fold belt. In: Hills RR, Muller RD, editors. Evolution and Dynamics of the Australian Plate. Geological Society of America Special Paper; Vol. 372. Boulder, USA: The Geological Society of America; 2003; Paper 372. pp. 383-404. DOI: 10.1130/0-8137-2372-8.383
  111. 111. Schellart WP, Lister GS, Toy VG. A late Cretaceous and Cenozoic reconstruction of the Southwest Pacific region: Tectonics controlled by subduction and slab rollback processes. Earth Science Reviews. 2006;76:191-233. DOI: 10.1016/j.earscirev.2006.01.002
  112. 112. Nicholson KN, Maurizot PP, Black PM, Picard C, Simonetti AA, Stewart A, et al. Geochemistry and age of the Nouméa Basin lavas, New Caledonia; Evidence for Cretaceous subduction beneath the eastern Gondwana margin. Lithos. 2011;125:659-674. DOI: 10.1016/j.lithos.2011.03.018
  113. 113. Sevin B, Maurizot P, Cluzel D, Tournadour E, Etienne S, Folcher N, et al. Post-obduction evolution of New Caledonia. In: Maurizot P, Mortimer N, editors. New Caledonia: Geology, Geodynamic Evolution and Mineral Resources. Vol. 51. London: Geological Society; 2020. pp. 147-188. DOI: 10.1144/M51-2018-74
  114. 114. Falvey DA, Mutter JC. Regional plate tectonics and the evolution of Australia’s passive continental margin. BMR Journal of Australian Geology and Geophysics. 1981;6:1-29
  115. 115. Régnier M. Lateral variation of upper mantle structure beneath New Caledonia determined from P-wave receiver function: Evidence for a fossil subduction zone. Geophysical Journal. 1988;95:561-577. DOI: 10.1111/j.1365-246X.1988.tb06704.x
  116. 116. Cluzel D, Bosch D, Paquette J, Lemennicier Y, Montjoi P, Menot R. Late Oligocene post-obduction granitoids of New Caledonia: A case for reactivated subduction and slab break-off. The Island Arc. 2005;14:254-271. DOI: 10.1111/j.1440-1738.2005.00470.x
  117. 117. Putirka K. Excess temperatures at ocean islands: Implications for mantle layering and convection. Geology. 2008;36:283-286. DOI: 10.1130/G24615A.1
  118. 118. Tatsumi Y. Some Constraints on Arc Magma Genesis in Inside the Subduction Factory. Vol. 138. Washington, D.C.: American Geophysical Union, Geophysical Monograph; 2003. pp. 277-292. DOI: 10.1029/138GM12
  119. 119. Stern RJ. Subduction zones. Reviews of Geophysics. 2002;40:3.1-3.38. DOI: 10.1029/2001RG000108
  120. 120. van Keken PE, Hacker BR, Syracuse EM, Abers GA. Subduction factory: 4. Depth-dependent flux of H2O from subducting slabs worldwide. Journal of Geophysical Research. 2011;116:B01401 DOI: 10.1029/2010JB007922
  121. 121. Cluzel D, Chiron D, Courme MD. Discordance de l’Eoce’ne supérieur et e’ve’nements pré-obduction en Nouvelle-Cale’donie (Pacifique sud-ouest): Compte Rendus de l’Académie des Sciences, Série II. Sciences de la Terre et des Planètes. 1998;327:485-449
  122. 122. Paquette J-L, Cluzel D. U–Pb zircon dating of post-obduction volcanic-arc granitoids and a granulite-facies xenolith from New Caledonia, Inference on Southwest Pacific geodynamic models. International Journal of Earth Sciences. 2007;96:613-622. DOI: 10.1007/s00531-006-0127-1
  123. 123. Yumul, Jr GP, Dimalanta CB, Marquez EJ, Queaño KL. Onland signatures of the Palawan microcontinental block and the Philippine mobile belt collision and crustal growth process: A review. Journal of Asia Earth Sciences. 2009;34:610-623. DOI: 10.1016/j.jseaes.2008.10.002
  124. 124. Yumul, Jr GP, Armada LT, Gabo-Ratiob JAS, Dimalanta CB, Austria RSP. Subduction with arrested volcanism: Compressional regime in volcanic arc gap formation along east Mindanao, Philippines. Journal of Asia Earth Sciences. 2020;4:100030. DOI: 10.1016/j.jaesx.2020.100030
  125. 125. Pearce JA, Peate DW. Tectonic implications of the composition of volcanic arc magmas. Annual Review of Earth and Planetary Sciences. 1995;23:251-285. DOI: 10.1146/annurev.ea.23.050195.001343
  126. 126. Shor GG, Kirk HK, Maynard GL. Crustal structure of the Melanesian area. Journal of Geophysical Research. 1971;76:2562-2586. DOI: 10.1029/JB076i011p02562
  127. 127. Dubois J, Aubertin RA, Louis J, Guillaume R, Launay J, Montadert L. Continental margins near New Caledonia. In: Burk CA, Drake CL, editors. The Geology of Continental Margins. New York: Springer Science and Business Media; 1974. pp. 521-1029
  128. 128. Kinzler RJ. Melting of mantle peridotite at pressures approaching the spinel to garnet transition: Application to mid-ocean ridge basalt petrogenesis. Journal of Geophysical Research. 1997;102:853-874. DOI: 10.1029/96JB00988
  129. 129. Faul UH. Melt retention and segregation beneath mid-ocean ridges. Nature. 2001;410:920-923. DOI: 10.1038/35073556
  130. 130. Zhu W, Hirth G. A network model for permeability in partially molten rocks. Earth and Planetary Science Letters. 2003;212:407-416. DOI: 10.1016/S0012-821X(03)00264-4
  131. 131. Rondenay S, Montési LGJ, Abers GA. New geophysical insight into the origin of the Denali volcanic gap. Geophysical Journal International. 2010;182:613-630. DOI: 10.1111/j.1365-246X.2010.04659.x
  132. 132. Korenaga J, Kelemen PB. Origin of gabbro sills in the Moho transition zone of the Oman ophiolite: Implications for magma transport in the oceanic lower crust. Journal of Geophysical Research. 1997;102:27729-27749. DOI: 10.1029/97JB02604
  133. 133. Sparks DW, Parmentier EM. Melt extraction from the mantle beneath spreading centers. Earth and Planetary Science Letters. 1991;105:368-377. DOI: 10.1016/0012-821X(91)90178-K
  134. 134. Spiegelman M. Physics of melt extraction: Theory, implications and applications. Philosophical Transactions of the Royal Society of London: Physical Sciences and Engineering. 1993;342:23-41. DOI: 10.1098/rsta.1993.0002
  135. 135. Katz RF, Spiegelman M, Langmuir CH. A new parameterization of hydrous mantle melting. Geochemistry, Geophysics, Geosystems. 2003;4:1073. DOI: 10.1029/2002GC000433
  136. 136. Lagabrielle Y, Labaume P, de Saint BM. Mantle exhumation, crustal denudation, and gravity tectonics during Cretaceous rifting in the Pyrenean realm (SW Europe): Insights from the geological setting of the lherzolite bodies. Tectonics. 2010;29:TC4012. DOI: 10.1029/2009TC002588
  137. 137. Tugend J, Manatschal G, Kusznir N. Spatial and temporal evolution of hyperextended rift systems: Implication for the nature, kinematics and timing of the Iberian–European plate boundary. Geology. 2015;43:15-18. DOI: 10.1130/G36072
  138. 138. Kröner A, Romer RL. Two plates-many subduction zones: The Variscan orogeny reconsidered. Gondwana Research. 2013;24:298-329. DOI: 10.1016/j.gr.2013.03.001
  139. 139. Regorda A, Lardeaux J-M, Roda M, Marotta AM, Spalla MI. How many subductions in the Variscan orogeny? Insights from numerical models. Geoscience Frontiers. 2020;11:1025-1052. DOI: 10.1016/j.gsf.2019.10.005
  140. 140. Beltrando M, Manatschal G, Mohn G, Dal Piaz GV, Vitale Brovarone A, Masini E. Recognizing remnants of magma-poor rifted margins in high-pressure orogenic belts: The Alpine case study. Earth Science Reviews. 2014;131:88-115. DOI: 10.1016/j.earscirev.2014.01.001
  141. 141. Rosenbaum G, Lister G S: The Western Alps from the Jurassic to Oligocene: Spatio-temporal constraints and evolutionary reconstructions: Earth Science Reviews 2005;69:281-306. DOI: 10.1016/j.earscirev.2004.10.001
  142. 142. Handy KR, Schmid SM, Bousquet R, Kissling E, Bernoulli D. Reconciling plate-tectonic reconstructions of Alpine Tethys with the geological–geophysical record of spreading and subduction in the Alps. Earth Science Reviews. 2010;102:121-158. DOI: 10.1016/j.earscirev.2010.06.002
  143. 143. Le Breton E, Brune S, Ustaszewski K, Zahirovic S, Seton M, Müller RD. Kinematics and extent of the Piemont-Liguria basin: Implications for subduction processes in the Alps. Solid Earth. 2021;12:885-913. DOI: 10.5194/se-12-885-2021
  144. 144. Schenker FL, Schmalholz SM, Moulas E, Pleuger J, Baumgartner LP, Podladchikov Y, et al. Current challenges for explaining (ultra)high-pressure tectonism in the Pennine domain of the Central and Western Alps. Journal of Metamorphic Geology. 2015;33:869-886. DOI: 10.1111/jmg.12143
  145. 145. Manzotti P, Bosse V, Pitra P, Robyr M, Schiavi F, Ballèvre M. Exhumation rates in the Gran Paradiso Massif (Western Alps) constrained by in situ U–Th–Pb dating of accessory phases (monazite, allanite and xenotime). Contributions to Mineralogy and Petrology. 2014;173:1-28. DOI: 10.1007/s00015-014-0172-x
  146. 146. Rubatto D, Gebauer D, Fanning M. Jurassic formation and Eocene subduction of the Zermatt-Saas-Fee ophiolites: Implications for the geodynamic evolution of the Central and Western Alps. Contributions to Mineralogy and Petrology. 1998;132:269-287. DOI: 10.1007/s004100050421
  147. 147. Manzotti P, Ballèvre M, Zucali M, Robyr M, Engi M. The tectonometamorphic evolution of the Sesia-Dent Blanche Nappes (Internal Western Alps): Review and synthesis. Swiss Journal of Geosciences. 2014;107:309-336. DOI: 10.1007/s00015-014-0172
  148. 148. Schmid SM, Kissling E, Diehl T, van Hinsbergen DJJ, Molli G. Ivrea Mantle Wedge, Arc of the Western Alps, and Kinematic Evolution of the Alps-Apennines Orogenic System. Swiss Journal of Geosciences. 2017;110:581-612. DOI: 10.1007/s00015-016-0237-0
  149. 149. Kästle ED, Rosenberg C, Boschi L, Bellahsen N, Meier T, El-Sharkawy A. Slab break-offs in the alpine subduction zone. International Journal of Earth Sciences. 2020;109:587-603. DOI: 10.1007/s00531-020-01821-z
  150. 150. Vissers RL, van Hinsbergen DJ, van der Meer DG, Spakman W. Cretaceous slab break-off in the Pyrenees: Iberian plate kinematics in paleomagnetic and mantle reference frames. Gondwana Research. 2016;34:49-59. DOI: 10.1016/j.gr.2016.03.006
  151. 151. Ziegler PA. Geological Atlas of Western and Central Europe. Vol. 121. The Hague, Shell Internationale Petroleum Maatschappij; 1982. pp. 371-372. DOI: 10.1017/S0016756800029344
  152. 152. Carey WS. The orocline concept in Geotectonics. Papers and Proceedings of the Royal Society of Tasmania. 1958;89:255-288
  153. 153. Van der Voo R, Spakman W, Bijwaard H. Tethyan subducted slabs under India. Earth and Planetary Science Letters. 1999;171:7-20
  154. 154. Le Pichon X, Sibuet JC. Western extension of boundary between European and Iberian plates during the Pyrenean orogeny. Earth and Planetary Science Letters. 1971;12:83-88. DOI: 10.1016/0012-821X(71)90058-6
  155. 155. Srivastava SP, Roest WR, Kovacs LC, Oakey G, Lévesque S, Verhoef J, et al. Motion of Iberia since the late Jurassic: Results from detailed aeromagnetic measurements in the Newfoundland basin. Tectonophysics. 1990;184:229-260
  156. 156. Srivastava SP, Sibuet JC, Cande S, Roest WR, Reid ID. Magnetic evidence for slow seafloor spreading during the formation of the Newfoundland and Iberian margins. Earth and Planetary Science Letters. 2000;182:61-76. DOI: 10.1016/S0012-821X(00)00231-4
  157. 157. Sibuet J-C, Srivastava SP, Spakman W. Pyrenean orogeny and plate kinematics. Geophysical Research. 2004;109:B08104. DOI: 10.1029/2003JB002514
  158. 158. Vissers RLM, Th MP. Mesozoic rotation of Iberia: Subduction in the Pyrenees? Earth Science Reviews. 2012;110:93-110. DOI: 10.1016/j.earscirev. 2011.11.001
  159. 159. Beaumont C, Muñoz JA, Hamilton J, Fullsack P. Factors controlling the alpine evolution of the central Pyrenees inferred from a comparison of observations and geodynamical models. Journal of Geophysical Research. 2000;105:8121-8145. DOI: 10.1029/1999JB900390
  160. 160. Roure F, Choukroune P, Berástegui X, Muñoz JA, Villien A, Matheron P, et al. ECORS deep seismic data and balanced cross sections: Geometric constraints on the evolution of the Pyrenees. Tectonics. 1989;8:41-50. DOI: 10.1029/TC008i001p00041
  161. 161. Souriau A, Chevrot S, Olivera C. A new tomographic image of the Pyrenean lithosphere from teleseismic data. Tectonophysics. 2008;460:206-214. DOI: 10.1016/j.tecto.2008.08.014
  162. 162. Bronner A, Sauter D, Manatschal G, Péron-Pinvidic G, Munschy M. Reply to problematic plate reconstruction by Tucholke B E, Sibuet J-C. Nature Geoscience. 2012;5:677
  163. 163. Clerc C, Lahfid A, Monié P, Lagabrielle Y, Chopin C, Poujol M, et al. High-temperature metamorphism during extreme thinning of the continental crust: A reappraisal of the North Pyrenean passive paleomargin. Solid Earth. 2015;6:643-668. DOI: 10.5194/se-6-643-2015
  164. 164. Mc Carthy A, Tugend J, Mohn G. Formation of the Alpine orogen by amagmatic convergence and assembly of previously rifted lithosphere. Elements. 2021;17:29-34. DOI: 10.2138/gselements.17.1.29
  165. 165. Michard A, Ouanaimi H, Chr H, Soulaimani A, Baidder L. Reply to the comment by Michard et al. on “Tectonic relationships of Southwest Iberia with the allochthons of Northwest Iberia and the Moroccan Variscides” by J F Simancas et al. (Comptes Rendus Geoscience 2009;341:103-113). Comptes Rendus Geoscience. 2010;342:170-174. DOI: 10.1016/j.crte.2010.01.008
  166. 166. Matte P. The Variscan collage and orogeny (480-290 Ma) and the tectonic definition of the Armorica microplate: A review. Terra Nova. 2001;13:122-128. DOI: 10.1046/j.1365-3121.2001.00327.x
  167. 167. Franke W, Cocks LRM, Torsvik TH. The Palaeozoic Variscan oceans revisited. Gondwana Research. 2017;48:257-284. DOI: 10.1016/j.gr.2017.03.005
  168. 168. Franke W. The mid-European segment of the Variscides: Tectonostratigraphic units, terrane boundaries and plate tectonic evolution. In: Franke W, Haak V, Oncken O, Tanner D, editors. Orogenic Processes: Quantification and Modelling in the Variscan belt. Vol. 179. London: Geological Society; 2000. pp. 35-61. DOI: 10.1144/GSL.SP.2000.179.01.05
  169. 169. Winchester JA, Pharaoh TC, Verniers J. Palaeozoic amalgamation of Central Europe. Vol. 201: 343. London: Geological Society; 2002. DOI: 10.1144/gsl.sp.2002.201
  170. 170. Kröner U, Hahn T, Romer RL, Linnemann U. The Variscan orogeny in the Saxo-Thuringian Zone – heterogeneous overprint of Cadomian/Paleozoic Peri-Gondwana crust. In: Linnemann UR, Nance D, Kraft P, editors. The Evolution of the Rheic Ocean: From Avalonian-Cadomian Active Margin to Alleghenian-Variscan Collision. Geological Society of America Special Paper; Vol. 423. Boulder, USA: The Geological Society of America; 2007. pp. 153-172. DOI: 10.1130/SPE423
  171. 171. Cogné J. The Cadomian Orogeny and its influence on the Variscan evolution of western Europe. In: D’Lemos RS, Strachan RA, Topley CG, editors. The Cadomian Orogeny. Vol. 51. London: Geological Society; 1990. pp. 305-311. DOI: 10.1144/GSL.SP.1990.051.01.19
  172. 172. Crowley QG, Floyd PA, Winchester GA, Franke W, Holland G. Early Palaeozoic rift-related magmatism in Variscan Europe: Fragmentation of the Armorican Terrane Assemblage. Terra Nova. 2000;12:171-180. DOI: 10.1046/j.1365-3121.2000.00290.x
  173. 173. Floyd PA, Winchester J, Seston R, Kryza R, Crowley QG. Review of geochemical variation in Lower Palaeozoic metabasites from the NE Bohemian Massif: Intracratonic rifting and plume-ridge interaction. In: Franke W, Haak V, Oncken O, Tanner D, editors. Orogenic Processes: Quantification and Modelling in the Variscan Belt. Vol. 179. London: Geological Society; 2000. pp. 155-174. DOI: 10.1144/GSL.SP.2000.179.01.1
  174. 174. Nowak I, Żelaźniewicz A, Dörr W, Franke W, Larionov AN. The Izera metabasites, West Sudetes, Poland: Geologic and isotopic U–Pb evidence of Devonian extension in the Saxothuringian Terrane. Lithos. 2011;126:435-454. DOI: 10.1016/j.lithos.2011.07.006

Written By

Mohamed A. Abu El-Rus, Ali A. Khudier, Sadeq Hamid and Hassan Abbas

Submitted: 24 November 2022 Reviewed: 08 December 2022 Published: 06 January 2023