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

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).

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.

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].

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].

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].

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].

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, 170, 171] 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).

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.

Acknowledgments

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

Conflict of interest

All figures are used with permission.