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Review Article
Crustal thickness controlled by plate tectonics: A review of crust–mantle interaction
processes illustrated by European examples
Irina M. Artemieva
a,
⁎, Rolf Meissner
b
a
University of Copenhagen, Denmark
b
Kiel University, Germany
abstractarticle info
Article history:
Received 16 February 2011
Received in revised form 22 December 2011
Accepted 29 December 2011
Available online 8 January 2012
Keywords:
Seismic reflectivity
Delamination
Extension
Magmatism
Variscides
Eclogitization
The continental crust on Earth cannot be extracted directly from the mantle, and the primary crust extracted
directly from an early magma ocean is not preserved on Earth. We review geophysical and geochemical aspects
of global crust–mantle material exchange processes and examine the processes which, on one side, form and
transformthe continental crustand, on the other side, chemically modifythe mantle residue fromwhich the con-
tinental crust has been extracted. Major mechanisms that provide crust–mantle material exchange are oceanic
and continental subduction, lithosphere delamination, and mafic magmatism. While both subduction and de-
lamination recycle crustal material into the mantle, mafic magmatism transports mantle material upward and
participates in growth ofnew oceanic and continental crusts andsignificant structuraland chemical modification
of the latter. We discuss the role of basalt/gabbro–eclogite phase transition in crustal evolution and the links
between lithosphere recycling, mafic magmatism, and crustal underplating. We advocate that plate tectonics
processes, together with basalt/gabbro–eclogite transition, limit crustal thickness worldwide by providing effec-
tive mechanisms of crustal (lithosphere) recycling.
The processesof crust–mantle interaction have created very dissimilar crustal stylesin Europe, as seen by its seis-
mic structure, crustal thickness, and average seismic velocities in the basement. Our special focus is on processes
responsible for the formation of the thin crust of central and western Europe, which was largely formed during
the Variscan (430–280 Ma) orogeny but has the present structure of an “extended”crust, similar to that of the
Basin and Range province in western USA. Major geophysical characteristics of the Variscan lithosphere are dis-
cussed within the frame of possible sequences of crust–mantle material exchange mechanisms during and after
main orogenic events in the European Variscides.
© 2012 Elsevier B.V. All rights reserved.
Contents
1. Introduction ...............................................................19
1.1. Primary crusts and planetary-scale differentiation ..........................................19
1.2. Primary crusts and plate tectonics on terrestrial planets .......................................20
1.2.1. Earth ........................................................... 20
1.2.2. Venus .......................................................... 20
1.2.3. Mars ........................................................... 21
1.2.4. Moon and Mercury ....................................................21
1.3. Geodynamic controls of crustal volume ...............................................21
2. Crust–mantle exchange processes: a global view ..............................................22
2.1. Crust–mantle compositional balance: a planetary-scale geochemical perspective ............................22
2.2. Making new continental crust ...................................................23
2.2.1. Growth of continental crust: convergent margins versus intraplate settings ..........................23
2.2.2. Growth rate of the continental crust ............................................24
2.3. Mafic magmatism: volcanism and intrusions .............................................24
2.3.1. Geodynamic causes of mafic volcanism ...........................................24
2.3.2. The shape of magmatic intrusions .............................................24
Tectonophysics 530-531 (2012) 18–49
⁎Corresponding author. Tel.: + 45 50882438; fax: +45 35322501.
E-mail address: irina@geo.ku.dk (I.M. Artemieva).
0040-1951/$ –see front matter © 2012 Elsevier B.V. All rights reserved.
doi:10.1016/j.tecto.2011.12.037
Contents lists available at SciVerse ScienceDirect
Tectonophysics
journal homepage: www.elsevier.com/locate/tecto
2.4. Subduction ............................................................ 25
2.4.1. Precambrian subduction .................................................. 25
2.4.2. Modern oceanic subduction ................................................ 27
2.4.3. Continental collision and subduction ............................................ 27
2.4.4. Recycling rate at subduction zones ............................................. 27
2.5. Thermo-mechanical recycling of the lithosphere ........................................... 28
2.5.1. Lithosphere delamination ................................................. 28
2.5.2. Delamination of the lower crust: geochemical evidence ................................... 29
2.5.3. Delamination of the lower crust: geophysical perspective .................................. 29
2.5.4. Thermal and thermo-mechanical erosion .......................................... 30
3. Peculiarities of crustal velocity structure in central and Western Europe ................................... 30
3.1. Making Paleozoic Europe...................................................... 30
3.2. Crustal thickness in Europe..................................................... 30
3.2.1. Crustal roots in young collisional orogens.......................................... 30
3.2.2. Cratonic crustal roots ................................................... 32
3.2.3. Crustal roots in Paleozoic orogens: Uralides......................................... 32
3.2.4. Missing crustal roots in Caledonides and Variscides ..................................... 34
3.2.5. Meso-Cenozoic rifts .................................................... 36
3.3. Seismic evidence for lower crustal delamination in the Variscides................................... 36
3.3.1. Mean crustal velocity in western Europe .......................................... 36
3.3.2. Thickness of crustal layers ................................................. 37
3.4. Structure of the Variscan lithospheric mantle ............................................ 37
3.4.1. Delaminated lithosphere ................................................. 37
3.4.2. Is the Variscan lithospheric mantle partly preserved? .................................... 37
4. Seismic reflections in the lithosphere of Western Europe .......................................... 38
4.1. Reflections from the lower crust .................................................. 38
4.1.1. Observations ....................................................... 38
4.1.2. Interpretations ...................................................... 39
4.2. Reflections in the upper mantle .................................................. 40
5. How was the Variscan crust created? ................................................... 40
5.1. Major characteristics of the present-day Variscan lithosphere..................................... 40
5.2. Why is the mafic lower crust missing? ............................................... 42
5.3. Unanswered questions ....................................................... 42
6. Conclusions ............................................................... 43
Acknowledgments............................................................... 44
References .................................................................. 44
1. Introduction
1.1. Primary crusts and planetary-scale differentiation
In contrast to undifferentiated bodies like most asteroids, all terres-
trial planets have a differentiated internal structure (e.g. Nimmo and
McKenzie, 1998; Robinson et al., 2008; Sohl et al., 2005; Solomon,
1980). Planetary-scale differentiation leads to formation of the core
and the bulk silicate shell represented by the mantle and the crust. The
fundamental differentiation of the terrestrial bodies into core, mantle
and crust has taken place already in the first 30–100 Ma after the crea-
tion of the solar system around 4.6 Ga (Bourdon et al., 2008; Boyet and
Carlson, 2005; Carlson and Lugmair, 1988; Harper et al., 1995; Harrison
et al., 2005; Iizuka et al., 2010; Lee and Halliday, 1995; O'Neill and
Palme, 1998). On the Earth, the initial internal structure was modified
later in the course of intense meteorite bombardment which continued
from the moment of the Earth's formation at ca. 4.56 Ga until ca. 4 Ga.
The largest collision with a planetary body (the “Giant Impact”)thathit
the Earth at ~4.53 Ga had a major effect on the compositional structure
of the Earth by separating the Moon (Canup, 2004; Kleine et al., 2005;
Lucey et al., 2006; Pritchard and Stevenson, 2000) with an anomalously
small core volume as compared to the Earth and thus leaving the Earth
with a somewhat modified ratio of the core to the bulk silicate shell vol-
umes. Heavy meteorite bombardment erased most of the evidence for
the pre-3.9 Ga geological and tectonic evolution of the planet (Ryder,
2003). Although (close to) none of the primary crust remains on the
Earth, recent development of experimental geochemistry including
high-resolution geochronology provides growing evidence on the
existence of the Earth's crust in the Hadean (Dauphas et al., 2007;
Harrison et al., 2005; Iizuka et al., 2010).
Planetary crusts are formed during and after freezing of a magma
ocean, either from low-density material produced by fractional crys-
tallization within the cooling magma ocean (primary crust), or by
magmatic additions (volcanic or intrusive) upon or into the pre-
existing crust which are produced by subsequent partial melting of
the mantle (Sharkov and Bogatikov, 2009; Taylor, 1989a). The lunar,
the Martian, and the Mercurial crusts are thick: 60 to 120 km on the
Moon, 50 to 100 km (and perhaps up to 150 km) on Mars (with a pro-
nounced crustal dichotomy—a thick, ca. 100 km, crust in the southern
hemisphere and a thin, ca. 50 km, crust in the northern hemisphere),
and 100 to 200 km on Mercury (Gudkova and Zharkov, 2004;
Neumann et al., 2004; Sohl and Schubert, 2007; Sohl et al., 2005;
Solomon, 2003; Solomon et al., 2005; Spohn et al., 2001; Turcotte et
al., 2002; Zuber, 2001). It is speculative how much of these crusts
are the primary crusts. Crustal thickness on Venus ranges from 20
to 50 km, perhaps even from 0 to 90 km (Anderson and Smrekar,
2006; Nimmo and McKenzie, 1998; Simons et al., 1997), with signif-
icant lateral variations as on the Earth, where the crustal thickness
has a global average of 20–25 km and ranges from 5.6–7.1 km in “nor-
mal”oceans (i.e. oceans where the bathymetry follows the square-
root of ocean-floor age dependence) to 30–50 km in the continents,
regionally thickening to 70–80 km in the Andes and Tibet (Beck,
2002; Murphy et al., 1997).
It is clear that the ratio of the crustal to the total mantle volume is
significantly different on various terrestrial planets: ~1.2% of on the
Earth, 1–4% on Venus, 5–11% (or more) on Mars, ~11–22% on the
19I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
Moon, and 10–30% on Mercury (Fig. 1b). The volume of the primary
crust extracted from the magma ocean during planetary-scale differ-
entiation is proportional to the mantle volume, but also depends on
mantle composition and P–T conditions in the planetary interior.
The mantles of the Earth, Mars, Venus, and the Moon are thought to
be similar in chemistry and mineralogy (Nimmo and McKenzie,
1998; Sohl and Schubert, 2007; Wieczorek et al., 2006), whereas on
Mercury where the ratio of the core relative to the bulk silicate frac-
tion is an unusually high mantle composition can be significantly dif-
ferent (Fig. 1)(Clark, 2007; Rothery et al., 2010; Taylor and Scott,
2004). Given different interior temperatures and sizes of the planets,
P–T conditions for crustal extraction are very dissimilar on all terres-
trial planets, except perhaps on Earth and Venus. We hypothesize
that plate tectonics processes, together with basalt–eclogite transi-
tion, additionally contribute to a significant difference in the ratio of
the crustal to the total mantle volume by limiting crustal thickness
on those planets where these mechanisms may operate.
1.2. Primary crusts and plate tectonics on terrestrial planets
1.2.1. Earth
Up to date, interpretations of the early plate tectonics on Earth re-
main highly controversial (see recent summaries by Condie and
Pease, 2008; Hatcher et al., 2007). Growing geological, geochemical,
tectonic, paleomagnetic, and seismic evidence indicates that some
kind of plate tectonics could have operated already on the Archean
Earth (Cawood et al., 2006). Some geological studies of the Archean
cratons place the start of plate tectonics at ca. 4.0 Ga (De Wit,
1998), whereas the oldest geophysically imaged signature of the
plate tectonics processes is ca. 2.7 Ga (Calvert et al., 1995) (see also
Section 2.4.1). The present-day interpretations of the geochemical
balance of the Earth provide the growing evidence that the crust
and the mantle may not be geochemically complementary reservoirs
if the Earth's composition is chondritic (e.g. Boyet and Carlson, 2005;
Jackson et al., 2010; Salters and Stracke, 2004; although recently
argued by Caro and Bourdon, 2010). Periodic switch of plate tectonics
could explain the geochemical disparity between the crustal and
mantle reservoirs (Carlson, 2011; Davies, 2011). This idea is sup-
ported by recent mantle convection models which show that hotter
mantle temperatures on the early Earth favor episodicity of Precam-
brian subduction and plate-driven tectonics (O'Neill et al., 2007b). A
periodicity in plate tectonics is caused by the time-dependent
changes in elastic lithospheric (the upper thermal boundary layer)
thickness and the convective stress, two factors that have opposite
effects on subduction initiation: only when the convective stresses
exceed the intrinsic strength of the lid, the lid fails and subduction
initiates (O'Neill et al., 2007a).
1.2.2. Venus
While the Venusian radius and the average crustal thickness are
similar to those on Earth, no topographic features resembling terrestrial
expressions of plate tectonics are known on Venus at present (Solomon
et al., 1992). Anhydrous, predominantly basaltic, composition of the Ve-
nusian crust (Surkov et al., 1986) implies its high viscosity and explains
major tectonic differences between the Earth and Venus (Mackwell et
al., 1998). On Earth weak boundaries separating strong plates play an
important role in plate tectonics (Bercovici et al., 2000). In contrast,
the present absence of water on Venus results in strong faults (Foster
andNimmo, 1996) and high mantle viscosity (Moresi and Solomatov,
1998; Zhong and Gurnis, 1996) and leads to stagnant lid convection
(confined to the deep interior below thick and strong lithosphere).
Given that the age of the Venusian surface is only 300–600 Ma
(Frankel, 1996; Herrick, 1994), plate tectonics could have existed on
Venus prior to the global resurfacing event; the latter could have
recycled the primary crust. Although rapid resurfacing on Venus could
have been caused either by plate tectonics processes or by plume-
related melt generation, plume activity alone seems to be insufficient
to produce global resurfacing, and plate tectonics processes seem to
be more plausible (Nimmo and McKenzie, 1998). Numerical models in-
dicate a possibly episodicpattern of plate tectonics with short periods of
massive subduction separated by periods of surface quiescence (O'Neill
et al., 2007b; Tackley, 2000; Turcotte, 1993).
While the Venusian crust could have been thinned to its present
thickness during the resurfacing event with an involvement of plate
tectonics processes, other processes (in the absence of the present-
day plate tectonics), presently restrict thickness of the Venusian
crust to depths similar to global average of the Earth's crust. Given
that the Venusian crust is basaltic, its present thickness is thought to
be controlled by basalt–eclogite transition (Nimmo and McKenzie,
1998), which occurs at depths of 30–70 km, depending on tempera-
ture (Green and Ringwood, 1967; Ito and Kennedy, 1971). However,
under dry conditions and in the absence of deformation basalt may re-
main metastable even to greater depths (Rubie, 1990).
Fig. 1. Sketch of the internal structure of the terrestrial bodies and the relative volume
of the crust, mantle and core as percent of the total volume of the bodies (based on data
of Gudkova and Zharkov, 2004; Nimmo and McKenzie, 1998; Solomon, 2003; Spohn et
al., 2001; Wieczorek et al., 2006; Zuber, 2001). Plate tectonics processes, together with
basalt–eclogite transition in the (lower) crust, may have a strong control on the crustal
thickness (b).
20 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
1.2.3. Mars
Alike on Venus, basalt–eclogite transition may also control thick-
ness of the Martian crust, which is proposed to have andesitic–basaltic
or basaltic shergottites composition (Sohl et al., 2005). Although satel-
lite observations indicate the absence of the present-day plate tecton-
ics on Mars (Zuber, 2001) and there is no robust evidence that it ever
operated there, some authors suggest that plate tectonics could have
operated on this planet in the past (Nimmo and Stevenson, 2000;
Sleep, 1994; although questioned by van Thienen et al., 2004),
depending on whether free liquid water was present on the Martian
surface (Catling, 2004; O'Neill et al., 2007a). The existence of plate tec-
tonics on Mars during its early evolution is supported by linear mag-
netization in the southern hemisphere (Connerney et al., 1999;
Lenardic et al., 2004).
1.2.4. Moon and Mercury
Both Moon and Mercury are likely to have been stagnant since their
formation until present due to the absence of surface water (Lewis,
1988; Lissauer, 1997). Basalt–eclogite phase transformation also
seems to be unimportant in limiting thicknesses of the non-basaltic
lunar and Mercurial crusts which are made chiefly of primordial anor-
thosites (e.g. Ashwal, 2010; Greenhagen et al., 2010; Jeanloz et al.,
1995). As noted earlier, both Moon and Mercury have a very high
ratio of crustal to mantle volume, which may be a consequence of the
absence of both plate tectonics and crustal delamination advanced by
eclogitization.
The observation that the terrestrial bodies with (i) continuous or
episodic plate tectonics and (ii) basaltic (lower) crust which may un-
dergo eclogitization have a very low ratio of crustal to mantle volume
(Fig. 1b) advocates a strong role of plate tectonics processes which,
together with basalt–eclogite transition, limit crustal thickness by
providing effective mechanisms of crustal (lithosphere) recycling
into the mantle. The next sections discuss some evidence in support
of this hypothesis by reviewing the crust–mantle material exchange
processes that operate on Earth.
1.3. Geodynamic controls of crustal volume
We now focus our discussion on the crustal structure on Earth, the
characteristic features of which are: (1) a (nearly) complete absence of
the primary crust extracted from the early magma ocean and (2) the
presence of the continental crust which cannot be extracted directly
from the mantle. Crustal thickness (and thus the total volume of the
planetary crust) is controlled by a dynamic interplay of crust–mantle
material exchange processes (crustal growth and crustal recycling),
and not all of them require plate tectonics. Major geodynamic processes
responsible for growth and recycling of the Earth's crust include (Fig. 2):
mafic magmatism, crustal (lithospheric) delamination, oceanic sub-
duction, and continental subduction. By continental subduction we
mean subduction associated with collisional orogens although, strictly
speaking, in most cases slabs correspond to pre-collisional subduction
of oceanic lithosphere that has been consumed from the surface.
(i) Crustal growth. Plume-related magmatic additions (volcanic or
intrusive) to the crust and crustal (or lithosphere) delamination
may exist both on the Earth and on terrestrial bodies with
stagnant-lid convection regime. In the presence of plate tecton-
ics, additional mechanisms of crustal volume growth through
mafic magmatism are those related to the presence of mid-
ocean ridges and subduction zones: (a) decompressional melting
of passively ascending mantle material at mid-ocean ridges re-
sponsible for formation of new oceanic crust (Fig. 3a) and (b)
magmatic processes at subduction zones responsible for forma-
tion of new continental crust (Fig. 3b). Scenarios for generation
of the continental crust are highly debated, given that granitic
magmas cannot be extracted directly from the upper mantle
and their petrogenesis remains controversial (Kemp and
Hawkesworth, 2003; Petford et al., 2000). Intraplate plume-
related magmatism (which does not require plate tectonics re-
gime) with associated magmatic underplating contributes not
only to the crustal growth but also to chemical and structural
modification of the continental crust. Major tectonic settings
where magmatism and the associated crustal growth can take
place are summarized in Fig. 4a.
(ii) Crustal recycling. Crustal growth is limited by two plate tectonic
processes, namely oceanic and continental subductions, as well
as by crustal (lithospheric) delamination (the latter may also be
caused by density changes due to phase transitions, e.g. eclogiti-
zation, and thus may operate in the absence of plate tectonics).
These processes recycle lithospheric material back into the man-
tle and thus control the crustal volume. In the “feed-back loop”,
the recycled crustal material, being chemically and mineralogi-
cally different from the mantle, stimulates mantle melting and
magmatism, in particular at convergent settings, but also at intra-
plate settings during lithospheric extension. Extensional process-
es themselves can be initiated and supported by passive mantle
upwellings generated by thermo-chemical convection or by reor-
ganization of lithospheric plates.
Diversity of tectonic settings within the European continent allows
to find regional illustrations to all major geodynamic processes respon-
sible for crustal growth and its recycling into the mantle. Phanerozoic
evolution of the crust, the main topic of our research in central and
western Europe, was governed and controlled by modern plate tectonic
processes which operated together with mantle dynamics. In western
Europe they included Phanerozoic orogenic events related to the colli-
sion of Baltica, Laurentia, and Avalonia, the closure of the Iapetus
Ocean and the Tornquist Sea, and subsequent amalgamation of a series
of terranesat the edge of the East European craton. We start with a gen-
eral overview of the geochemical perspective on crust–mantle links
(Section 2); it provides grounds for better understanding of the tectonic
Fig. 2. Sketch of major processes controlling crust–mantle material exchange. Vertical and horizontal dimensions are not to scale.
21I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
and geodynamic processes that are imaged by seismic methods. We
next summarize geophysical (mostly seismic) evidence for crust–
mantle material exchange processes using continental Europe as an
example (Section 3). Our special focus is on the Variscan crust which,
despite its Paleozoic age, has crustal structure typical for present-day
extensional settings. We discuss the peculiarities of the structure of
the Variscan crust, its past and modern tectonic analogues, and tectonic
processes which may have created the modern lithospheric structure of
the Phanerozoic Europe (Section 4).
2. Crust–mantle exchange processes: a global view
2.1. Crust–mantle compositional balance: a planetary-scale geochemical
perspective
Planetary-scale differentiation resulted in the formation of two
fundamentally different portions of the Earth: the iron-rich core
that was formed at 33±2 My after the beginning of the Solar System
(Kleine et al., 2002) and the outer silicate portion of the Earth (the
mantle plus the crust) that was termed by Hart and Zindler (1986)
as the Bulk Silicate Earth (BSE). The BSE consists of several geochem-
ical reservoirs (McDonough and Sun, 1995; O'Neill and Palme, 1998;
Taylor and McLennan, 1995):
•the primitive (undifferentiated) mantle that, by definition, has
composition of BSE;
•the depleted convecting upper mantle (with a mass fraction of
~0.50± 0.15 of BSE) that is the source of mid-ocean ridge basalts
(MORB);
•the enriched upper mantle that is the source of ocean–island basalts
(OIB);
•the oceanic crust;
•the suboceanic lithospheric mantle;
•the continental crust (with the mass fraction of 0.0054 of BSE);
•the subcontinental lithospheric mantle (with a mass fraction of no
more than 0.04 of BSE);
•hydrosphere and atmosphere.
Decompression-induced partial melting of depleted mantle is the
process thought to be responsible for generation of mid-oceanridge ba-
salts (MORB) and thus of the oceanic crust (Fig. 3a). Modern production
rate of MORB is estimated to be 20 km
3
/year (Hoffman, 1989)(Fig. 4b).
The products of melt segregation, namely the oceanic crust formed by
the melt material and the residual depleted mantle, are chemically
complementary (O'Neill and Palme, 1998). Oceanic crust with an aver-
age thickness of ~7 km can be produced by adiabatic upwelling of the
mantle with a potential temperature of ~1300 °C, while higher poten-
tial temperature (1450–1500 °C) generates enough basaltic melt to
produce oceanic crust up to 25 km thick (e.g. McKenzie and Bickle,
1988). Cooling of the oceanic plate, when it moves away from the
mid-ocean ridge, produces oceanic lithosphere which has the same
composition as the surrounding mantle. A significant amount of the
oceanic crust and the associated residual depleted mantle is recycled
back into the depleted mantle in subduction zones. However, island-
arc magmatism that plays an important role in formation of the conti-
nental crust, brings back some of the most highly incompatible trace
elements to the near-surface settings (Fig. 3a).
Fig. 3. Sketch illustrating geochemical relations between mantle, oceanic crust and
continental crust. Melting of the depleted convecting upper mantle generates mid-
ocean ridge basalts and produces oceanic crust. A significant amount of the oceanic
crust together with the associated residual depleted mantle is recycled back in subduc-
tion zones refertilizing the mantle and producing island-arc magmatism which plays
an important role in formation of the continental crust. The enriched upper mantle is
the source of ocean-island basalts. Large-scale mantle upwellings (plumes) as well as
small-scale convective instabilities (not shown) transport mantle material into the
continental lithosphere and lead to crustal growth, particularly notable in LIPs. Vertical
and horizontal dimensions are not to scale. Fig. 4. Major types of magmatism. (a) Diagram showing major types of magmatism as-
sociated with specific tectonic settings. (b) Global rates of Cenozoic magmatism in var-
ious tectonic settings (based on data of Wilson, 1995). Later studies indicate greater
role of crustal production at convergent margins. See text for discussion.
22 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
In contrast to the oceanic crust, the continental crust cannot be di-
rectly differentiated from the mantle, and most geochemical models
propose a two-stage process with the initial extraction of basaltic
magmas from the mantle and their subsequent remelting or fractional
crystallization (Rudnick, 1995, 2003; Taylor and McLennan, 1985).
Trace-element abundances in the average continental crust and normal
mid-oceanridge basalts (N-MORB) are complementary: enrichments of
the continental crust in incompatible elements are mirrored by deple-
tions in the same elements in N-MORB (Taylor and McLennan, 1985).
This observation led to the key concept in mantle geochemistry: the
upper mantle is depleted (as compared to the primitive mantle) pri-
marily due toextraction of the continental crust. Geochemical data indi-
cate the depleted mantle should fill not only the upper mantle, but also
asignificant portion of the lower mantle (McCulloch and Bennett,
1998). However, the composition of the continental crust, that has
~60% of SiO
2
, is unlikely to be in the equilibrium with the mantle com-
position(e.g. Hawkesworth andKemp, 2006a). In particular, granitic (in
a broad sense) rocks play an important role in crustal evolution (Douce,
1999). Their petrogenesis remains controversial leading to different,
highly debated, scenarios for the continental crust generation.
2.2. Making new continental crust
2.2.1. Growth of continental crust: convergent margins versus intraplate
settings
Since granitic magmas cannot be extracted directly from the upper
mantle, three geochemical processes by which the granitic crust can be
formed include (Kemp and Hawkesworth, 2003): “(i) fractional crystal-
lization of primary basaltic liquid, (ii) mixing between partial melts of
pre-existing crust and mantle-derived magmas, or their differentiates,
and (iii) partial melting of young, mantle-derived mafic protoliths in
the crust”(i.e. remelting within the crust). Geodynamic processes (mu-
tually non-exclusive) by which the continental crust can be produced
include:
•subduction-related processes at convergent margins (island-arcs
settings) (Fig. 3a) and
•plume-related magmatism with fractional crystallization of primary
basaltic liquid (Fig. 3b).
At compressive tectonic settings, growth of the continental crust oc-
curs by additions of new continental crust (produced by magmatic
underplating and mixing between partial melts of pre-existing crust
and mantle-derived magmas) and also by accretion of oceanic plateaus,
submarine turbidite fans, and fragments of juvenile oceanic crust. In in-
traplate settings, plume-related magmatism with associated magmatic
underplating contributes to the crustal growth, if granitic magmas
contain a component of juvenile mantle-derived material. Since the
lowermost crust typically does not reach temperatures high enough
(800–900 °C) for partial melting (Artemieva and Mooney, 2001)
(Fig. 5), generation of a significant volume of granitic magmas on time
scales of up to 100 My requires heat advection to the crustal base, for
example by underplating or interplating by mafic magmas (Clemens,
2006)(seeSection 2.3).
Primitive island-arc basalts (IAB) and ocean–island basalts (OIB)
represent two end-member compositions of basaltic magmas that
make new material added to the continental crust. The former constrain
compositions of basaltic magmas typical for subduction settings, while
the latter—composition of the present-day intra-plate magmas (e.g.
Sun and McDonough, 1989). Despite IAB and OIB represent two end-
member basaltic compositions that participate in the formation of
new continental crust, average major element composition ofthe latter
does not correspond to a simple mixture between IAB and OIB. Instead
it is represented by average composition of the lower crust that can be
modeled by a mixture of ~92% of IAB and ~8% of OIB. It implies that
throughout the Earth's history the continental crust is primarily
generated at convergent margins (island-arcs settings) (Barth et al.,
2000; Rudnick, 1995).
Nevertheless, plume-related magmatism could have been a princi-
pal mechanism for generation of a new crust in the Archean (Fig. 6),
when plate tectonics was only beginning to operate (the debate about
the start of plate tectonics is presented by Condie and Pease, 2008 and
is outside the scope of the present review). In particular, the existence
of convergent margins with the associated subduction zones in the
Archean is not indisputably proven; an exception may be a 2.69 Ga sub-
duction imaged by seismic reflection in the Canadian Shield (Calvert
etal., 1995). Plume-related magmatism provides a convenient explana-
tion for the origin of tonalite–trondheimite–granodiorite (TTG)
Fig. 5. Typical temperatures at the base of the continental crust versus tectono-thermal
age (based on data of Artemieva and Mooney, 2001). Boxes correspond to different tec-
tonic structures. The plot accounts for significant regional variations in crustal thickness.
Fig. 6. Sketch showing relative contributions of magmatism caused by mega-plumes and
mantle overturns (red) and arc-related magmatism at convergent settings (blue) to gen-
eration of the crust. While plume-related magmatism may have been important on the
early Earth, the role of arc-related magmatism became dominant with the start of plate
tectonics. The upper plot shows distribution of superplume events (from Condie, 2004).
If peaks in crustal growth at ca. 1.9 Ga and 1.1 Ga were caused by mega-plume events
(red dashed lines), the relative contribution of subduction-related magmatism to crustal
growth should have been reduced (blue dashed lines).
23I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
gneisses which dominate the Archean middle crust, although subduc-
tion related processes in combination with plume-related melting can
also explain the origin of the TTG and growth of the early continental
crust (e.g. Kemp and Hawkesworth, 2003; Zegers andvan Keken, 2001).
While massive mantle melting caused by heads of mega-plumes
or mantle overturns may have been the dominant mechanism at the
early stages of the lithosphere formation, the proportion of the crust
generated by plume-related oceanic plateau basalts probably de-
creased from the Archean to the present day (Condie, 1994) due to
a secular decrease of mantle potential temperature (Campbell and
Griffiths, 1992). Arc-related magmatism at convergent margins is
expected to become dominant with mantle cooling and at present
most of the continental crust is generated at arc settings (Fig. 6).
In case the dominant mechanism of continental crust production
has changed during the evolution of the Earth, a systematic difference
between the bulk compositions of the Archean and post-Archean
crust with respect to key chemical indices or elemental ratios should
be expected. However, isotope analysis of Nb/La ratios, which are sen-
sitive to the proportion of plume component in volcanic rocks, does
not indicate that the contribution of plume-related magmas to crust
generation has decreased since the Archean (Hofmann et al., 1986;
Kemp and Hawkesworth, 2003). This conclusion is in apparent con-
tradiction with the earlier mentioned results (Barth et al., 2000;
Rudnick, 1995) and Figs. 4, 6.
2.2.2. Growth rate of the continental crust
There is an on-going debate on secular variations in growth and
recycling rates of the crust and the lithospheric mantle, that is on the
relative roles of different tectonic processes responsible for upward-
and downward-directed material exchange, which apparently have
not been constant through the Earth's evolution. Based on geological
evidence, Dewey (2007) argues that downward-directed material ex-
change processes could have been most dramatic at the early stages of
the Earth's evolution and it was not until the mid-Archean that crustal
growth and crustal recycling became roughly compensated. Some
geochemical evidence (Harrison et al., 2005) suggests that the early
basaltic/komatiitic crust could have been entirely recycled back into
the mantle in the Hadean (>3.5 Ga).
There is a significant controversy in the estimates of the crustal
growth rate from the Hadean until present. Global geophysical and geo-
chemical evolutionary models suggest that crustal growth rate was
~2–7km
3
/yr during the Hadean and less than 2 km
3
/yr in the post-
Hadean (b3.5 Ga) time (e.g. Armstrong, 1981; Dewey and Windley,
1981; Reymer and Schubert, 1984). Global and continent-scale isotope
studies (although commonly significantly biased by sampling) shift the
peak in crustal growth to the Archean–Paleoproterozoic (with a rate of
~2–7km
3
/yr) and indicate low Phanerozoic growth rates, b1km
3
/yr
(Allègre, 1982; McLennan and Taylor, 1982; Veizer and Jansen, 1985).
In contrast, some isotope studies (based on Nd and U/Pb isotopes,
and also significantly biased by sampling in the Laurasia continents) in-
dicate episodic growth of continental crust: major peaks at 2.7–2.5 Ga,
1.9–1.7 Ga, and 1.1 Ga with growth rates of juvenile crust of
~6–7km
3
/yr are separated by “quiet epochs”with the crustal growth
rate of ~1 km
3
/yr (Condie, 1998; McCulloch and Bennett, 1994). How-
ever, recent study based on ca. 13,800 integrated U–Pb and Hf-isotope
analyses of(largely detrital)zircons suggests little episodicity in the for-
mation of the continental crust (Belousova et al., 2010). It argues that
ca.60% of the present volume of the continental crust could have been
separated from the mantle in the pre-Proterozoic time (before
2.5 Ga), whereas three major isotope peaks represent the time of
major magmatic events and the associated crustal reworking. As
noted by Condie (1998), the peaks at 2.7–2.5 Ga, 1.9–1.7 Ga, and
1.1 Ga are well correlated with global plate tectonic events such as as-
sembly and dispersion of supercontinents and the time of the emplace-
ment of giant dyke swarms (GDS) and komatiites (Fig. 6). Variations in
volume of the present-day continental lithospheric mantle (that reflects
its “survival rate”, i.e. the difference between the growth rate and the
recycling rate) also show similar peaks (Artemieva, 2006). Note that re-
cent numerical simulations of mantle convection indicate that, due to
higher mantle temperatures on the early Earth, the Precambrian
plate-driven tectonics is likely to have been episodic, with pulses of
crustal growth being a direct consequence of rapid plate motions
(O'Neill et al., 2007b).
2.3. Mafic magmatism: volcanism and intrusions
2.3.1. Geodynamic causes of mafic volcanism
Basaltic magmatism plays an important role in melt generation on
the terrestrial planetary bodies, and it can exist on the terrestrial
bodies both with or without plate tectonics. Both in the absence and
in the presence of plate tectonics, mafic magmatism can be caused
by large-scale temperature anomalies in the mantle associated either
with large- or small-scale mantle convection or with mantle plumes.
An interaction of anomalously hot mantle with the upper thermal
boundary layer (the lithosphere) can lead to its thermo-mechanical
erosion, extension, and delamination (Davies, 1999; Ebinger and
Sleep, 1998; Korenaga and Jordan, 2002). When plate tectonics oper-
ates, mafic magmatism can be a consequence of lithospheric subduc-
tion (Hoernle et al., 2008; Scholl and von Huene, 2007).
Basalt is the primary melt formed during mantle melting on Earth.
The role of mantle melting and the associated mafic magmatism as a
major mechanism of growth of new crust has been discussed above.
Here we mention briefly another role that mafic magmatism and plu-
tonism play in crustal (and lithosphere) evolution, that is to supply
mantle material into the lithosphere by intrusions and volcanism in in-
traplate settings (Fig. 4a). However, the volume of mantle material
transported into the crust in intraplate settings is small as compared
to the voluminous magma generation at the plate boundaries
(Fig. 4b). In the oceans, intraplate magmatism is responsible for forma-
tion of oceanic islands and ocean plateaus, that is “anomalous oceans”
where the bathymetry does not follow the square-root of age depen-
dence. The largest known oceanic large igneous province (LIP) is the
Ontong-Java Plateau.
On the continents, which are the principal focus of our study, intra-
plate magmatism is responsible for epeirogenesis and crustal thicken-
ing (Kay et al., 1992; McKenzie, 1984; Mo et al., 2007). The most
voluminous magmatic additions to the continental lithosphere include
LIPs (with continental flood basalt provinces) and, mostly Precambrian
in age, GDS (Mahoney and Coffin, 1987); both are formed by mantle
melting caused by large-scale or small-scale mantle convection instabil-
ities (King and Anderson, 1995; Yale and Carpenter, 1998). The Archean
ages of the oldest known continental rifting events in the Kalahari cra-
ton and in Western Australia (e.g. Olsson et al., 2010; van Kranendonk
et al., 2010) and the oldest giant mafic dyke swarms in the southern
West Greenland (Nilsson et al., 2010) indicate that thermo-
mechanical interaction between the ascending mantle material and
the continental lithosphere has played an important role in the evolu-
tion of the continents since their very formation. Commonly maficmag-
matism has a regional coverage; for example, small-scale convection
instabilities termed “baby-plumes”were proposed to explain anoro-
genic Cenozoic igneous provinces in Europe (Lustrino and Wilson,
2007; Ritter and Christensen, 2007).
2.3.2. The shape of magmatic intrusions
Emplacement of high-density, mafic–ultramafic intrusions into
low-density crustal rocks is widely observed in many tectonic set-
tings. Geological observations as well as numerical modeling of intru-
sion of partially molten mantle rocks produced in a sub-lithospheric
magmatic source indicate that geometry of intrusions critically de-
pends on lithosphere rheology and ranges from dikes to deep-
seated balloon-shaped intrusions and to flat sills (Gerya and Burg,
2007). In case of cold lithospheric temperatures (typical for cratonic
24 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
settings) and elasto-plastic crustal rheology, upward magma propa-
gation is controlled by crustal faulting and makes finger-shaped
dikes. Higher lithospheric temperatures favor lateral spreading of
magma above the Moho with the formation of flattened shallow mag-
matic sills (Fig. 7). Such geometry of mafic intrusions is typical for the
crust of Western Europe, but it can also be observed in continental rift
zones and in some Precambrian terranes (e.g. large mushroom-
shaped plutons are imaged by the BABEL seismic profiles within the
crust of the southern Swedish Baltic Shield (Korja and Heikkinen,
2005)).
Emplacement of basaltic magmas into the continental crust leads
to its heating and consequent melting and can generate a large vol-
ume of silicic (felsic) magmas which may form granitic plutons (e.g.
Coldwell et al., 2011; de Saint Blanquat et al., 2011; Guo et al., 2007;
He et al., 2011)(Fig. 8). For example, a 70 km thick crust is produced
in the Andes mostly due to sialic intrusions. Most of granitic rocks are
formed by an interaction of basaltic melts with metamorphic rocks of
supracrustal origin (Douce, 1999). Theoretical models indicate that a
500 m basaltic sill can generate the silicic magma layer with a thick-
ness ranging from 300 to 1000 m depending on crustal temperatures
(Huppert and Sparks, 1988). The time-scales for granitic magma gen-
eration in basalt-induced crustal melting are very short, hundreds to
thousands of years, and even when large volume of magmas are pro-
duced, granite magmatism is a rapid process with timescales less than
100,000 years, irrespective of tectonic setting (Petford et al., 2000). In
comparison, thermal relaxation of the continental crust takes place on
the time-scale of tens of millions of years and the lifetimes of large si-
licic magma systems exceed millions of years. Progressive heating of
the crust by basaltic intrusions may result in cycles of mafic-to-
progressively silicic magmatism commonly observed in nature (de
Saint Blanquat et al., 2011). Crustal heating by mafic intrusions also
weakens the lithosphere and may lead to the lower crustal flow
(Fischer, 2002). Being commonly associated with lithosphere exten-
sion and stretching (Le Bas, 1987), magmatism can lead to crustal
thinning due to removal of the lower portions of lithosphere as ob-
served in large parts of central and western Europe (Matte, 1991)
(Section 2.5).
Recent seismic study of the Baikal rift zone indicates an important
role of magmatic intrusions in continental rifting. In contrast to clas-
sical models of continental rifting, no crustal thinning is observed
across the Baikal rift zone. The flat Moho can be explained by
magma intrusions into the lower crust as indicated by the presence
of the extremely high-velocity (7.4–7.6 km/s) and highly reflective
zone in the lower crust (Thybo and Nielsen, 2009). Observed seismic
reflectivity can be explained by a layered sequence of strong reflec-
tors (ca. 400 ±150 m thick with velocity contrast of ~ 0.4 km/s)
which occupy around 50 vol.% of the 13-km-thick lower crust. As a re-
sult, the estimated extension β-factor increases from 1.3 to 1.7 if the
lower-crustal intrusions are taken into account. Similarly, seismic ob-
servations in the Devonian Dnieper-Donets rift zone in the south-
western part of the East European craton (Wilson et al., 2004)indi-
cate aflat Moho and high crustal reflectivity, which can be explained
by the presence of mafic intrusions in the lower crust that may form
intermixed layering of primarily olivine- and pyroxene-rich rocks
with layers of plagioclase-rich components (Lyngsie et al., 2007).
Intraplate magmatism is associated not only with mafic intrusions
into the crust, but also with intrusions of basaltic magmas into the lith-
ospheric mantle; many of them may never reach the surface. Infiltration
of basaltic magmas into depleted cratonic lithosphere during large-
scale lithosphere–mantle interaction may lead to its compositional
modification (melt-metasomatism) as evidenced by mantle-derived
peridotite xenoliths from the Wyoming craton in North America and
the Sino-Korean craton in China (e.g. Eggler et al., 1988; Griffinetal.,
1998). Similarly, the proposed melt-metasomatism of the lithospheric
mantle of the southern part of the East European craton (as evidenced,
in the absence of mantle xenoliths data, by its post-Devonian subsidence
history and by gravity and buoyancy modeling) may be associated with
the Devonian plume beneath the Dnieper-Donets Rift and the Peri-
Caspian-northern Caspian depressions (Artemieva, 2003; Artemieva et
al., 2006; Kaban et al., 2003).
2.4. Subduction
2.4.1. Precambrian subduction
A large number of numerical studies have investigated the param-
eter space that defines physical conditions on the terrestrial planets
favorable for a plate tectonics regime. These studies demonstrated
that three qualitatively different styles of mantle convection are pos-
sible: stagnant-lid, episodic, and mobile convection (plate tectonic re-
gime) (Fig. 9). In the latter case, the negative buoyancy of the thermal
boundary layer on the top of the mantle causes subduction and drives
plate tectonics. The ratio of driving to resisting stresses defines the
parameters critical for plate tectonics regime: when the convective
Fig. 7. Sketch showing the “stability fields”for mafic intrusions as a function of magma
viscosity and rheology of the lower crust (based on numerical modeling of Gerya and
Burg, 2007). “Cold”and “hot”crust correspond to Moho temperatures of 400 °C and
600 °C, respectively, for the crustal thickness of 40 km.
Fig. 8. Sketch showing evolutionary scheme for a formation of a silicic magma system by basalt emplacement into the crust (after Huppert and Sparks, 1988). At early stages most
basalt reaches the surface. At the next stage, the crust becomes heated by basaltic magmas and its melting starts. Basalt, which is mostly trapped within the crust, does not form
volcanic fields; silicic magmas which are episodically produced can either erupt or form shallow intrusions. At the final stage, a large volume of the crust reaches melting temper-
atures producing large volumes of silicic magma, while basaltic magmatism can still be present in the peripheral parts.
25I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
stresses produced due to the formation of sinking thermals exceed
the intrinsic strength of the lid (the upper thermal boundary layer,
or the lithosphere), the latter fails initiating subduction (Solomatov
and Moresi, 1997). The convective viscous stresses depend on the
vigor of mantle convection which is largely controlled by mantle tem-
perature, whereas the resisting stresses are controlled by rheology of
the overlying lid, which essentially depends on temperature and the
presence of free water at the surface (Bird et al., 2008; O'Neill et al.,
2007a). Since mantle temperatures and the concentration of heat
producing elements are thought to be significantly different on the
early Earth (Michaut and Jaupart, 2004; Nisbet and Fowler, 1983),
the convection style on the early Earth could have also been different
than today.
Numerical simulations examine the effect of high mantle tempera-
tures on the style, and even on the possibility, of oceanic subduction
in the Archean. In particular, due to increased mantle temperatures in
the Archean, lithospheric stresses were lower, and could lead to episo-
dicity of plate tectonics processes on the early Earth when rapid pulses
of subduction were interspersed with periods of relative quiescence
(O'Neill et al., 2007a). Other mantle convection simulations (Sizova et
al., 2010) suggest that no subduction of oceanic plate could take place
if mantle potential temperature (hypothetical temperatures of mantle
adiabatically brought up to the surface without melting) was 250 °K
hotter than the present day mantle, whereas in case mantle tempera-
ture was only ~150–200 °C above the present value, lithosphere defor-
mation was characterized by “pre-subduction”tectonic regime with
shallow underthrusting of the oceanic plate under the continental
plate (Fig. 9). Thus, estimates of mantle potential temperatures in the
Precambrian are critical for understanding of the style of oceanic sub-
duction on the early Earth.
Thebestknowntemperatureestimateisbasedondataforthe
Archean komatiites which was interpreted to indicate that the Archean
mantle was at 1600–1800 °C (Nisbet and Fowler, 1983), while potential
temperature of the present-day mantle is commonly estimated to be be-
tween 1300 °C and 1450 °C (e.g. McKenzie et al., 2005). Early estimates of
Archean mantle temperatures have been questioned during the past de-
cades. In particular, studies of MORB-like ophiolite suites and greenstone
belts suggest that Archean mantle was at temperatures of 1430–1500 °C
(Abbott et al., 1994), while melting models for hydrous komatiites (Grove
and Parman, 2004) and mantle convection models (Campbell and
Griffiths, 1992) suggest that it could have been only ~50 °C hotter than
the present mantle. In case the Archean mantle temperature was similar
to the present day value, oceanic subduction could have started very early
in the Earth's evolution. However, in case of high mantle temperatures in
the Archean, no oceanic subduction existed on the Archean Earth, and the
transition from “no subduction”to “pre-subduction”tectonic style has
probably occurred at 3.2–2.5 Ga.
Even if oceanic subduction has operated already in the Archean, its
geodynamic signature was significantly different from modern sub-
duction (Abbott and Hoffman, 1984): (i) in contrast to steep post‐
Archean subduction, Archean subduction was buoyant, flat, and shal-
low and (ii) melting occurred in the slab wedge rather than in the
mantle wedge as in post‐Archean tectonics. Secular variations in
ridge length, plate velocities, and in the intensity of push‐pull forces
led to secular variations in the rate of material exchange between
the lithosphere and the mantle in oceanic subduction zones. In partic-
ular, if Proterozoic plate velocities were slower than at present
(Abbott and Menke, 1990), the lithosphere–mantle material exchange
in the Archean‐Paleoproterozoic could have been less vigorous than in
Phanerozoic.
The oldest seismically imaged subduction zone is ~2.69 Ga (Calvert
et al., 1995). It has been imaged by the LITHOPROBE seismic reflection
studies in the Superior Province of the Canadian Shield as dipping re-
flectors extending 30 km into the mantle with the geometry similar to
the reflectivity patterns in modern subduction zones and collisional
sutures. There is no indisputable evidence for Archean subduction
within the European continent. However, BABEL seismic profiles in
the Baltic Sea and in the North Sea imaged several distinct dipping
sub-Moho reflectors at the margins of the East European craton that
extend 30–80 km into the mantle (Fig. 10)(Abramovitz et al., 1998;
BABEL WG, 1990, 1993). These reflections dip at a 15° to 35° angle,
are traced laterally over distances of up to 100 km, and in two occur-
rences are accompanied by a sharp 5–7 km Moho offset. Their ages
vary from Proterozoic (ca. 1.9 Ga in the northern part of the Bothnian
Gulf and 1.6–1.4 Ga at the southern, Proterozoic, margins of the Baltic
Shield) to Phanerozoic (ca. 440 Ma for the reflector in the North Sea)
Fig. 9. Sketch illustrating three possible tectonic styles at convergent margins (based
on numerical modeling of Sizova et al., 2010). At mantle temperature ~250 °C higher
than the present-day mantle, oceanic subduction does not develop (bottom figure).
Mantle temperature ~150–200 °C higher than the present-day mantle favors “pre-sub-
duction”tectonic regime with shallow underthrusting of the oceanic plate under the
continental plate (middle figure).
Fig. 10. Observations of subcrustal distinct dipping seismic reflectivity in the Baltic
Shield (modified after Artemieva, 2011; Balling, 2000). Arrows show dip direction;
the age and the depth range of subcrustal reflectivity is shown in boxes. These dipping
seismic reflectors are interpreted as images of relict subduction zones and collisional
sutures. TESZ = Trans-European suture zone.
26 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
and are constrained by the ages of known major tectonic events in the
region.
2.4.2. Modern oceanic subduction
Lithosphere sinking in modern oceanic subduction zones is the
most widespread and powerful process of both growth of the crust
(see Section 2.2) and recycling of crustal/lithospheric material at ac-
tive margins. Recent estimates indicate that ~80% of the thickness of
the Ontong Java oceanic plateau has been subducted and recycled
into the mantle if the past convergence rate between the plateau
and the Solomon Island arc was the same as at present (Mann and
Taira, 2004).
In modern subduction zones, basaltic oceanic crust of subducting
slabs starts melting or transforms into denser eclogite phases at a rela-
tively shallow depth. Thus oceanic subduction transports eclogite as
well as serpentinite from themantle wedge deep into the upper mantle.
This process produces a large impedance-contrast boundary at slab–
mantle interface and ananisotropic boundary that are recognized in re-
cent seismic surveys of the Cascadia subduction zone in western North
America and Japan (Nikulin et al., 2009; Peacock, 2001). Eclogite addi-
tion to lithospheric mantle during oceanic subduction has been pro-
posed to explain the presence of eclogites in cratonic kimberlites that
are not in equilibrium with harzburgite (McGreggor and Manton,
1986). However, kimberlitic eclogites in cratonic settings can also be
residues from the process of Archaean granitoid crust formation or
crustal (lithosphere) delamination (Rollinson, 1997).
The bulk of the subducting oceanic lithospheric plate may reach,
or even cross, the 660 km discontinuity, as indicated by seismic tomo-
graphic models (e.g. van der Hilst et al., 1997). Mantle flow and vis-
cous drag may also have a strong effect on the penetration depth of
slabs (Boutelier and Cruden, 2008). In some cases, the subducting
plate may break, creating a slab window, that facilitates magmatism
and promotes additions of new mantle material to the crust. This pro-
cess is believed to operate along the Pacific coast of North America in
Mesozoic and Cenozoic, where the subducting Farallon plate was de-
tached and split into several microplates (Atwater and Stokes, 1998),
that are imaged by seismic tomography models of the upper mantle
of western USA (van der Lee and Nolet, 1997) and are reflected in
wide‐spread magmatism in the area (Breitsprecher et al., 2003).
2.4.3. Continental collision and subduction
Alike oceanic subduction, continental subduction is another effi-
cient process of transporting lithospheric material back into the man-
tle. However, its initiation requires high stresses and commonly
continental subduction is facilitated by a preceding oceanic subduc-
tion as it happened in several Phanerozoic orogens of Europe. In
Cenozoic Europe, oceanic subduction processes are thought to be the
forerunner and active carrier of the subducting continental plates in
the Alps, Carpathians, Pyrenees, and beneath Iberia. Numerous region-
al P-wave and S-wave refraction and tomography models indicate the
presence of several subduction zones in the Alpine–Mediterranean
region that extend down to a depth of 200–250 km (e.g. Cavazza et
al., 2004; Piromallo and Morelli, 2003). High-resolution teleseismic
P-wave tomography of the Alps (Lippitsch et al., 2003) suggests the
existence of two subduction zones: beneath the eastern Alps, the con-
tinental Adriatic lower lithosphere subducts northeastwards beneath
the European plate, whereas beneath the western and central Alps,
the continental European lower lithosphere steeply subducts south-
eastwards beneath the Adriatic microplate. Localized linear high-
velocity blocks in the upper mantle interpreted as subducting slabs
have been imaged by seismic tomography beneath the Ligurian-
Tuscany and southern Iberia (Blanco and Spakman, 1993).
A 30 km wide block extending down to a depth of ca. 80–100 km
with 2% lower velocities, which has been imaged by regional P-wave
tomography models beneath the central and eastern Pyrenees, has
been interpreted as subduction of the lower crust of Iberia (Sibuet
et al., 2004; Souriau and Granet, 1995); eclogitization of the lower
crust during its subduction produced weak negative residual gravity
anomalies observed in the Pyrenees (Vacher and Souriau, 2001). Sim-
ilarly, a linear belt of positive residual gravity anomalies along the
TESZ suggests the presence of an ancient subducting slab beneath
the western margin of the EEC as indicated by a regional S-wave to-
mography model (Nolet and Zielhuis, 1994; Zielhuis and Nolet,
1994). However, there is no seismic sign of a subducting slab beneath
the Caucasus (Artemieva et al., 2006).
In contrast to oceanic subduction, subducting continental plates
do not reach the depth of subducting oceanic plates because of an
early start of melting and a smaller slab pull. In the early days of
plate tectonics, the process of continental subduction was not even
considered to be possible because of buoyancy problems associated
with low-density continental crust. Presently, about 20 cases of Phan-
erozoic continental subduction have been confirmed worldwide by
the presence of ultrahigh-pressure metamorphic rocks (UHPM) near
continent–continent collision zones, where rocks of the continental
crust have been subjected to pressures of up to 5 GPa (i.e. transported
to a depth of more than 150 km). Even depths as great as 300 km
have been inferred for subduction of continental material from sam-
ples of garnet peridotite that show the presence of metamorphic
microdiamond and coesite inclusions (Kaneko et al., 2000; Ye et al.,
2000). In particular, UHPM rocks of the Kokchetav Massif of Kazakhstan
indicate subduction of Precambrian crustal protoliths which, at around
530 Ma, were transported to depths of ~200 km (pressures of ~7 GPa).
Subsequent exhumation of these UHPM rocks can occur by different
tectonic mechanisms, for example by wedge extrusion (Boutelier et
al., 2004; Kaneko et al., 2000).
There is no undisputable evidence for the operation of continental
subduction in the early-mid Precambrian. Due to a greater buoyancy
of the continental lithosphere in the Archean–Paleoproterozoic, sub-
duction of oceanic lithosphere, but not of continental lithosphere, is
a more likely process to operate on the early Earth. Seismic reflections
in the cratonic mantle commonly interpreted as relict subduction
zones within the continental lithosphere (BABEL WG, 1990; Warner
et al., 1996)(Fig. 10) could have rather been formed by oceanic sub-
duction during closure of paleooceans. Note, however, that in the
Baltic Sea a dipping seismic reflector crosses the Moho and in the
same place sub-Moho Vp velocities change from 7.8 to 8.2 km/s over
a distance of ca. 25 km, suggesting the operation of continental sub-
duction in Paleoproterozoic (Abramovitz et al., 1997).
2.4.4. Recycling rate at subduction zones
While oceanic slabs have a large variety of dip angles and penetra-
tion depths (Anderson, 2007), not much is known about the shape of
subducting continental plates. The collision of the Indian and Asian
plates provides the most powerful and the best studied example of
the continental subduction (e.g. Kind et al., 2002; Molnar and
Tapponier, 1975; Wittlinger et al., 2004). Seismic images suggest
that, after the subduction of large parts of the Tethyan oceanic plate,
the continental lithosphere of the Indian subcontinent has been (and
still is) pushed below the Himalaya and Tibet down to a 200–250 km
depth in the mantle (e.g. Huang et al., 2003). If the convergence rate
(its part not absorbed by thrusting along the Himalayas front) has
been constantly similar to the present-day rate of ~5 cm/year, about
2500 km of the Indian plate with a lithospheric thickness of ~200 km
may have been recycled in the subduction zone beneath the Tibet
over the past ~50 My.
Despite this huge number, the recycling rate of the continental
lithosphere is much smaller than of the oceanic lithosphere. The con-
tinental lithosphere is thicker, buoyant, and commonly tectonically
protected from subduction zones due to its location within the conti-
nental interior. Recent geochemical constraints suggest that no more
than 30% of the modern mass of the continents has subducted into the
mantle during the entire Earth's history. Furthermore, about 50–70%
27I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
of the subducted continental material has been emplaced back either
as magmatic products into the continental lithosphere or as meta-
morphic material into the lower crust (Coltice et al., 2000). In con-
trast, according to geophysical data, a volume of continental crust
equal to its present-day volume could have been recycled since 2.5 Ga
in case recycling rates in modern subduction zones are representative
of the past rates (Scholl and von Huene, 2007). Presently, the rate of
crustal recycling at non-accreting margins is ~1.3 km
3
/year which is
about an order of magnitude higher than at accreting margins (compare
with Fig. 4b). The recycling rate of the entire lithospheric plate can be
several times higher than for the crust but commonly it is not included
into estimates of recycling rates.
Subduction of continental lithosphere, although less effective than
oceanic subduction and more limited regionally, plays an important
role in the chemical and mineralogical modification of the upper man-
tle. The amount of subducted crustal material depends on the stress
regime and the rheology of the colliding lithospheric plates. Sub-
ducted crustal material may mobilize the asthenosphere and trigger
magmatism as observed in Tibet (Mo et al., 2007), but mantle material
is not always quickly returned to the surface. We discuss later the tec-
tonic similarity between the modern Tibet and the Variscan Europe as
proposed by different authors (Le Pichon et al., 1997; Menard and
Molnar, 1988), and discuss whether subduction-related processes
could produce Paleozoic regional intrusions of basaltic magmas into
the lower crust of west/central Europe.
2.5. Thermo-mechanical recycling of the lithosphere
2.5.1. Lithosphere delamination
Gravitational (or Rayleigh–Taylor-type) instability, which de-
velops at an interface between two viscous materials of different
densities when the lighter material underlies the heavier material,
is responsible for lithosphere delamination. The process which has
long ago been proposed to explain the Cenozoic uplift of the Colorado
Plateau (Bird, 1979) has recently been imaged seismically (Levander
et al., 2011). Importantly, it is likely that lithosphere delamination
always includes delamination of both the lower crust and the litho-
spheric mantle, for example due to viscous coupling between the
crust and mantle as proposed for the on-going removal of the dense
batholithic root beneath the southern Sierra Nevada mountains in
California (Zandt et al., 2004). In some cases, lithosphere (crustal) de-
lamination may result in magmatism and epeirogenesis (Kay and Kay,
1993) such as proposed for some parts of the Andes (McGlashan et al.,
2008).
A density inversion that causes gravitational instability may have
thermal or compositional origin, such as produced by pondering of
hot mantle material beneath the lithospheric base, lithosphere thick-
ening beneath orogenic belts, and phase transitions in the crust
(Fig. 11, see below). Rheologically weak lithosphere and low conver-
gence rate favor development of the Rayleigh–Taylor instability in
collisional settings, where convergence is accommodated by pure
shear thickening. As an orogen is formed, the entire lithosphere is
thickened, and a cold and dense lithospheric root sticks down into
the hotter asthenosphere, creating lateral temperature instability. It
drives convective flow which can remove (delaminate) the lower
portions of the thickened lithospheric mantle. Similarly, the presence
of density inversion also causes delamination of negatively buoyant
lithospheric material, which is swept into descending drips or blobs.
To maintain the mass balance, hotter mantle material replaces dela-
minated lithosphere, causing adiabatic decompressional melting and
triggering magmatism.
The efficiency of lithosphere delamination by Rayleigh–Taylor-type
instability is controlled by lithosphere viscosity. Depending on the
thickness of the lithospheric root, lithosphere delamination requires
lower crustal viscosity between 10
20
to 10
23
Pa s if viscosity of the
upper crust is 10
23
Pa s (Schott and Schmeling, 1998). Numerical
modeling of the development of convective instability associated with
mechanical thickening of cold, dense lithosphere beneath mountain
belts indicates that only the lower 50–60% of the lithosphere may be
delaminated in case of dry olivine rheology but more than 80% of the
lithosphere in case of wet olivine rheology (Houseman and Molnar,
1997). Removal of the lithospheric root in mountain belts built by crust-
al thickening adds potential energy to the lithosphere and causes col-
lapse of orogens by normal faulting and horizontal crustal extension.
Commonly this process starts within ~30 My after the end of the crustal
thickening.
Lithosphere delamination is proposed to drive the Cenozoic tec-
tonic evolution of the Colorado Plateau (Bird, 1979; Levander et al.,
2011), the Sierra Nevada (Harig et al., 2008), the southern Puna pla-
teau of the Andes (Kay and Kay, 1993), Tibet and Himalaya (England
and Houseman, 1989; Turner et al., 1996). In Europe, lithosphere de-
lamination can play an important role in Cenozoic geodynamics of
Fig. 11. Gabbro/basalt–eclogite phase transitions in the crustal rocks. Rainbow shading—
eclogite stability field, colors refer to lithospheric temperatures (purple for cold, red
for hot). Pressure–depth conversion is made assuming crustal density of 2.90 g/cm
3
.
(a) Bold black lines—phase diagram (after Spear, 1993). Shaded area and gray boxes—
extrapolated stability fields of eclogite, garnet granulite, and pyroxene granulite–gabbro
based on experimental data for the quartz tholeiite composition (Ringwood and Green,
1966). Thindashed lines—typicalcontinental reference geotherms(Pollack and Chapman,
1977); numbers—surface heat flow in mW/m
2
. (b) Depth togabbro/basalt–eclogite phase
transition (thick gray line)in different continentalsettings plotted versuscontinental ref-
erence geotherms labeled in heatflow values (after Artemieva, 2011). Tectonic provinces
are markedon the top in accordance with typical heat flow values. Gabbro/basalt–eclogite
phase transition limits crustal thickness to 40–45 km in cold stable platforms and to
~30 km in Phanerozoic basins.
28 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
the Vrancea zone in the Carpathians and the Alboran Sea in western
Mediterranean (Houseman and Molnar, 1997). The presence of,
roughly cylindrical in shape, drip-structures between ~60 and
300 km depth with high seismic velocities is imaged beneath both re-
gions and is interpreted, although not univocally, as delaminated lith-
osphere (Calvert et al., 2000; Koulakov et al., 2010; Wortel and
Spakman, 2000). The process may have also played an important
role in the evolution of the Variscan lithosphere (see Section 3).
2.5.2. Delamination of the lower crust: geochemical evidence
Average major element composition of the continental crust (see
Section 2.2) is directly related to crust–mantle material exchange pro-
cesses. In particular, composition of the upper crust corresponds to
only ~14% partial melting of average new basaltic crust (Hawkesworth
and Kemp, 2006a), whereas typically the upper crust makes ~ 25–35%
of the volume of the continental crust. This misbalance implies that a
significant amount of the lower crust (with the composition of a new
continental crust), from which the upper crustal layer was produced,
is missing and was recycled into the mantle. The residence time of ele-
ments in the lower crust can be ~5 times shorter than those in the
upper crust (Hawkesworth and Kemp, 2006b).
The mechanism of lower crustal recycling may be gravitational
instability of the missing lower crustal portions that originally could ex-
tend down to a 100 km depth as required by volume balance. Elevated
P–T conditions typical for depths of 50–100 km favor phase transitions
that produce heavy minerals, such as garnet, and thus promote recy-
cling of the lower crustal material into the mantle (Kay and Kay,
1991). Delamination of lower crustal material should be more dramatic
in those tectonic settings where crustal growth is most rapid, that is at
convergent margins.
2.5.3. Delamination of the lower crust: geophysical perspective
Gravitational instability develops when continental lithosphere is
in unstable mechanical equilibrium because one of its upper layers
becomes denser and heavier then the underlying layer. Experimental
studies indicate that throughout large regions of the normal conti-
nental crust eclogite is more stable than gabbro and garnet granulite
(Ringwood and Green, 1966). Phase transition of gabbro/basalt to
eclogite in a deep mafic crust can be responsible for delamination of
the lower crust (and in some cases, even of the upper crust, Di
Luzio et al., 2009) since it can produce a significant density inversion
(densities and Vp seismic velocities of gabbro/basalt and eclogite at
the lithostatic pressure and temperatures typical of the lower crust
are ρ~2850–3000 kg/m
3
, Vp~6.2–7.0 km/s and ρ~ 3500 kg/m
3
,
Vp~7.6–8.1 km/s, respectively (Christensen and Mooney, 1995)).
Since the transition from gabbro [plagioclase+pyroxene ± olivine±
spinel] to eclogite [garnet+pyroxene ± quartz] occurs through the
garnet granulite [garnet+pyroxene(s) + plagioclase](Fig. 11a), the
increase in density and seismic velocity from gabbro to eclogite
(caused by an increase in the proportion of garnet and a decrease in
the proportion of plagioclase with a pressure increase) is uniformly
smeared throughout the garnet granulite transition interval (with
ρ~3150 kg/m
3
and Vp~6.8–7.2 km/s at lower crustal P–T conditions).
Given that eclogitic material is denser than the underlying mantle
peridotite (density of the latter is 3300–3400 kg/m
3
), the densified
lower crust starts sinking, for example by crustal “dripping”into the
mantle (Gögüs and Pysklywec, 2008), when a significant amount of
eclogite is formed. Thus, delamination of the lower crust limits the
thickness of the (petrological) continental crust by the depth where
the lower layer transforms into dense eclogite- or garnet–pyroxenite
facies.
Gabbro/basalt–eclogite phase transition typically starts at depths
30–45 km (Mengel and Kern, 1992) with significant regional depth
variations due to a strong temperature‐dependence of the reaction
(Fig. 11). However, the actual P–T stability field of the system at
low temperatures is unknown because the rate of gabbro/basalt–
eclogite transformation is very slow and also strongly temperature-
dependent. For this reason, laboratory experiments have been made
at temperatures T >800–1000 °C (Ito and Kennedy, 1971; Ringwood
and Green, 1966), and then extrapolated to lower temperatures. Al-
though the real rate of the transition is unknown, it is sensitive to
both temperature and the presence of fluids, so that even trace
amounts of water may lead to rapid eclogitization (Austrheim et al.,
1997), such as expected in subducting slabs where metamorphic
transformation of the basaltic and gabbroic oceanic crust to eclogite
occurs on the order of 10
6
–10
8
years, that is at distances of
10–100 km along the slab for typical subduction rates (Ahrens and
Schubert, 1975). The composition of basalts has some effect on the
phase transition pressures and temperatures; for example experi-
mental studies indicate a pressure difference of up to 5 kbar for the
gabbro–garnet granulite transition in alkali olivine basalts and tholei-
ite basalts (Ringwood and Green, 1966).
In cratonic regions with exceptionally cold lithospheric tempera-
tures (Fig. 5), the basalt–eclogite transition is expected to occur only
at depths >40–45 km (Fig. 11b). These values are in a perfect agree-
ment with global seismic observations that show that crustal thick-
ness in Precambrian shields is limited to ~50–55 km with an average
thickness of ~40–45 km: deeper crustal roots become gravitationally
unstable with phase transition advance and are mechanically delami-
nated. In some cases, some eclogite material can still be present in the
deeper parts of the crust, but it would be seismically indistinguishable
from the mantle material (in particular, because not 100% of the lower
crustal rocks transforms to eclogite). Due to the extreme slowness of
the gabbro–eclogite phase transition at low temperatures, thick
seismically-distinguishable metastable crustal roots may be occasion-
ally preserved in cold and dry intracratonic environments, like in cen-
tral Finland, where the crustal root extends down to 55–60 km, well
into the eclogite stability field. Lower crustal xenoliths from the Kola
Province of the Baltic shield confirm the presence of variably retro-
gressed and metasomatized eclogites at depths between 30 and
48 km (the deepest sampled depth) (Downes, 1993). Although it is
unknown if gabbro can remain metastable for a billion years, the
time scale of processes in Precambrian terranes, solid state diffusion
cannot produce the transition in geologically meaningful times
under low temperatures (b600–800 °C) and at anhydrous conditions
that are thought to be typical of the cratonic lithosphere (Ahrens
and Schubert, 1975).
At high crustal temperatures, typical for Phanerozoic regions,
gabbro/basalt–eclogite phase transition requires time on the order of
million years (Artyushkov, 1993). For this reason, mafic crustal roots
deeper than 40–45 km in young mountain ranges are transitory,
metastable features that are either being formed during on‐going con-
tinental collision/subduction or existing as long as the age of the
newly‐formed crustal roots is younger than the time‐scale of the
gabbro‐eclogite phase transition. Additionally, the timing of lower
crustal eclogitization may be controlled not only by crossing of the
phase transition depth, but rather by the time when fluids are intro-
duced into the anhydrous system (Austrheim et al., 1997). In particu-
lar, crustal doubling without eclogitization may exist in collisional
orogens with a dry crust such as proposed for the southernmost
Tibet where low seismic Vp (~ 6.3 km/s) in the underthrusting Indian
crust do not indicate crustal eclogitization (Zhao et al., 1993).
Due to a peculiar shape of gabbro/basalt–eclogite phase transition
on P–T diagram, metamorphism can start already at ~30 km depth in
young platforms with moderate crustal temperatures, such as in the
Variscan Europe (Fig. 11b). A surprisingly uniform thickness of
28–32 km typical of the European Variscan crust (see Section 3.2.4)
is in a striking accord with the depth of eclogitization; yet other expla-
nations for the thin Variscan crust are also plausible. In contrast, due to
the shape of basalt stability field, young tectonic regions with high
lithospheric temperatures may have stable lower crustal roots down
to a 35–40 km depth.
29I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
2.5.4. Thermal and thermo-mechanical erosion
Two other thermo-mechanical mechanisms to destroy (and de-
laminate) the lithosphere include (i) thermal erosion by mantle
plumes or convection and (ii) thermo-mechanical erosion by stirring
and friction (basal drag) associated with relative movement of the
lithosphere over the underlying mantle (e.g., Artemieva and
Mooney, 2002; Sleep, 2003). Similar to gravitational instability, the ef-
ficiency of these processes strongly depends on lithosphere rheology.
Particular physical and chemical conditions required for preservation
of 2.5–3.0 Ga old cratonic lithospheric roots that may extend down
to at least 250–300 km depth (Artemieva and Mooney, 2001;
Kustowski et al., 2008; Lévy et al., 2010) have been a subject of numer-
ous studies. Numerical models of mantle convection indicate that the
continental lithosphere can sustain against convective erosion if litho-
sphere viscosity is three orders of magnitude higher than that of the
convective mantle (Lenardic and Moresi, 1999; Shapiro et al., 1999).
Dry depleted composition of cratonic lithosphere as proposed by
Pollack (1986) can provide the high viscosity contrast required for
the cratonic lithosphere to maintain large equilibrium thickness
(Doin et al., 1997). Thermo-mechanical erosion is thought to be the
cause of the missing cratonic lithosphere beneath the North China cra-
ton (Fan et al., 2000). However, thermal erosion of the continental
lithosphere contributes not only to recycling of crustal material into
the mantle (Gao et al., 2004; Xu et al., 2004), but causes mafic magma-
tism which leads to crustal growth (Section 2.3).
3. Peculiarities of crustal velocity structure in central and Western
Europe
3.1. Making Paleozoic Europe
To facilitate the following discussion, we start with a brief summa-
ry of the European tectonic evolution in the Paleozoic. The Caledo-
nides form the oldest of the Paleozoic mountain belts in Europe
(Fig. 12). After the closure of the Iapetus Ocean, a complex conti-
nent–continent triple plate collision (Baltica, Laurentia, and Avalonia)
began in the Ordovician–Silurian time (McKerrow et al., 1991)
(Fig. 13a). Large parts of west-central Europe, including southern
Great Britain, the southern North Sea, and parts of northern Germany
(the German–Polish Caledonides) are thought to be remnants of East
Avalonia (Thybo et al., 2002)(Figs. 12, 14). The Caledonian conti-
nent–continent-collision, that lasted until 400–390 Ma, was accom-
panied by significant continental subduction as indicated by seismic
data and the presence of numerous eclogites and ultrahigh pressure
(UHP) rocks in Norway (Boutelier et al., 2004); some of these rocks
were exhumed from a depth of more than 100 km (Johnston et al.,
2007). The collision resulted in the Caledonian–Appalachian Orogeny
as a result of which the Caledonian foldbelt was formed in western
Europe and East Greenland and the Appalachian orogen was formed
in eastern North America (Fig. 14).
The Variscan Orogen is a broad, 500 to 1000 km wide, mountain
belt which stretches NE–SW across most of west-central Europe for
about 3000 km and runs from Portugal and Spain to Poland and
England (Fig. 15a). It was formed mainly between 430 and 280 Ma
as a result of the closure of the Rheic Ocean and the subsequent colli-
sion of the Gondwana and Laurasian continents (Fig. 13b). At around
400 Ma Avalonia collided with Baltica (Fennoscandia and Denmark)/
Laurentia (North America, Greenland and northern Britain), forming
the new supercontinent Laurasia (Lliboutry, 1999; McKerrow et al.,
1991; Tanner and Meissner, 1996).
Although it is clear that oceanic and, perhaps, continental subduc-
tions played a key role at the initial stages of the formation of the
Variscan orogen, how the convergence in the Variscan orogeny really
took place is much debated. It is unknown if it resulted in the forma-
tion of huge plateaus and high mountain ranges or only in relatively
small mountain belts (Franke, 2006; Oncken, 1998). Strong seismic
anisotropy in the lithosphere and some mantle xenoliths indicate a
complicated pattern in the closure of oceanic domains and the conse-
quent formation of the Variscan orogen and the associated subduc-
tion zones (Babuška et al., 2002; Christensen et al., 2001; Downes,
1993; Fuchs and Wedepohl, 1983; Plenefisch et al., 2000; Wilson et
al., 2004). The total crustal shortening during the Variscan orogeny
exceeds 600 km. The terranes of Proterozoic to Carboniferous ages,
including the Armorican, Iberian, Bohemian, Ardennes, and Massif
Central (Fig. 15b) were deformed and partly metamorphosed. Large
volumes of granitoides have been emplaced between 370 and
280 Ma (Matte, 1986); a few eclogitic rocks were found in the Sax-
othuringian sub-terrane (Franke, 1994). The present-day Variscan
structures are represented by the Rhenohercynian, Saxothuringian
and Moldanubian sub-terranes, separated by subparallel NE–SW-
striking sutures (Fig. 15a). The sub-terranes have distinctly different
lithospheric structure which is discussed in the next section.
3.2. Crustal thickness in Europe
We next review the major characteristics of the Paleozoic crust of
western Europe to discuss how it could have been formed and which
of the crust–mantle material exchange processes played the dominant
role in its structural evolution. A detailed review of crustal structure of
Europe is outside the scope of the present study. Recent summaries
can be found in Pavlenkova (1996),Ziegler and Dèzes (2006),
Artemieva (2007),Tesauro et al. (2008),Artemieva and Thybo
(2012). The general tectonic evolution and the present-day structure
of the European lithosphere is summarized in Blundell et al. (1992)
and Artemieva et al. (2006). Here we would like to emphasize only
few points related to the present discussion, with an emphasis on
the Paleozoic crust.
Typical continental crust has an average thickness of ca. 40 km
(Christensen and Mooney, 1995). Although this number produced
by worldwide averaging over tectonic provinces with diverse tectonic
origin should not necessarily be meaningful, it provides some refer-
ence frame for global comparisons. In general, any continental crust
significantly thinner or thicker than 40 km is usually considered as
“anomalous”. In this section we demonstrate that most of the conti-
nental crust in Europe is “anomalous”, either in terms of comparisons
with aglobal average of ca. 40 km, or with expectations coming from
comparisons with tectonic analogues of the same ages, or both. In
particular, continental crust with a 40 km thickness is largely absent
in Europe (Fig. 16).
3.2.1. Crustal roots in young collisional orogens
Deep crustal roots are commonly present beneath young collision-
al orogens worldwide. In Europe they have been seismically imaged
beneath the Alps, the Carpathians, and the Pyrenees where they are
parts of subducting continental slabs (Artemieva et al., 2006 and ref-
erences therein). The formation of these orogenic belts and associated
subduction zones is related to the convergence of the Eurasian and
African plates which resulted in the collision of the Adriatic and the
Iberian terranes with the Variscan continent, producing the main up-
lift of the Alpine belt at ~23 Ma (Castellarin and Cantelli, 2000;
Schmid et al., 1996). Although the pre-orogenic history of the Alps
and the Pyrenees is different, the deep seismic structure looks similar
and both mountain belts have developed deep crustal roots, extend-
ing down to 55–60 km (Behm et al., 2007; Lippitsch et al., 2003;
Pfiffner, 1990; Willingshofer and Cloetingh, 2003).
Beneath the Alpine mountain belt, the lower crust is felsic (with
low P-wave seismic velocities) and seismically laminated, apparently
as a heritage of the subducted Variscan crust (see below) (Figs. 17a,
18). The downward curvature of the lower crust towards the center
of the mountain belt has been interpreted as an indicator that its age
is older than the last orogenesis (TRANSALP Working Group, 2002).
In counter play with downward material transport by subduction
30 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
processes, magmatic activity in the mountain belts provides some up-
ward material transport as well. The famous Ivrea body in northern
Italy is an intrusive mafic complex with P-velocities of more than
7.4 km/s. The body reaches the surface where it is surrounded by “hy-
brid”rocks which have apparently assimilated some crustal material
during their ascend. The composition of the magmas suggests they
originated in the upper mantle with perhaps some input from the low-
ermost crust (Sinigoi et al., 1991; Voshage et al., 1990).
In the Carpathians, the origin of which is also, at least in part, asso-
ciated with the African–Eurasian plate collision so that some authors
consider them to be an eastward extension of the Alps (Cloetingh et
al., 2004), the subduction-related orogenic crustal root is less pro-
nounced than beneath the Alps and the Pyrenees with Moho at
39–43 km depth (Sroda et al., 2006). Today the Carpathians display a
wide range of strong crust–mantle interaction processes, which in-
clude a complex interplay between the eastward extrusion of the
Fig. 13. Sketch illustrating Paleozoic assemblage of Laurasia and the formation of the Caledonian and the Variscan orogens (after Torsvik et al., 1996).
Fig. 12. Tectonic map of Europe (modified after Artemieva et al., 2006). Dashed lines—crustal cross-sections discussed in Fig. 19 (profile C–V follows a 50 N latitude from the Atlantic
margin (Cornwell) to the Voronezh massif; profile I–K extends from the Iberian massif to the Kola province). Black diamonds—DEKORP/Basin profile (see Fig. 19e). Black thin lines—
seismic profiles discussed in Fig. 18 (1—TRANSALP, 2—Swiss, 3—ECORS). TESZ = Trans-European Suture Zone, VP = Voring Plateau, PB = Pannonian Basin, S = Skagerrak area.
31I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
Alps, various steep subduction segments and an asthenospheric up-
welling beneath the Pannonian basin, below which some remnants
of slabs are also observed (Dando et al., 2011).
Similar to the Carpathians, there is no clear seismic evidence for
the presence (or the depth extent) of the crustal roots beneath the
Dinarides, the Apennines, and the Caucasus although compressional
tectonics and high topography suggest their presence. Beneath the
Apennines the crust is apparently very thin (30–35 km), while be-
neath the Dinarides is can be 40–45 km thick. The deep crustal struc-
ture beneath the Caucasus is unknown due to the lack of seismic data,
but the presence of the subducting slab with associated crustal root is
suspected from a strong negative gravity anomaly.
3.2.2. Cratonic crustal roots
Several unusual crustal structures with well preserved crustal
roots extending deeper than 50 km and locally reaching 55–60 km
are documented in the cratonic parts of Europe (Figs. 16, 19a, c).
While the presence of a deep crustal root at the Archean–Svecofen-
nian suture in the Baltic Shield is well known (e.g. Korja and
Heikkinen, 2005; Korja et al., 2009), it is not broadly acknowledged
that similar crustal roots, down to a 55–60 km depth, are also present
beneath some terranes of the Volga-Uralian sub-craton and the Ukrai-
nian shield (e.g. Trofimov, 2006). The age of these crustal roots can be
as old as Paleoproterozoic and their origin remains speculative
(Artemieva et al., 2006; Artemieva and Thybo, 2008). However, it is in-
triguing if such deep crustal roots extending down into the eclogite
stability field can be preserved for more than a billion years. As dis-
cussed in Section 2.5.3, in the cold and dry cratonic lower crust meta-
morphic reactions can be very slow because of a critical dependence of
the rate of gabbro/basalt–eclogite phase transition on temperature
and fluid regime, but lack of experimental data leaves the question
open, if gabbro can remain metastable for a billion years.
3.2.3. Crustal roots in Paleozoic orogens: Uralides
By analogy with modern zones of continent–continent collision , deep
crustal roots extending down to more than 50 km should be expected to
be formed in the Paleozoic orogens as well. In particular, the Urals oro-
gen has preserved a thick crustal root until present (Fig. 20c). The
Uralides orogeny took place in the Palaeozoic (at 450–250 Ma) and the
orogen was formed by accretion of series of island arcs, volcanic com-
plexes and fragments of microcontinents to the eastern edge of the
East European craton during the collision of the European and the
Kazakhstan plates starting in the Early Ordovician–Carboniferous time
(Sengör et al., 1993; Zonenshain et al., 1990). Although much of the to-
pographic elevation is lost (the modern topography of the Urals came
into existence only during the Tertiary–Quaternary (Morozov, 2001)),
the orogen has a pronounced, more than 50 km thick, maficcrustal
root with very high average crustal velocities (Berzin et al., 1996;
Juhlin et al., 1998), reaching down to 65 km in the Polar and in the
Southern Urals (Carbonell et al., 1996; Druzhinin et al., 1990). The pres-
ence of the crustal root is considered as an evidence of missing post-
orogenic extension (Steer et al., 1998).
In the southern part of the orogen, along the URSEIS seismic profile,
lower crustal velocities of more than 7.8 km/s have been observed;
they transform gradually to velocities of ca. 8.0 km/s (Echtler et al.,
1996) interpreted as evidence for eclogitic crustal root. The survival
of the deep crustal root below the depth of gabbro/basalt–eclogite
phase transition since Paleozoic (i.e. for several hundred million
years) raises questions of its mechanical stability. Southern Urals, in
the part where a deep crustal root is observed, is characterized by
anomalously low surface heat flow (ca. 25 mW/m
2
,Kukkonen et al.,
1997). Although, various reasons for the heat flow anomaly have
been proposed (e.g. paleoclimatic signal, low crustal heat production
anomaly, groundwater circulation), one cannot rule out that a part
of the anomaly has a deep origin and that the lower crustal tempera-
ture is very low, thus providing long-term mechanical stability of the
crustal root. In particular, thermal models with a low crustal heat
Fig. 14. The Appalachian–Caledonian belt, shown in a Pangea fit (after Cocks et al.,
1997; McNamara et al., 2001).
Fig. 15. (a) Variscan massifs (in blue) and megazones of the Variscan orogen (after Bromley and Holl, 1986). HF = Hercynian Front. (b) Map of Variscan massifs; massifs with sim-
ilar geological features are shown by the same colors. The Iberian Peninsula is rotated to its Paleozoic position (modified after Martinez Catalan, 1990). 1—allochthonous terranes of
various origins including the continental margin of Gondwana, non-Gondwana domains and oceanic crust; 2–6—allochthonous or parallochthonous domains of the Gondwana
realm; 2—Saxothuringian zone; 3—Central Armorican and Central Iberian zone, the Pyrenees; 4—Moravo-Silesian and Rhenohercynian zones, SW England and Ireland, and the
South Portuguese zone; 5—West-Asturian-Leonese zone; 6—Cantabrian zone. Abbreviations: AM = Armorican Massif; AR = Ardennes Massif; B = Brabant Massif; BF = Black For-
est (Rhine Graben); BO = Bohemian Massif; CW = Cornwall; Hz = Harz Massif; IM = Iberian Massif; MC = Massif Central; MS = Moravo-Silesian Massif; PY = Pyrenees; RM =
Rhenish Massif; SPZ = South Portuguese zone; V = Les Vosges (Rhine Graben).
32 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
production in island arc complexes of the Uralian crust suggest Moho
temperatures (at a depth of ca. 60 km) of 550–600 °C (Kukkonen et al.,
1997). However, the presence of even trace amounts of water, which
one can expect to be still present in paleo-subduction zones, would
significantly accelerate the rate of gabbro/basalt–eclogite phase tran-
sition (Ahrens and Schubert, 1975; Austrheim et al., 1997).
Fig. 16. Structure of the European crust shown as deviation of the Moho depth from 40 km depth. The map is based on a 5 deg× 5 deg interpolation of all published seismic data
smoothed with a low-pass filter. Data sources: Pavlenkova (1996),Ziegler and Dèzes (2006),Artemieva (2007),Kelly et al. (2007),Artemieva and Thybo (2012). Dashed line—the
TESZ.
Fig. 17. Vp-velocity structure of the European crust. The map is based on a 5 deg× 5 deg interpolation of all published seismic data; in case of multiple interpretations of the same
data, more recent results were used. Data sources: Pavlenkova (1996),Ziegler and Dèzes (2006),Artemieva (2007),Kelly et al. (2007),Artemieva and Thybo (2012). Dashed line—
the TESZ. (a) Thickness of the mafic lower crust (crustal layers with Vp >6.8 km/s); (b) variations in average Vp seismic velocities (recalculated to room P–T conditions) in the
consolidated crust in Europe. Zero corresponds to average in situ Vp =6.6 km/s in a region with a platform geotherm (surface heat flow ~55 mW/m
2
). The correction to lateral tem-
perature variations is based on the regional thermal model (Artemieva, 2007) for ∂Vp/∂T=−0.39 e−3 m/s/K (the value reported by Christensen and Mooney (1995) for granite
and basalt). Dashed line—the TESZ.
33I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
3.2.4. Missing crustal roots in Caledonides and Variscides
Compared to most other Paleozoic mountains (the Caledonides,
the main Appalachians, the Variscides), the Uralides are unique in
preserving the crustal root (Fig. 20cd). Although, similar to other
continent–continent collision zones, crustal roots should have been
formed underneath the Caledonian and the Variscan mountain belts
in the Paleozoic, at present they are absent in the Norwegian Caledo-
nides, the Caledonides of Scotland and North America, and the Varis-
cides (Clegg and England, 2003; Cook et al., 1988; Stratford et al.,
2009)(Figs. 16, 19, 20). A local crustal root with intermediate veloc-
ities is still preserved only in the very south of the Appalachians
(Hawman, 2008). However, in the Newfoundland Appalachians the
crust is thin (Fig. 20d) and seismic data were interpreted as evidence
for partial eclogitization (Chian et al., 1998). Since the Urals orogen
was not easily accessible for geologic and geophysical studies to
western researchers and international publications on its structure
remained scarce until 1990-ies, it has long been believed that a thin
crust such as in the Caledonides, Variscides, and the northern Appala-
chians (Behr and Heinrichs, 1987;Meissner, 1986;Nelson, 1992)is
typical for all Paleozoic orogens.
3.2.4.1. European Caledonides. Today the crust in the Norwegian Cale-
donides shows no sign of the orogenic crustal root and gradually
thickens eastwards towards the Baltic Shield from about 28 km at
the coast of Norway (Ottemoeller and Michs, 2004) to ca. 40 km be-
neath the southern dome (Stratford et al., 2009). Similar crustal thick-
ness, of about 36–40 km, is typical for the northern Appalachians
(Hughes and Luetgert, 1991)(Fig. 20d). In southern East Greenland,
within the Caledonian deformation zone, crustal thickness calculated
by receiver functions varies between 24 and 32 km (Dahl-Jensen et
al., 2003; Kumar et al., 2007). Similar to the Newfoundland
Appalachians (Clowes et al., 2010), no crustal roots are known in
the Irish and British Caledonides, where the crustal thickness is
30–35 km (see summary by Kelly et al., 2007) and reduces to less
than 20 km at the Atlantic shelf area (Roberts et al., 2009).
The German–Polish Caledonides form a relatively narrow belt at
the northern margin of the Variscan structure (Fig. 12). Toward the
end of the Carboniferous, the northward Variscan compression initiat-
ed orthogonal extensional stresses and created numerous rifts and ba-
sins with magmatic intrusions in the vast area which extends from
northern Germany and Denmark to the southern coasts of Norway
and Sweden (Bayer et al., 1999; Marotta et al., 2001; Thybo, 1997;
Ziegler and Cloetingh, 2004). The formation of these rifts and basins
started long after the Caledonian orogeny had finished and the
Variscan orogeny had started (Heermans et al., 2004). In the northern
foreland of the Variscides the most prominent Paleozoic extensional
structures are the Oslo Graben, the Skagerrak area, and the North
German Basin. Huge amounts of mafic magmas have entered the
stretched lower crust in all Paleozoic rifts and basins at the northern
margins of the Caledonian–Variscan orogen, sometimes producing re-
markable gravity highs as at the border between the Danish- and the
North German Basins (Sandrin and Thybo, 2008). Seismic data from
the northern edge of the former Caledonian orogeny in Denmark
show high sub-Moho velocities interpreted as the presence of a possi-
bly eclogitic material (Abramovitz et al., 1998).
In the North German Basin, quick subsidence began in the Latest
Carboniferous–Earliest Permian, preceded by extensive faulting and
magmatism, and followed by several tectonic events which led to for-
mation of Meso-Cenozoic sub-basins (Heermans et al., 2004; Ziegler
and Dèzes, 2006). Thickness of sediments in the North German
Basin reaches 10 km and its sub-basins contain a thick (up to 9 km
in thickness) layer of the Upper Permian (Zechstein) salt. Seismic
Fig. 18. Present subduction in the Alps and the Pyrenees (simplified from Choukroune, 1989; Gebrande et al., 2006; Pfiffner et al., 2001; Souriau and Pauchet, 1998; TRANSALP
Working Group, 2002; Willingshofer and Cloetingh, 2003). Laminated lower crust (and lithospheric mantle) is partly broken, partly introduced into mantle. Profile locations are
shown in Fig. 12.
34 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
data indicate the presence of massive high-velocity mafic intrusions
in the lower part of the crust and a flat and shallow (possibly new)
Moho at a depth of 30–32 km, as in the adjacent Variscan terranes
(Bayer et al., 1995; Bilgili et al., 2007; Rabbel et al., 1995; Wilson et
al., 2004)(Figs. 19, 20a).
3.2.4.2. Variscides. The crustal structure of the Variscan orogen is truly
unusual.
(i) The crust in the Variscan terranes of Europe is surprisingly
thin, ca. 28–32 km only (Figs. 16, 19ac, 20a). Only some of
the Paleozoic massifs show slightly larger crustal thicknesses
(up to 35 km). In continental regions, similar crustal thick-
ness is observed at present only in the recently extended
regions, such as the Cenozoic Basin and Range province in
western USA (Fig. 20b), where recent teleseismic receiver
function studies indicate Moho depths variations from
29.5 km in the east to 36.5 km in the west; wide-angle refrac-
tion/reflection survey for the same region gives comparable
Moho depths of 32–37 km (Gashawbeza et al., 2008). By com-
parison, the crust of the East European craton is typically ca.
15–20 km thicker (Fig. 16).
(ii) The Variscan crust has a rather uniform thickness with only
small regional variations (Fig. 19a, c, e). An important charac-
teristic of the present-day Variscan crust is lack of relationship
between the crustal structure and pre-existing terrane
boundaries. The Moho is sharp and subhorizontal over the en-
tire western Europe, even where it crosses the major sutures of
the Variscan orogen (Figs. 16, 19e).
(iii) The Variscan terranes have a significantly reduced thickness of
the mafic crust, which commonly is 3–5 km thick as compared
to 10–25 km thick mafic layer of the cratonic crust (Figs. 17a,
19). As a consequence, low average basement velocities are
typical of the Variscan crust (Fig. 17b and Section 3.3).
(iv) A characteristic feature of the Variscan orogen is a widespread
reflectivity in the lower crust (Fig. 19e and Section 4.1).
Such a crustal structure, with a thin crust, generally with a seismi-
cally laminated lower crust and a sharp subhorizontal Moho which
ignores the pre-existing tectonic elements, is often interpreted as an
indication that a large part of the lower crust, and probably of the lith-
ospheric mantle, has been delaminated after the main compressional
events took place in the Paleozoic orogenies. These processes are dis-
cussed in detail later.
The Trans-European Suture Zones (TESZ), one of the most important
tectonic boundaries in Europe which separates the Variscan terranes
and the German–Polish Caledonides from the Precambrian, mainly
Archean–Paleoproterozoic in age, East-European craton (Fig. 11),
marks a sharp change in the crustal structure from a thin and rather uni-
form crust of western and central Europe to a 45–50 km thick cratonic
crust with the “classical”three-layered structure, including a high-
velocity (mafic) lower crust (EUROBRIDGE SWG, 1999; Korja and
Fig. 19. (a–d) Two cross-sections through the European crust constrained by all available seismic data averaged within 600 km-wide corridors along the profiles (the locations of
the profiles are marked in Fig. 12). Upper plots (a, c) show the subdivision of the lithosphere into compositional layers as based on P-wave seismic velocities (Mengel et al., 1991;
Wedepohl, 1995): granites and gneisses (upper crust) Vp b6.4–6.5 km/s; felsic granulites (middle crust) Vp~ 6.4–6.8 km/s; mafic granulites (lower crust) Vp ~ 6.8–7.2 km/s; pyrox-
enites and eclogite (lowermost crust) 7.2–7.6 km/s; spinel lherzolites and harzburgites (lithospheric mantle) Vp >7.8 km/s. For data sources see Pavlenkova (1996),Artemieva et
al. (2006),Ziegler and Dèzes (2006),Artemieva (2007),Kelly et al. (2007),Artemieva and Thybo (2012). Lower plots (b, d) show variations in mean P-wave velocity in the base-
ment of the European crust (i.e. the crust without the sediments) based on seismic data. Dashed lines refer to in situ conditions (as sampled by seismic methods) and reflect var-
iations in both crustal composition and average crustal temperatures. Solid lines—Vp variations corrected for lateral temperature variations in the crust (based on Artemieva, 2003,
2006), which reflect variations in the average crustal composition and anisotropy (in case it is present). Zero corresponds to average in situ Vp =6.6 km/s in a region with a platform
geotherm (surface heat flow ~55 mW/m
2
). TESZ = Trans-European Suture Zone; DDR = Dnieper-Donets paleorift; NGB = North German basin. (e) P-wave seismic velocity struc-
ture of the European Variscides and Caledonides (North German Basin) along the profile DEKORP/BASIN 9601. Seismic velocities are derived from wide-angle seismic data and
shown in relation to the line drawing of the seismic reflection data (based on Bayer et al., 1999). Profile location is shown in Fig. 12.
35I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
Heikkinen, 2005; Pavlenkova, 1996 and many others) (Figs. 19c, 20c).
However, in intracratonic zones of paleoextension (Riphean to Devoni-
an paleorifts within the craton interior) the crustal structure resembles
the Variscan–Caledonian crust (Fig. 17a): Moho depth reduces to
35–40 km and becomes dominated by a thick (>20 km) upper crust,
with the middle crust being almost absent and the lower crust signifi-
cantly thinned (Artemieva, 2007; Kostyuchenko et al., 1999).
3.2.5. Meso-Cenozoic rifts
One of the most prominent Mesozoic rifts of Europe is the rift sys-
tem of the North Sea, which includes the Viking Graben in the north
and the Central Graben in the south. Both of them are deeply buried
under thick Tertiary sediments. The formation of the North Sea rifts
has probably begun at the late stages of the Caledonian orogeny, but
the major phase was connected to Mesozoic rifting at the Atlantic pas-
sive continental margin (Ziegler, 1992). Mesozoic rifting, which began
in Triassic–early Jurassic, continued for an unusually long time of ca.
175 Ma. The subsequent post-rift thermal subsidence, partly associat-
ed with delayed thermal reactions due to late metamorphic reactions
in the uppermost mantle, occurred during the Tertiary (Vejbaek,
1990; Ziegler, 1992). A vast amount of data on the crustal structure
of the North Sea (see e.g. summaries by Klemperer and Hobbs, 1991;
Ziegler and Dèzes, 2006) indicates crustal thickness of ca. 25–30 km,
which reduces to 24–26 km along the Viking Graben (central North
Sea).
Alike European Caledonides, a large part of the Variscides has been
reworked by Meso-Cenozoic events as a result of large relative move-
ments of the Eurasian and African plates and tectono-magmatic activity
in the Central European Rift System. The cause of Cenozoic volcanism in
the latter (which includes the Rhinegraben, western Rhenish Massif
and Massif Central) is still much debated. The proposed mechanisms
include plume-related active rifting, passive rifting in response to colli-
sional processes in the Alps and Pyrenees, back-arc rifting, or slab pull
associated with Alpine subduction. Similar to the rift system of the
North Sea, the crust in the Central European Rift System is thin, 24 to
30 km (e.g. Ziegler and Dèzes, 2006)(Fig. 16).
3.3. Seismic evidence for lower crustal delamination in the Variscides
3.3.1. Mean crustal velocity in western Europe
Seismic velocity structure of the crust in continental Europe is
highly heterogeneous due to variations in crustal structure, composi-
tion and in thermal state of the crust (Fig. 17); additionally some part
of crustal velocity variations can be caused by anisotropy. The crust of
the East European craton with mean basement velocities typically be-
tween 6.4 and 6.6 km/s (Figs. 19c, 20c) provides a convenient region-
al reference frame for examining the anomalous structure of the crust
of western Europe, since mean basement velocities of ca. 6.5 km/s are
statistically representative of stable continental crust (Christensen
and Mooney, 1995). [By crustal basement we mean the consolidated
Fig. 20. Profiles of a typical 1-D P-wave velocity structure with depth extracted from seismic velocity models for the continental crust of Paleozoic Europe and North America: (a) for
three Variscan sub-terranes of western Europe (RH = Rhenohercynian zone, ST = Saxothuringian zone, and MD = Moldanubian zone) (Blundell et al., 1992) and differen t parts of
the North German Basin (Bayer et al., 1999; DEKORP-BASIN RG et al., 1999; Tryggvason et al., 1998); (c) for Southern Urals (the axial zone, Carbonell et al., 1996;Morozov, 2001);
(d) for Caledonides of Southern Norway (Stratford et al., 2009); northern and southern Appalachians (Taylor, 1989b), and Newfoundland Appalachians (Clowes et al., 2010). Ve-
locity–depth profiles for extensional crust of the Basin and Range Province (western USA, Lerch et al., 2007) and typical crust of the East European cratons are shown for a com-
parison (b–c). The velocity–depth range typical for the Variscan Europe is shown in (b–d) by red shading.
36 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
crust, i.e. the crust without the sedimentary layer]. The discussion of
the details of the velocity structure of the East European craton is out-
side the scope of the present review and can be found in Artemieva
(2007).
The transition from the stable cratonic crust to Variscan and Cale-
donian crust of western Europe is marked by a sharp and pronounced
contrast in mean basement velocity constrained by various seismic
surveys (Fig. 19d). New seismic data from the Norwegian Caledonides
indicate mean basement velocity of ca. 6.5–6.6 km/s beneath most of
the Norwegian mountains (Stratford et al., 2009). However, in their
southern, most uplifted part, mean velocity is low (6.3–6.4 km/s)
suggesting that a large part of the lower crust, and probably of the lith-
ospheric mantle, has been delaminated during or after the Paleozoic
orogeny (Figs. 17ab, 20d). Recent and still on-going uplift of the south-
ern Norwegian dome (Anell et al., 2009; Japsen and Chalmers, 2000)
provides an indirect evidence for the timing (Neogene, and perhaps
even Plio-Pleistocene) when the crustal root of the Norwegian Caledo-
nides may have been lost.
Similar to the southern part of the Norwegian Caledonides, mean
basement crustal velocity in the Irish and British Caledonides is
6.2–6.4 km/s (Kelly et al., 2007). Alike European Caledonides (no seis-
mic data exist so far for the Caledonides of Greenland), low mean
basement velocities (6.2–6.4 km/s) are typical for the Variscan crust
of Europe. The Central European rift system which extends from
Massif Central to the Central Graben of the North Sea has extremely
low mean basement velocities (b6.3 km/s) (Fig. 17b).
Due to highly variable crustal geotherms (owing to variations in
crustal heat production, geodynamic regime, heat flow from the man-
tle, the presence of recent magmatism, Fig. 5), the contribution of tem-
perature variations into seismic velocity variations recorded in seismic
surveys can be significant. Average crustal temperatures in the conti-
nental crust of Europe may differ by up to 200 °C (Artemieva, 2007);
note that the variation could have been much larger if it were not for
a generally inverse correlation between the thermal regime of the
crust and crustal thickness in Europe. When the effect of temperature
variations is removed, mean P-wave velocities in the basement reflect
primarily structural and compositional variations in the crust
(Figs. 17b, 19b, d). For example, in the Pannonian Basin in situ seismic
velocities in the crustal basement are similar to the Variscides; howev-
er, temperature-corrected Vp seismic velocity is ca. 0.2 km/s higher
than typical velocities in the Variscan basement, suggesting that Ceno-
zoic magmatism controls the present-day velocity and density struc-
ture of the Pannonian crust.
The general patterns of the mean basement velocity variations are
similar at in situ and at temperature-adjusted conditions, but the am-
plitudes of the anomalies are stronger in the latter case. As with in situ
velocities, the effect of crustal anisotropy can also be present. Howev-
er, we omit the latter from the discussion due to insufficient seismic
data for many tectonic provinces of western Europe. Several regions
can be mentioned due to distinct anomalies in crustal velocity struc-
ture. The mean velocity structure of the Variscides is apparently rela-
tively homogeneous with a ca. −0.4 km/s Vp anomaly as compared
to the stable platform; similar low velocities are typical of the
British–Irish Caledonides and southern Norway (Fig. 17b). The lowest
mean basement velocities are determined for the Massif Central and
the Atlantic shelf area. In contrast, very high mean velocities (with a
+0.3 km/s Vp anomaly) in the basement of the Voring Plateau at
the Norwegian shelf can be associated with mafic intrusions in the
lower crust (compare with Fig. 17a).
3.3.2. Thickness of crustal layers
Structure of the crust, represented by thicknesses of individual
crustal layers, provides further insights into tectonic evolution of
western Europe and explanations for mean basement velocity varia-
tions (Figs. 17, 19). There is a remarkable correlation between the
thicknesses of the upper and lower crust and the mean basement
velocity. In Paleozoic Europe, a belt of the thickest upper crust
(20 km and more) extends from the Irish–British Caledonides
through the Armorican Massif to Iberia; the lower crust in these re-
gions is nearly absent (Fig. 17a). In contrast, the crust of the Norwe-
gian Caledonides has a significant thickness of the middle-crustal
layer, which explains its higher mean basement velocity than in the
Variscan Europe. However, the lower crust is apparently absent be-
neath the southern Norwegian mountains as well (Stratford et al.,
2009).
There is a striking difference between the stable platform crust of
the East European craton and the Caledonian–Variscan crust of
Europe, in particular in thicknesses of the middle and the lower crustal
layers (Figs. 17a, 19ac): while on the platform the middle crust is more
than 10 km in thickness and the lower crust is typically 8 to 12 km
thick (with significant regional variations), in the Caledonian–
Variscan structures the middle crust is much thinner (4–8 km) and
the lower crustal layer is largely absent or only few kilometers thick.
Lower crustal xenoliths from different Variscan locations in Germany
indicate that the deep crust is much more felsic (anorthosites, felsic
granulites, metagranitoids, high-grade metasediments) than expected
in continental settings and mafic granulites, which are commonly be-
lieved to be the major component of the lower crust, make only a few
kilometers thick layer above the Moho (Fig. 21). This observation
provides additional evidence for post-Variscan delamination of the
lower crust in the Paleozoic parts of Europe (Abramovitz et al., 1998;
Aichroth et al., 1992).
3.4. Structure of the Variscan lithospheric mantle
3.4.1. Delaminated lithosphere
Despite the complicated tectonic evolution of the Variscides, their
lithospheric structure is surprisingly simple and homogeneous, provid-
ing a strong argument that orogenic lithospheric structures did not
survive since Paleozoic (Artemieva et al., 2006). As discussed in
Sections 3.2–3.3, the sharp, subhorizontal Moho of the Variscides
crosses pre-existing terrane boundaries (Franke, 2006; Meissner et al.,
1986), indicating that these Paleozoic mountain belthas lost its former
high-velocity lower crustal roots. It is difficult to imagine a geodynamic
process which would remove only the lower crust without affecting the
lithospheric mantle formed during the Paleozoic orogeny.
In support of lithosphere delamination scenario, seismic data indi-
cate a rather uniform seismic velocity structure of the Variscan litho-
spheric mantle. Furthermore, dominating subhorizontal wide-angle
reflectors are imaged in the upper 90 km of the Variscan lithosphere
(Faber and Bamford, 1979; ILIHA DSS Group, 1993), although the res-
olution on these studies is relatively low (due to the >3 km intervals
between the seismic stations along the refraction profiles). These ob-
servations are interpreted as indicating that the Hercynian structures
have not been preserved in the European lithospheric mantle and
that the Paleozoic lithospheric mantle of the Variscides or its signifi-
cant part has been lost.
An example of on-going lithosphere delamination is provided by
seismic data from the Vrancea Zone in the eastern Carpathians. A
prominent deep, near-vertical lithospheric slab with high seismic ve-
locities down to a 300 km depth and with a strong mantle seismicity
underlies the eastern Carpathians at the sharpest bend of the moun-
tain belt (Koulakov et al., 2010). Whether the slab is oceanic or conti-
nental, and whether this high-velocity cylindrical anomaly represents
a subducting slab or a delaminated lithosphere is hotly debated
(Fillerup et al., 2010; Lorinczi and Houseman, 2009; Martin et al.,
2006; Tondi et al., 2009).
3.4.2. Is the Variscan lithospheric mantle partly preserved?
Seismic data on Pn anisotropy and SKS shear wave splitting indi-
cate that the Variscan lithospheric mantle is strongly anisotropic. Seis-
mic anisotropy of 6.5 to 15% for P-wave velocities has been interpreted
37I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
as evidence for paleosubduction zones associated with the closure of
the oceanic domains and the consequent Hercynian orogeny
(Plomerova et al., 1998). By the pattern of seismic anisotropy in the
lithospheric mantle, the Variscides can be subdivided into two do-
mains with NW- and SE-dipping anisotropic structures. The boundary
between the two domains has a general SW–NE orientation and ap-
proximately corresponds to the suture between the Saxothuringian
and Moldanubian terranes; the suture also corresponds to a ca.
40 km difference in lithospheric thickness and to a different depth of
the electrical conductor in the mantle (Plomerova et al., 1998; Praus
et al., 1990). NW- and SE-dipping anisotropic structures in the litho-
spheric mantle are interpreted as traces of two divergent systems of
Paleozoic subduction zones with olivine orientations inherited from
subducted ancient lithosphere (Babuška and Plomerová, 1992).
Studies of spinel lherzolite xenoliths from the Bohemian Massif
(which sample the Variscan lithosphere down to a depth of ca.
70 km) show that the Variscan lithospheric mantle (at least, in some
locations) has a layered structure: the upper layer has a horizontal
foliation whereas the lithospheric mantle below ca. 45 km depth has
a vertical (or steeply dipping) layering (Christensen et al., 2001). A
horizontal a-axis of mantle olivine in the lower lithospheric layer (be-
tween depths of 45 and 70 km), with an approximately E–W strike,
parallel to the observed fast shear wave direction, could have been
inherited from the Variscan convergence.
Even more complicated patterns are observed in the upper mantle
structure of the Armorican massif (north-western France) which has
two distinct patterns of S-wave seismic anisotropy (Judenherc et al.,
2002). The upper mantle at a 90–150 km depth has + 3% P-wave ve-
locity anomaly in the southern domain and −3% P-wave velocity
anomaly in the northern domain. Furthermore, while the southern
domain has a NW–SE orientation of Pn and SKS fast directions,
which are parallel to the strike of the South Armorican shear zone
and interpreted as orogen-related, SKS fast directions in the northern
domain are not parallel to the strike of major Hercynian shear zones.
This pattern of seismic anisotropy is interpreted to be formed during
a pre-Hercynian subduction, when two parts of the Armorican massif
were merged. Thus, in contrast to expectations, some parts of the
Variscan and even pre-Variscan lithospheric mantle apparently are
still preserved in the western Europe, at least within some of the
Variscan massifs.
4. Seismic reflections in the lithosphere of Western Europe
Structural and compositional heterogeneities in the lithosphere
produced by the processes of crust–mantle interaction are best
detected by seismic reflection surveys (e.g. Meissner et al., 1991).
Especially strong reflections, caused by high impedance contrasts,
occur from the Moho, from mafic/ultramafic sill-like intrusions in
the sialic or intermediate lower crust (the seismic lamellae)
(Meissner, 1986; Warner et al., 1996), and from fluid–solid interfaces
(Stratford and Stern, 2004). Fluid-saturated shear zones in the lower
crust that has experienced high-pressure metamorphism (eclogitiza-
tion) may also have high reflection coefficients (Fountain et al.,
1994). We next discuss seismic reflection observations in the crust
and lithospheric mantle of western Europe and their geodynamic sig-
nificance and relation to crust–mantle interaction processes.
4.1. Reflections from the lower crust
4.1.1. Observations
A characteristic feature of the crust in most areas of west-central
Europe is the strongly reflective (laminated) lower crust, which has
been documented by seismic reflection surveys in many countries
(Figs. 19e, 22a). Dense reflections (lamellae) are widespread from
the Faroe Islands in the north to the Alps and the Iberian Peninsula
in the south, and from the continental boundary of the Mid-Atlantic
ridge in the west to the TESZ in the east. They were intensively studied
around Britain (by numerous BIRPS surveys), in Germany (DEKORP
project), France (ECORS project), Iberia (IAM and ESCI projects), and
Switzerland (see summaries by Meissner, 1986; Meissner and Kern,
2008; Meissner and Rabbel, 1999; Mooney and Meissner, 1992;
Ziegler and Cloetingh, 2004). The only part of Phanerozoic Europe
which is free from lower-crust reflectivity includes the terranes of
East Avalonia in the North Sea (MONA LISA WG et al., 1997; Tanner
and Meissner, 1996).
Worldwide, the most remarkable seismic lamellae are observed in
extensional continental areas with thin crust and low mean basement
velocities (Fig. 23), such as the Basin and Range Province in western
USA (Prussen, 1991) and at passive margins. Some reflectivity is ob-
served in young collisional orogens such as Central Tibet and the
Alps (Ross et al., 2004) and in some parts of the Precambrian shields
(e.g. Baltic and Canadian Shields) that have undergone Proterozoic
orogenies, extension, and basin formation (BABEL WG, 1993;
Kukkonen and Lahtinen, 2006; Prussen, 1991; Wu and Mereu, 1991),
and not all of lower crustal reflectivity is associated with extension.
In contrast to west-central Europe, no significant lower crustal re-
flectivity has been observed in the thick and mafic lower crust of the
East European craton, except for the Paleozoic Dnieper-Donbas rift
(Maystrenko et al., 2003). In rifts and basins, strong extension and
massif voluminous intrusions may create a “reflection-unfriendly”sit-
uation within the regional “reflection-friendly”lower crust, where a
wide-spread, but moderate extension initiated mafic sills. An example
Fig. 21. Model profiles of crustal composition in three regions in Germany constrained by Vp seismic velocities (red lines) and crustal xenoliths (based on Downes, 1993; Mengel et
al., 1991).
38 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
of a “reflection-unfriendly”situation caused by massive magmatic in-
trusions is an African caldera where no reflections are observed within
the volcanic complex, but many are observed outside it (Bauer et al.,
2003). In contrast, “reflection-friendly”patterns are observed in
some of the European rifts. For example, clear lamellae are observed
in the lower crust of the Viking Graben but no reflections are seen in
the crust of the surrounding regions (McBride et al., 2004); the
thinned parts of the North Sea crust are strongly reflective within oth-
erwise non-reflective crust (Thybo et al., 1998). It is interesting that
some thin lamellae, possibly produced by ultramafic/piclogitic intru-
sions, are seen at the bottom of the mafic high-velocity (~ 7.2 km/s)
lower crust in the North German Basin (Bayer et al., 1995, 1999) and
in the Dnieper-Donets rift (Maystrenko et al., 2003).
Seismic lamellae are nearly always restricted to the continental
lower crust. Only two examples of strong seismic reflectivity outside
the lower crust are known at present. One of them (a “sill-like struc-
ture”which is 3–5 km thick and 175 km long) is located in the middle
crust of the Variscan southern Spain (Simancas et al., 2003). The other
structure (with numerous high-velocity sills) is observed in the upper
crust of the south-western Baltic Shield down to a 15 km depth
(Juhojuntti et al., 2001). By analogy with other (more isolated)
upper crustal sills, this anomalous reflectivity structure, that crosses
the Caledonian Deformation Front and the Precambrian Trans-
Scandinavian Igneous Belt, is interpreted to result from a Paleozoic
extensional episode (Ferré et al., 2002).
4.1.2. Interpretations
The preferred subhorizontal layering of the intruding magmas in
the low viscosity lower crust is caused by the development of exten-
sional stresses (Mooney and Meissner, 1992;Petford et al., 2000).
Moderate and long-time extension favors emplacement of multiple
sills which create high impedance contrasts with the surrounding
(mostly sialic) host rocks and produce the “lamellae”. Clearly, as evi-
denced by crustal reflectivity, the widespread ascend and intrusion of
mantle magmas into the lower crust is one of the important processes
of an upward transport of mantle material into the crust. Melt sills
associated with high crustal reflectivity are also observed in oceanic
crusts on top of the Moho around rift axes and are reported for
some prominent ophiolites (Canales et al., 2009; Korenaga and
Kelemen, 1997; Singh et al., 2006; Zagorevski et al., 2006).
Observations of seismic reflectivity underline the key role of exten-
sion and the volume of intruded mafic magmas in producing crustal
Fig. 22. Examples of seismic lamellae in the Variscan crust. (a) Line drawing of reflections from the migrated profile of DEKORP-2S. Thick lines—high quality reflections; vertical
exaggeration 1.5 (after Behr and Heinrichs, 1987). (b) Sketch illustrating a possible interpretation of crustal reflectivity by high- and low-velocity lamellae in the lower crust
(after Mooney and Meissner, 1992). Physical origin of seismic lamellae can be due to lithologic (mafic intrusions) and metamorphic layering.
Fig. 23. (a) Seismic lamellae in the lower crust in various tectonic provinces where normal-incidence and wide-angle observations are available (based on compilation of Meissner
et al., 2006). Four boxes refer to different tectonic settings: “Variscides”include both the Variscan European crust and Cenozoic crust of the Basin and Range province (USA); both
tectonic provinces have undergone significant lithosphere extension; “Pt cratons”include seismic data from the Paleo-Mesoproterozoic structures of the Canadian and Baltic
Shields; “Central Tibet”(only some selected measurements) and “Alps”are based on seismic data for two Cenozoic collisional structures. (b) Typical temperatures in the lithosphere
of different continental tectonic structures (based on Artemieva and Mooney, 2001). Colors match the corresponding structures in plot (a). Cold lithospheric temperatures in the
Tibet and the Alps are associated with subducting lithospheric slabs. Gray shading approximately marks the depths where seismic reflectivity is observed. As the plot illustrates,
seismic reflectivity is commonly restricted to a depth with temperatures between 300 and 500 °C.
39I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
reflectivity. In case large-volume massive intrusions are emplaced
during continental extension, they will form a totally new lower
crust with no impedance contrast between the former host rock and
the intruded magmas. Only when the volume of mafic–ultramafic in-
trusions (and the lithosphere extension) is moderate, can an alterna-
tion between the (low-velocity) host rock and the intruded layers
develop (Fig. 22b).
Strong impedance contrasts and dominating horizontal layering of
seismic lamellae have a striking similarity with reflectivity patterns
produced by intruded mafic sills. Commonly, the thickness of an indi-
vidual sill-layer in the lower crust is estimated to be 10 m to 100 m
(Handy and Brun, 2004; Holliger and Levander, 1994; Meissner and
Kern, 2008; Mooney and Meissner, 1992). Seismic studies in Canada
allowed to recognize different scaling laws for sills, laccoliths and plu-
tons, with sizes ranging from several cm to 600 m (Cruden and
McCaffrey, 2001). These values are of the same order of magnitude
as the estimates of sill size in the European crust based on seismic re-
flectivity (lamellae). A comparison of seismic data to synthetic data
generated from a synthetic stochastic velocity model allows for esti-
mates of geometry and seismic velocity contrast within high- and
low-velocity lamellae in the lower crust. In particular, the Variscan
crust can be modeled by a 400 ±300 m thick layers with a random ve-
locity contrast of ±0.05 to ±0.50 km/s with respect to mean crustal
velocity (e.g. Wenzel et al., 1987). A similar study based on the PASS-
CAL Basin and Range experiment (Holbrook, 1990) explains seismic
data by a 300 m thick and ca. 1 km long high-velocity anomalies
(Vp~6.3 km/s) as compared to the background Vp velocity of
6.0 km/s (Poppeliers, 2007). Greater thickness (400–600 m) of mafic
sills in the lower crust has been reported for the Dnieper-Donets rift
and the Baikal rift zone (Lyngsie et al., 2007; Thybo and Nielsen, 2009).
The depth at which lamellae develops is controlled by crustal rhe-
ology and, nearly always, a sialic/intermediate lower crust is the
weakest part of the lithosphere (Burov and Watts, 2006; Handy and
Brun, 2004; Meissner and Strehlau, 1982). The intruding magmas,
however, may change the crustal rheology. Cooling magmas, which
form subhorizontal sills connected by near-vertical (seismically invis-
ible due to their orientation) feeding dykes, may form a three-
dimensional solid network (Meissner and Kern, 2008). In contrast to
predictions of classical rheological models, this three-dimensional
“corset”makes the lower crust rheologically strong, at least in areas
of pronounced lamellae. This conclusion is supported by observations
of undisturbed, continuous structure of seismic crustal lamellae across
most of western/central Europe, below the Alpine and Pyrenees com-
pressional orogens, and below several Mesozoic extensional basins
around the British isles (Dyment et al., 1990). The presence of strike-
slip earthquakes down to a 30 km depth north of the Alps (including
occurrences in the laminated lower crust) further indicates that
today the lower crust is rheologically strong (Meissner and Kern,
2008). A comparison of the depth range of crustal reflectivity with
typical continental geotherms indicates that the reflectivity zone is
limited by temperature of ca. 500 °C (Fig. 23b). This temperature
approximately corresponds to the upper limit estimate for brittle–
ductile transition temperatures in quartz and feldspar-rich rocks, typ-
ical of the continental crust.
It is worth mentioning that deep crustal reflectivity may have a
multi-genetic origin. In particular, it can be caused not only by mag-
matic intrusions, but also by the presence of eclogite-facies rocks in
the lower crust. In particular, high reflection coefficients (0.04–0.14)
calculated for strongly foliated eclogite-facies rocks in a shear zone
within the Norwegian Caledonides indicate that deep crustal seismic
reflectors can be explained by the presence of eclogite-facies rocks in
those portions of the crust that were subject to high-pressure condi-
tions but did not experience high temperatures (Fountain et al.,
1994). In addition to lithologic/igneous/metamorphic layering, crustal
reflectivity can also be caused by the presence of shear zones (e.g.
Austrheim et al., 1997; Mooney and Meissner, 1992).
4.2. Reflections in the upper mantle
Compared to the lower crust, the upper mantle usually lacks a co-
herent reflectivity. Although occasionally upper mantle reflections are
observed (Yang, 2003), not all of them can be linked to the tectonic
history of a special location (Asencio et al., 2003). Impedance contrasts
are mostly positive (Warner et al., 1996) although some clear negative
contrasts have been observed, for example, in a volcanic back-arc zone
in New Zealand (Stratford and Stern, 2004).
Where isolated, often dipping, mantle reflectors are observed in
the continental lithospheric mantle, they are related to past collision-
al and subduction processes. A spectacular reflectivity example is
imaged by the URSEIS profile in the southern Urals (Knapp et al.,
1996). While the uppermost mantle shows several dipping reflectors
(apparently remnants of subduction zones), several strong, nearly
sub-horizontal reflectors are observed at depths between 75 km and
175 km. They can be associated with the continent–continent colli-
sion at the eastern edge of the East European craton but, because of
the complex tectonic history of the Uralides orogen, any tectonic in-
terpretation of mantle reflectivity should be made with caution.
Several locations in Europe with prominent sub-Moho reflections
should be mentioned. One is the spectacular upper mantle reflectivity
in northern Scotland, where two strong, sub-horizontal reflectors have
been imaged by BIRPS near-vertical reflection surveys (Warner et al.,
1996)andlaterconfirmed by wide-angle and receiver function studies
(Asencio et al., 2003). The reflectors have a positive polarity and appar-
ently represent a pre-Caledonian subduction zone. Similarly, several
dipping upper mantle reflections observed in the Baltic Shield (in the
North Sea and the Baltic Sea) at depths of 30–90 km (Fig. 10) are inter-
preted as remnants of slabs subducted in the Proterozoic and Paleozoic
(Abramovitz et al., 1997; Babel WG, 1990, 1993; Mona Lisa WG, 1997).
In contrast, no prominent mantle reflectors have been found below
the Variscan terranes of western Europe (Figs. 19e, 22a), where reflec-
tivity always terminates below the laminated lower crust, at the
Moho, and even tests with reflection recording times up the 30 s did
not show mantle reflectivity (Meissner, unpublished experiment). De-
lamination of the lowermost crust/lithospheric mantle in the post-
Variscan time together with high lithospheric temperatures and an
intensive magmatism homogenized much of the upper mantle and
destroyed the evidence of the previous tectonic processes. Possible
location of former collisional zones can be constrained by several dip-
ping reflectors that are present only in the rigid upper crust (Meissner
and Bortfeld, 1990).
5. How was the Variscan crust created?
Several types of significantly different material exchange processes
between the continental crust and the upper mantle have taken place
in various tectonic units and at various time in Europe. Here we dis-
cuss some questions related to these interaction processes, concen-
trating on one of the largest and most problematic tectonic unit in
Phanerozoic Europe, the Variscides.
5.1. Major characteristics of the present-day Variscan lithosphere
The characteristic and unusual features of the present-day litho-
sphere structure in Variscan Europe, which are described in detail
earlier, can be summarized in the following:
(1) a thin (only 30 km) continental crust;
(2) a flat and sharp Moho, which continues across major tectonic
boundaries,
(3) the absence of a high-velocity lower crust,
(4) wide-spread seismic reflectivity (lamellae) in the lower crust
with a sharp reflection Moho,
(5) general absence of upper mantle seismic reflectivity.
40 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
Apparently, all these patterns are typical for the “extended”crust
and are inter-related. For example, seismic refraction and reflection
experiments demonstrate similar crustal features in the extended
crust of the Basin and Range Province, western USA, where strong
orogenic processes preceded the extension, as it was in the Paleozoic
Variscides (Allmendinger et al., 1983, 1987; Gashawbeza et al., 2008;
McCarthy and Thompson, 1988). Regional extension in the Basin and
Range province is commonly explained by crustal “collapse”; the
lower crustal flow further smoothed out crustal thickness variations
caused by differential crustal extension and led to the creation of
the anomalous crustal structure (McKenzie et al., 2000). Similar con-
cepts supported by modeling (McKenzie, 1984) are favored for the
Variscan area.
A characteristic feature of the extensional regime of the Basin and
Range Province is the presence of metamorphic core complexes. The
latter are surface exposures of middle and lower continental crust
exhumed in association with largely amagmatic extension (Block
and Royden, 1990). Typically metamorphic core complexes consist
of an exposed high-grade metamorphic basement terrane and an
unmetamorphosed cover, between which are low-angle detachment
faults (mylonite shear zones) (Fig. 24a). Several metamorphic core
complexes are recognized in Europe, e.g. in the Western and Central
Carpathians and in the French Massif Central (Ciulavu et al., 2008;
Echtler and Malavieille, 1990; Franke et al., 2011; Janák et al., 2001).
The origin of the metamorphic core complex in the Massif Central is
related to the collapse of the Variscan orogen at ca. 300 Ma when
the compressional regime was followed by crustal extension, uplift,
and denudation (Fig. 24b–d).
Clearly, today the Variscan crust of Europe appears to be more “ex-
tensional”than orogenic (Menard and Molnar, 1988). The Variscan
orogenic roots were lost by oceanic and continental subduction, pos-
sibly supported by lithosphere delamination. In collisional orogens
formed by closure of an ocean basin (such as the Paleozoic Variscides
in Europe and the Tethyan Fold belt-Himalayas in Asia Minor-Central
Asia) continental or arc subduction can develop in case of strong lith-
osphere rheology and high convergence rate. In contrast, if the litho-
sphere is weak and the convergence rate is low, the development of
the Rayleigh–Taylor instability is a more likely scenario. In the former
case, convergence is accommodated by internal deformation, while in
the latter case by pure shear thickening. Extensional collapse of over-
thickened crust and the uppermost lithosphere (Menard and Molnar,
1988) and gravitational collapse caused by formation of eclogites (Le
Pichon et al., 1997; Mengel and Kern, 1992) have been proposed as
mechanisms responsible for the formation of thin crust in the Europe-
an Variscides. Post-orogenic wrench tectonic and extension, probably
initiated by the westward movement of Gondwana in the Permian,
has transformed the Variscan orogen into a typical extensional struc-
ture (Matte, 1991; Menard and Molnar, 1988).
The post-orogenic extension in the Variscides is commonly
explained by a stretching factor βof 1.5 to 1.7, assuming 45–50 km
Fig. 24. Metamorphic core complexes (MCC), asymmetrical dome like structures, usually formed in regions of continental extension after gravitational collapse of thick crust.
(a) Sketch (not to scale) illustrating the basic structure of a MCC, which consists of a metamorphic basement and an unmetamorphosed cover, separated by a decollement zone
with mylonitic fabric. (b) Location of the Montagne Noire MCC in the southern Massif Central, France. (c) Sketch (not to scale) illustrating various deformation styles associated
with the Late Carboniferous extensional tectonics in the Massif Central. (d) Evolutionary model for the various stages of Variscan deformation of the Montagne Noire MCC. Two
upper plates refer to orogenic shortening, while three lower plates refer to late orogenic extensional tectonics and uplift of the thickened core zone. The latter is advanced by
migmatization and anatectic granite emplacement. Development of Permian sedimentary basins (bottom plate) is controlled by high-angle normal faulting ((c–d) modified
after Echtler and Malavieille, 1990).
41I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
for the average thickness of the pre-extensional crust of the Variscan
orogen and 30 km for its present-day thickness. Similar values of the
β-factor are commonly estimated for extensional areas worldwide,
including the Basin and Range Province (Gans, 1987). This simple
stretching model allows for an additional 7–8 km of crustal thickening
by mafic magmas. This value is certainly a large overestimation since
the total thickness of mafic intrusions in the present-day Variscan
lower crust does not exceed 1–2 km. Larger volumes of intrusions
would have increased average seismic velocities, in contrast to seismic
observations in the Variscides (Fig. 17b). However, possibly an equiv-
alent of 7–8 km thick layer of mafic magmas is introduced not into the
lower crust but accumulated below the Moho (underplated).
5.2. Why is the mafic lower crust missing?
Two scenarios can explain the absence of the crustal roots beneath
the Paleozoic orogens on Europe and the missing mafic lower crust in
extensional areas. They describe a possible sequence of crust–mantle
material exchange mechanisms during or after main orogenic events
and are directly applicable to the European Variscides.
According to a scenario originally proposed for the main Appala-
chians (Nelson, 1992) and for the Paleozoic orogens of Europe
(Franke, 1992) and later examined numerically for the Variscan and
the Himalayan orogenies (Schott and Schmeling, 1998), the crustal
thinning occurs at the final stages of compression by subduction(s)
and delamination of thickened, eclogitized lower crust and uppermost
lithospheric mantle (Fig. 25). The presence of eclogites in lower crust-
al xenoliths from the Variscan Europe (Downes, 1993) suggests that
this scenario is plausible. Post-orogenic extension initiates the rise of
mantle magmas, produced by decompression melting of the upwell-
ing asthenosphere. These rising asthenospheric (mafic–ultramafic)
magmas replace delaminated lithosphere and form a new shallow
upper mantle with a sharp and flat Moho, identical in both reflection
and refraction seismics (Meissner and Bortfeld, 1990). Heat from the
rising magma initiates melting of the sialic crust leading to the forma-
tion of granitic bodies in the crust (Fig. 7), widespread in Variscan
Europe (Lliboutry, 1999; Oncken, 1994). Mafic–ultramafic magma
that enters the crust forms sills in its lower, rheologically favorable,
horizons. Thus formed lamellae and the new widespread flat and shal-
low Moho are present today below most (extensional) areas of west-
central Europe, including rifts and basins. High mantle temperatures
during the extension phase have dissolved sharp structural and com-
positional boundaries in the lithospheric mantle so that no mantle re-
flectors exist at present. An indirect support for this scenario comes
from the isotope data on the lower crustal xenoliths which indicate
that the age of the lower crust beneath the Variscan terranes is often
younger that the age of near-surface rocks. These results are inter-
preted as an evidence that magmatic intrusions and underplating
during periods of subduction, rifting and orogenic collapse in the
Variscides added a significant portion of the mafic lower crust to the
base of the pre-existing crust (Downes, 1993).
An alternative scenario for the missing mafic lower crust beneath
the Paleozoic orogens of western Europe has been proposed by
Berthelsen (1998). In this model, the former mafic lower crust un-
dergoes partial eclogitisation. A new seismic Moho is formed on top
of the eclogite-facies layer which (being made of a mixture of basaltic
and eclogitic rocks) can be seismically undistinguishable from the un-
derlying mantle peridotite, whereas the “petrologic Moho”remains at
the base of the eclogite-facies layer. Although this scenario may ex-
plain seismic observations of a thin, two-layered crust beneath the
Variscides, it is difficult to imagine that large-scale lower crustal eclo-
gitization was able to modify the crust of most of western Europe
during a few hundred millions of years; tectonic analogues operating
on such a scale are not well documented but similarities have been
proposed for the Variscides, Tibet and the Andes (Le Pichon et al.,
1997). Someanisotropy patterns observed within the Variscan terranes
also cannot be easily explained by this mechanism (Section 4.2), and
this scenario is in conflict with younger isotope ages of the lower crustal
rocks than of the overlying rocks (Downes, 1993).
5.3. Unanswered questions
Both scenarios described in Section 5.2 provide mechanisms how
the crust which formed during a large-scale Paleozoic orogeny may
have lost or chemically changed its lower parts to become thin and
uniform as it is seismically observed at present. It is, however, still a
mystery what was the sequence of tectonic events that formed such
a uniform “extensional”crust over the whole of western and central
Europe. Apparently, only a combination of massive tectonic exten-
sional stresses with an extensive, continent-scale, Variscan or/and
post-Variscan heating event could possibly have created a uniform
crust over such a huge area, and one has to acknowledge a widespread
Fig. 25. Sketch illustrating the transition from compressional to extensional tectonic regime. (I) Tectonic compression leads to mountain building and subduction of lithospheric
plate, enhanced by eclogitization of the lower crustal material. (II) As a result of delamination of thickened, eclogitized lower crust and uppermost lithospheric mantle, crustal thin-
ning occurs at the final stages of compression. (III) Decompression melting of the upwelling asthenosphere during post-orogenic extension initiates the rise of mantle magmas,
which replace delaminated lithosphere and form a new shallow upper mantle with a sharp and flat Moho. (IV) Crustal heating by rising magmas initiates its melting with the for-
mation of granitic plutons (not shown), while mafic–ultramafic magmas that enter the lower crust form reflecting sills (seismic lamellae).
42 I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
large-scale material exchange processes between the crust and mantle
of Europe.
Many more problems still remain to be explained for the Variscan
lithosphere of Europe. Some of them are listed below.
(1) Wide-spread eclogitization of the Variscan lower crust is prob-
lematic. The timing of the eclogitization raises further questions.
The situation in the North German Basin is especially challenging:
majorsubsidenceinthebasinoccurredinthePermo-Triassic
when more than 10 km of sediments were accumulated there
(Fig. 20a). Given that the present Moho is at ca. 30 km depth, it
means that the thickness of the crystalline crust was reduced to
less than 20 km, that is by more than 30%. Furthermore, a ca.
10-km thick high-velocity layer with P-wave velocity of
6.9–7.5 km/s has been detected in the lower crust north of the
Elbe river in an area of ca. 150 ×180 km by a network of deep
seismic refraction profiles in Northern Germany (Rabbel et al.,
1995). It is difficult to imagine that a significant portion of the
lower crust could have been eclogitized during the last few
hundred million years even though the rate of the eclogitization
processisnotwellconstrained.
We thus suspect that the formation of the shallow Moho is rela-
tively recent and it occurred when rising (probably piclogitic)
magmas achieved gravitational equilibrium. This hypothesis is
in accord with an explanation of the origin of the high-velocity
lower crustal body beneath the North German Basin by infiltra-
tion of mafic magmas during the extensional stage (Rabbel et
al., 1995), although greenshist facies to amphibolite facies meta-
morphism of the lower crust was also proposed as a possible
mechanism (Brink, 2005). Anderson (2007) argues that litho-
sphere heating may cause a “Yo-Yo”tectonics giving rise to eclo-
gitic blobs and their mixing with peridotite to form “piclogite”.
Major element composition of these piclogites is intermediate
between dunites and basalts but with a larger basalt fraction
than peridotites, so that their P-velocity is ca. 8 km/s (e.g. Given
and Helmberger, 1981). These piclogites are likely to be a main
component of rising asthenospheric material that forms the
new Moho, the new upper mantle and also mafic intrusions
into the lower crust. Their possible presence, seismically unde-
tectable, however should have a strong effect on crustal/litho-
spheric buoyancy.
(2) The sequence of tectonic processes described above (Fig. 25),
from compression to extension, is not the only way to create
extensional continental crust. Heterogeneous seismic velocity
structure of Phanerozoic Europe (Fig. 17b), although to a
large extent due to significant lateral temperature variations,
suggests that some remnants of delaminated and subducted
lithosphere can still be present in the upper mantle of the Var-
iscides (Bijwaard and Spakman, 2000; Piromallo and Morelli,
2003). In particular, mantle Vp/Vs velocity models (that are
weakly sensitive to regional temperature variations in the
mantle but are more sensitive to compositional variations) in-
dicate significant compositional heterogeneity in the upper
mantle of the Variscides down to at least 200 km depth (e.g.
Artemieva, 2009; Artemieva et al., 2006), although sharp struc-
tural and compositional boundaries in the mantle (as evi-
denced by the absence of mantle seismic reflections) may
have been eliminated by very high (homogenizing) tempera-
tures during lithosphere extension.
(3) The sequence of the formation of granites and seismic lamellae
in the crust poses another problem. There are no definite signs
for the generation of granites inside the lamellae of the lower
crust, and one has to assume that the mafic sills were formed
after granitic plutons, at least in those areas where granitic
intrusions are present and seismic lamellae have a very regular
pattern.
The regularity of the lamellae raises another question. Surpris-
ingly, the lamellae do not show any preference (or systematic
variation) regarding their depth distribution and intensity
within the lower crust. Since the lowermost layer is the hottest
part of the crust and its viscosity should be the lowest, it should
facilitate intrusions of magmas leading to an increase of seis-
mic reflectivity. Surprisingly, the regular depth pattern of la-
mellae, which indicates a similar density and thickness of
intruding sills, suggests a constant lower crustal viscosity.
(4) A special problem for the Variscan orogeny is the question
whether the colliding Paleozoic terranes that formed most of
the western-central European crust (Figs. 13–15) had a
“shield-like”three-layer pre-collisional crust. One cannot ex-
clude the possibility that the terranes that came from the
northern rim of Gondwanaland were possibly already “exten-
sional”and the mafic lower crust was already absent there.
Such an assumption could provide an additional explanation
for the absence of a thick mafic lower crust in west-central
Europe and for the timing of post-Variscan tectonic processes
in Europe, reducing the role of crustal eclogitization in the for-
mation of the present-day crust.
6. Conclusions
Global interaction processes between crust and mantle, many of
which are manifestations of plate tectonic processes, have deter-
mined the development of the crustal structure of the Earth. A com-
parison of the ratio of crust-to-mantle volume on the terrestrial
planets suggests that plate tectonic processes, together with crustal
eclogitization, may be important in limiting the total thickness and
total volume of the crust. On Earth, crustal volume is effectively con-
trolled by subduction (oceanic or continental) and delamination
which recycle crustal (lithospheric) material into the mantle and re-
duce its volume. Working in the opposite direction, magmatic intru-
sions, underplating and exhumation create new crust and can
increase the volume of the existing crust. Consequently, crustal thick-
ness may be reduced during magmo-tectonic events, given that mag-
matism is nearly always associated with lithosphere extension.
Crustal structure as revealed by seismic data preserves the record
of past tectonic processes responsible for its formation and structural
and chemical modification. The role of material exchange processes in
shaping the structure of the continental crust is illustrated using the
European continent as an example, with focus on the structure of
the present-day crust of the Variscan terranes the origin of which is
still highly debated. This crust, created as a result of large-scale Cale-
donian and Variscan orogenies, did not preserve orogenic crustal
roots but instead has all present-day characteristics of the extended
continental crust. Understanding the origin of the anomalous “exten-
sional”crust in the European Variscides requires understanding:
(i) the origin of flat Moho, thin crust, and low lower crustal
velocities;
(ii) the origin of lower crustal reflectivity (lamellae) in the European
Variscides and absence of mantle reflectivity in most of Europe;
(iii) the role of basalt/gabbro–eclogite phase transition in limiting
crustal thickness worldwide, including depth–temperature
variations for this transition in different tectonic provinces of
Europe and conditions for a long-time preservation of thick
crustal roots beneath stable continental interior;
(iv) the links between crustal recycling, mafic magmatism, sills and
crustal underplating, and their effects on crustal rheology.
We conclude that there are several possible scenarios of the post-
Variscan evolution of the crust of western-central Europe which leave
the open questions that still await to be answered.
43I.M. Artemieva, R. Meissner / Tectonophysics 530-531 (2012) 18–49
Acknowledgments
Two anonymous reviews and comments of H. Thybo to the early
version of the manuscript are appreciated. The study is supported
by FNU-Denmark grant to IMA.
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