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Lead isotope evolution of the Central European upper mantle: Constraints from the Bohemian Massif

Authors:
  • Czech Academy of Sciences and Brno University of Technology

Abstract and Figures

The Pb isotope composition of the upper mantle beneath Central Europe is heterogeneous due to the subduction of regionally contrasting material during the Variscan and Alpine orogenies. Late Variscan to Cenozoic mantle-derived melts allow mapping this heterogeneity on a regional scale for the last ca. 340 My. Late Cretaceous and Cenozoic anorogenic magmatic rocks of the Bohemian Massif (lamprophyres, volcanic rocks of basanite/tephrite and trachyte/phonolite series) concentrate mostly in the Eger Rift. Cretaceous ultramafic lamprophyres yielded the most radiogenic Pb isotope signatures reflecting a maximum contribution from metasomatized lithospheric mantle, whereas Tertiary alkaline lamprophyres originated from mantle with less radiogenic 206Pb/204Pb ratios suggesting a more substantial modification of lithospheric source by interaction with asthenospheric-derived melts. Cenozoic volcanic rocks of the basanite/tephrite and trachyte/phonolite series define a linear mixing trend between these components, indicating dilution of the initial lithospheric mantle signature by upwelling asthenosphere during rifting. The Pb isotope compositon of Late Cretaceous and Cenozoic magmatic rocks of the Bohemian Massif follows the same Pb growth curve as Variscan orogenic lamprophyres and lamproites that formed during the collision between Laurussia, Gondwana and associated terranes. This implies that the crustal Pb signature in the post-Variscan mantle is repeatedly sampled by younger anorogenic melts. Most Cenozoic mantle-derived rocks of Central Europe show similar Pb isotope ranges as the Bohemian Massif.
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Research Paper
Lead isotope evolution of the Central European upper mantle: Constraints
from the Bohemian Massif
Simona Krmí
ckov
a
a
,
b
, Luk
a
s Krmí
cek
a
,
b
,
c
,
*
, Rolf L. Romer
d
, Jaromír Ulrych
b
a
Department of Geological Sciences, Faculty of Science, Masaryk University, Kotl
a
rsk
a 2, CZ-611 37, Brno, Czech Republic
b
Institute of Geology of the Czech Academy of Sciences, Rozvojov
a 269, CZ-165 02, Prague 6, Czech Republic
c
Brno University of Technology, Faculty of Civil Engineering, AdMaS Centre, Veve
rí 95, CZ-602 00, Brno, Czech Republic
d
Deutsches GeoForschungsZentrum GFZ, Telegrafenberg, D-144 73, Potsdam, Germany
ARTICLE INFO
Handling Editor: Christopher J Spencer
Keywords:
Lead isotopes
Lamprophyres
Volcanic rocks
Mantle components
Bohemian massif
Variscan orogeny
ABSTRACT
The Pb isotope composition of the upper mantle beneath Central Europe is heterogeneous due to the subduction of
regionally contrasting material during the Variscan and Alpine orogenies. Late Variscan to Cenozoic mantle-
derived melts allow mapping this heterogeneity on a regional scale for the last ca. 340 Myr. Late Cretaceous
and Cenozoic anorogenic magmatic rocks of the Bohemian Massif (lamprophyres, volcanic rocks of basanite/
tephrite and trachyte/phonolite series) concentrate mostly in the Eger Rift. Cretaceous ultramac lamprophyres
yielded the most radiogenic Pb isotope signatures reecting a maximum contribution from metasomatised lith-
ospheric mantle, whereas Tertiary alkaline lamprophyres originated from mantle with less radiogenic
206
Pb/
204
Pb
ratios suggesting a more substantial modication of lithospheric source by interaction with asthenospheric-
derived melts. Cenozoic volcanic rocks of the basanite/tephrite and trachyte/phonolite series dene a linear
mixing trend between these components, indicating dilution of the initial lithospheric mantle signature by up-
welling asthenosphere during rifting. The Pb isotope composition of Late Cretaceous and Cenozoic magmatic
rocks of the Bohemian Massif follows the same Pb growth curve as Variscan orogenic lamprophyres and lamp-
roites that formed during the collision between Laurussia, Gondwana, and associated terranes. This implies that
the crustal Pb signature in the post-Variscan mantle is repeatedly sampled by younger anorogenic melts. Most
Cenozoic mantle-derived rocks of Central Europe show similar Pb isotope ranges as the Bohemian Massif.
1. Introduction
Central Europe has been affected by two major extensional events of
Permo-Carboniferous and Tertiary age (e.g., Meier et al., 2016).
Permo-Carboniferous extension began in Central Europe because of
post-collisional changes in the kinematics between the Laurussia and the
Gondwana plates (Kroner et al., 2016) and resulted in the formation of
the Central European Extensional Province (CEEP) with numerous
volcano-sedimentary basins in Europe as well as the Oslo Rift (e.g.,
Neumann et al., 2004; Wilson et al., 2004). This extension was coeval
with the opening of the Palaeo-Tethys Ocean farther to the east and the
closure of the remaining Rheic Ocean farther to the west (Kroner and
Romer, 2013). Late Cretaceous to Tertiary lithospheric extension was
initiated as a response to the tensional reactivation of Variscan
Permo-Carboniferous fracture systems by the Alpine collision, eventually
resulting in the formation of a vast rift system in Western and Central
Europe. The European Cenozoic rift system includes the following indi-
vidual rifts: Valencia Trough in Spain, the Gulf of Lions, the Sa^
one,
Limagne and Bresse grabens in south-eastern France, the Rhine, Ruhr and
Leine grabens in Germany, and the Eger Rift in the Bohemian Massif
(Ziegler, 1992,Fig. 1).
Formation of continental rift systems associated with horizontal
movements of plates and subsequent lithosphere thinning is commonly
accompanied with magma generation by decompression melting of
lithospheric and asthenospheric mantle that is passively upwelling
beneath the thinned lithosphere (e.g., Berkesi et al., 2019). Melts
generated in rifts by decompression are separated from their residue and
ascend from the zone of melting in the upper mantle and are emplaced
within the overlying continental crust or are extruded as lava ows. The
volume of the produced melts depends on the amount of lithospheric
* Corresponding author. Institute of Geology of the Czech Academy of Sciences, Rozvojov
a 269, CZ-165 02, Prague 6, Czech Republic.
E-mail address: lukas.krmicek@gmail.com (L. Krmí
cek).
Peer-review under responsibility of China University of Geosciences (Beijing).
HOSTED BY Contents lists available at ScienceDirect
Geoscience Frontiers
journal homepage: www.elsevier.com/locate/gsf
https://doi.org/10.1016/j.gsf.2019.09.009
Received 27 April 2019; Received in revised form 2 August 2019; Accepted 25 September 2019
Available online 23 October 2019
1674-9871/©2019 China University of Geosciences (Beijing) and Peking University. Production and hosting by Elsevier B.V. This is an open access article under the
CC BY-NC-ND license (http://creativecommons.org/licenses/by-nc-nd/4.0/).
Geoscience Frontiers 11 (2020) 925942
extension and on the temperature of the asthenosphere. The primary
products of volcanic activity in rifts are mainly basaltic. Continental
intra-plate rifting produces predominantly olivine/nepheline-bearing
alkaline basalts (e.g., White and McKenzie, 1989; Wilson and Downes,
1992; Lustrino and Wilson, 2007; Haase and Renno, 2008).
For the Cenozoic rift system of the Central European Volcanic Prov-
ince (CEVP), there are two end-member type models. Some authors
favour asthenospheric melting in response to large-scale upwelling of
mantle plumes or small-scale plumelets (e.g., Wilson and Downes, 1991;
Hegner et al., 1995). Other studies favour derivation of the intraplate
volcanic rocks from the metasomatised lithospheric mantle, either by
preferential melting of phlogopite/amphibole-bearing vein assemblages
hosted in lherzolitic mantle or by very low degree melting of highly
metasomatised domains, resulting in the generation of strongly
SiO
2
-undersaturated melts (Hegner et al., 1995; Jung et al., 2005;
Pf
ander et al., 2018).
Late Variscan to Cenozoic mantle-derived rocks, such as lamp-
rophyres, lamproites, and alkali-basalts, largely sampled the same mantle
on a regional scale and at variable depth during above mentioned
extensional events (e.g., Mayer et al., 2014; Krmí
cek et al., 2016; Ulrych
et al., 2018). Although mantle-derived lamprophyres and associated
intrusive rocks are relatively rare within the volcanic complexes, these
small-volume melts are particularly important for studying processes and
geochemical heterogeneities of the upper mantle beneath Central Europe
as they preferentially sample the metasomatic component, which was
induced during the Variscan orogeny when continental collision brought
continental crust to mantle depth (e.g., Kroner and Romer, 2013; Bor-
ghini et al., 2018; Pf
ander et al., 2018).
The Sr, Nd, and Pb isotope composition of mantle-derived rocks may
identify contributions from different mantle sources, such as depleted
asthenospheric and enriched lithospheric upper mantle and lower
mantle, to the generated magmas (e.g., White and McKenzie, 1989; Zou
et al., 2000). Isotope constraints on the mantle sources of Mesozoic and
Cenozoic volcanic rocks of the Bohemian Massif largely rely on SrNd
isotope data (e.g., Alibert et al., 1983, 1987; Blusztajn and Hart, 1989;
Bendl et al., 1993; Vokurka, 1997; Ulrych et al., 2002, 2008, 2011, 2013,
2016, 2018; Haase and Renno, 2008; Cajz et al., 2009; Sk
ala et al., 2014,
2015). In contrast, Pb isotope data are rare (Blusztajn and Hart, 1989;
Fig. 1. (A) Distribution of major volcanic areas
within Central Europe (modied after Blusztajn and
Hegner, 2002); EL Elbe Line, OFZ Odra Fault
Zone. (B) Map of volcanic centres in the Eger Rift
with sampling areas (modied after Zachari
a
s et al.,
2008). Abbreviations: BRB Berzdorf Radomierzyce
Basin; CDG ChebDoma
zlice Graben; CSVC
Cesk
e
St
redoho
rí Volcanic Complex; DHVC Doupovsk
e
Hory Volcanic Complex; MB Most Basin; RPVC
Ralsk
a Pahorkatina Volcanic Complex; SB Sokolov
Basin; ZB Zittau Basin.
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
926
Haase and Renno, 2008; Ulrych et al., 2016). As typical mantle has low
Pb contents, the Pb isotope composition of the mantle is readily affected
during metasomatism. Because of the contrasting Pb contents in crust and
mantle rocks, small contributions of continental material are more
sensitively recorded in the Pb isotope composition than in the Sr and Nd
isotope composition (Cohen and ONions, 1982; Davies and Macdonald,
1987; Sun and McDonough, 1989; Conticelli et al., 2002).
Our paper focuses on Pb isotope data and whole-rock geochemistry of
intrusive and extrusive volcanic rocks of the Bohemian Massif that
sampled the upper mantle. We pay special attention whether Late
Palaeozoic to Quaternary Central European mantle-derived rocks
sampled different mantle sources on a local to regional scale and through
time.
2. Geological setting
The upper mantle beneath Central Europe is heterogeneous, which is
largely due to Variscan and Alpine subduction events during which
regionally contrasting subducted material of both continental and
oceanic crust metasomatically modied the upper mantle (e.g., Wit-
t-Eickschen and Kramm, 1997; Downes, 2001; Ackerman et al., 2009;
Kroner and Romer, 2013; Pf
ander et al., 2018). Older tectonic structures
reactivated during the late Variscan to Alpine events were associated
with asthenospheric upwelling and subsequent extension-related mag-
matism that sampled this modied mantle (Wilson and Downes, 1991).
2.1. Variscan to post-Variscan development
The Variscan orogen is the result of the DevonianCarboniferous
collision of Gondwana with Laurussia that started with the collision of
the Armorican Spur, which is part of segmented peri-Gondwana. The
collision of the Armorican Spur with Laurussia resulted in the closure of
the Rheic Ocean in the area of future Central Europe (Kroner and Romer,
2010, 2013). The subducted material consisted primarily of relative thin
lithosphere covered by Palaeozoic volcano-sedimentary rocks (Kroner
et al., 2007). In contrast, the unsubductable parts of the peri-Gondwana
(Cadomian) magmatic arc (i.e., thick crustal fragments of the Armorican
Spur) caused reorganisation within the plate boundary zone (Kroner and
Romer, 2013). The Bohemian Massif consists of high- and low-strain
domains that behaved differently during the Variscan orogeny. The
Tepl
a-Barrandian Unit and Lusatia (parts of Gondwana) and the Bruno-
vistulian Terrane (part of Laurussia) are low-strain domains that collided
during the closure of the Rheic ocean (e.g., Kalvoda et al., 2008; Kalvoda
and B
abek, 2010). The Saxo-Thuringian and Moldanubian Zones are
high-strain domains that represent former Gondwana shelf and that were
subducted and exhumed during the Variscan orogeny (e.g., Kroner and
Romer, 2013; Krmí
cek et al., 2016;
Z
ak and Sl
ama, 2018).
Plate tectonic processes that may have changed the composition of
the upper mantle of the Bohemian Massif include (i) Cadomian subduc-
tion followed by Late Neoproterozoic back-arc spreading and early
Palaeozoic rifting in northern peri-Gondwana, (ii) intra-oceanic sub-
duction and formation of oceanic arcs that were later accreted to form
part of Variscan Europe, (iii) possible subduction beneath the northwards
migrating terranes, (iv) subduction and collision at the southern margin
of Laurussia, and (v) Carboniferous extension of Central Europe (Wilson
et al., 2004; McCann et al., 2006; Pin and Waldhausrov
a, 2007; Tim-
merman, 2008; Abdelfadil et al., 2013; Dostal et al., 2019a,b). Subduc-
tion of both oceanic and continental crust peaked at 340 Ma and led to
the establishment of isotopically contrasting domains in the meta-
somatised lithospheric mantle (Krmí
cek et al., 2016). During the late
stage of the Variscan orogeny, these metasomatised mantle domains
underwent partial melting resulting at ca. 340300 Ma in the intrusion of
potassic to ultrapotassic dykes (lamprophyres and lamproites) along
deep-fault zones related to initial crustal extension eventually leading to
the formation of the CEEP (Awdankiewicz, 2007, 2009; Krmí
cek et al.,
2011, 2014, 2016; Abdelfadil et al., 2013; Hrouda et al., 2016).
Post-Variscan magmatic activity throughout Variscan Europe was
associated with the extension of thickened Variscan crust and accom-
panied by progressively increasing contributions from the astheno-
spheric mantle (Lorenz and Nicholls, 1984; Timmerman et al., 2009).
Permian extension led to the formation of rift basins, such as the North
Germany Basin or the Oslo Rift (Benek et al., 1996; McCann et al., 2006).
2.2. Mesozoic to Cenozoic development
Within the framework of the AfricanEurasian plate collision, the
extensive European Cenozoic Rift System (ECRIS) formed in Western and
Central Europe stretching from Spain and France through Germany to the
Czech Republic and Poland (Prodehl et al., 2006). The ECRIS recorded
intermittent sub-volcanic/volcanic activities that started in the Late
Cretaceous and have lasted to the present (Lustrino and Wilson, 2007).
Episodic volcanism occurred mainly in the Oligocene to Miocene with
waning phases of anorogenic volcanic activity reaching to the
Plio-Pleistocene (Nowell et al., 2006). Magmatic activity in the ECRIS is
concentrated in intrusive complexes and volcanic elds within the gra-
bens and their shoulders (Fig. 1A). Rifts formed at reactivated Variscan
suture zones separating large different lithospheric segments, indicating
structural control on the location of Cenozoic volcanic activity (D
ezes
et al., 2004).
The Bohemian Massif is transected by the nearly 300 km long
ENEWSW trending Eger Rift, by the transverse NWSE striking Elbe/
LabeOdra Fault System in the north (Fig. 1A), and by the NWSE
trending ChebDoma
zlice Graben in the west (
Spa
cek et al., 2011). The
Eger Rift represents the easternmost part of the Cenozoic rift system of
the Central European Volcanic Province (e.g., Ziegler, 1994; Lustrino and
Wilson, 2007). The magmatic rocks of the Eger Rift are predominantly
SiO
2
-undersaturated alkaline rocks of intra-plate origin (Ulrych et al.,
2002,2011;Lustrino and Wilson, 2007; Dostal et al., 2017). The volcanic
activity is a result of reactivation of Variscan structures in the Bohemian
Massif during the Alpine orogeny (Babu
ska and Plomerov
a, 1992, 2001,
2010). Whether the model of mantle plumes sensu Wilson and Paterson
(2001), with Alpine exure and lithospheric extension followed by
adiabatic decompression, decompression melting, and injection of
mantle-derived magmas into the crust, played a major role is debated
(Wilson and Downes, 1991; Wedepohl et al., 1994; Lustrino and Wilson,
2007; Ulrych et al., 2011) and is reected in terms like Common Mantle
Reservoir (CMR Lustrino and Wilson, 2007) or European Astheno-
spheric Reservoir (EAR Cebri
a and Wilson, 1995) used to characterise
magmas from a possible sub-lithospheric source region.
Ulrych and Pivec (1997) and Ulrych et al. (2011) dened three phases
of Cenozoic volcanic activity based on KAr dating and palaeostress
mapping: (i) a pre-rift period (~8049 Ma), (ii) a syn-rift period (4216
Ma) and (iii) a late-rift period (160.26 Ma). Ulrych and Pivec (1997) also
dened two coeval alkaline series: (i) a volumetrically dominant neph-
elinitebasanitetephritephonolite series of strongly to mildly alkaline
rocks and (ii) a subordinate and only locally occurring weakly alkaline
alkali basalttrachybasalttrachyandesitetrachyterhyolite series. The
Eger Rift comprises several volcanic centres (Fig. 1B), namely the
Cesk
e
St
redoho
rí Volcanic Complex (CSVC), the Doupovsk
e Hory Volcanic
Complex (DHVC; access to volcanic edice exposures is restricted as they
are predominately located in the Doupovsk
e Hory military training zone),
the Ralsk
a Pahorkatina Volcanic Complex (RPVC) and a great number of
isolated volcanoes in the western part of the Bohemian Cretaceous Basin
and in the ChebDoma
zlice Graben (CDG). The youngest volcanoes occur
in the Cheb/Eger Basin, a sub-basin of the Eger Rift.
The rst manifestation of volcanic activity in the Eger Rift is Late
Cretaceous ultramac lamprophyres and related melilitic rocks associ-
ated with subsurface intrusion of melilitolite composition that occur near
the intersection of the marginal fault of the Eger Rift and the Lusatian
Fault within the future RPVC in the northern part of the Bohemian
Cretaceous Basin (Ulrych et al., 2014). Most of the Cenozoic volcanic
rocks, predominantly of basanitic composition, are concentrated in the
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
927
CSVC and DHVC, whereas hypabyssal intrusions spatially associated with
alkaline lamprophyre dykes of Cenozoic age are restricted to the CSVC
and RPVC (Sk
ala et al., 2014). Sr-Nd-(Pb)-isotope compositions of rocks
from the western Eger Rift, i.e., the DHVC and CDG, were recently
published by Ulrych et al. (2016) and Haase et al. (2017). For locations of
investigated samples and characteristics of sampled areas see Fig. 1B and
Supplement A.
3. Methods
Representative samples of the studied volcanic and subvolcanic rocks
are described petrographically using conventional optical microscopy
and characterised geochemically by whole-rock and mineral composition
using ICP-ES, ICP-MS, and electron microprobe, respectively.
Whole-rock chemical analyses from fresh samples lacking signs of
alteration or wall-rock assimilation were carried out at Bureau Veritas
(former ACME) Analytical Laboratories Ltd. (Vancouver, Canada) using
inductively coupled plasma emission spectrometry (ICP-ES; major oxides,
Ba, Ni, Cu, Pb,Zn) and inductively coupled plasmamass spectrometry (ICP-
MS; Co, Cs,Hf, Nb, Rb, Sr, Ta, Th,U, V, Zr, Y and REE). Losson ignition (LOI)
was determined by weight difference after ignition at 1000 C. The Pb, U,
and Th concentrations are used to recalculate the initial Pb isotope
composition. The errors for the Pb, U and Th concentrations correspond to
0.1 ppm, 0.1 ppm, and 0.2 ppm, respectively. For further analytical
details and detection limits see www.acmelab.com.
The composition of characteristic dark minerals (amphibole, clinopyr-
oxene, and mica) was analysed using a CAMECA SX 100 electron micro-
probe (Institute of Geology of the CAS, Prague) operated in wavelength-
dispersive mode. Measurements were performed using a 15 keV accelera-
tion voltage, 10 nA beamcurrent and 2
μ
m beam diameter. Both naturaland
synthetic minerals were used as reference standards. Concentrations of
following elements were measured (standard, spectrometer crystals and
detectionlimit for analysedelements aregiven in parentheses): Si (diopside,
LTAP, 222ppm), Ti (rutile,LPET, 357 ppm), Al (jadeite, LTAP,272 ppm), Cr
(MnCr spinel, LIF, 910 ppm), Fe (haematite, LIF, 1047 ppm), Mn (rhodo-
nite, LIF, 965 ppm), Ni (Ni
2
Si, LTAP, 1404ppm), Mg (periclase, LTAP, 422
ppm), Ca (diopside, LPET, 341 ppm), Na (jadeite, LTAP, 262 ppm), K
(leucite, LPET, 300 ppm), F (uorite, PC0, 1575 ppm), Cl (tugtupite, LPET,
320 ppm), Rb (RbCl, LTAP, 217 ppm) and Ba (barite, LPET, 503 ppm).
Countingtimes on peaks were10 s for Mg, Al, K, Ca,Cl, Ti; 20 s for Na, Si,Rb,
Ba, Ni andMn, Ni; and 30 s for Cr, Ca andAl. The X-phi correction procedure
(Merlet, 1992) was used for spectra processing.
The Pb isotope compositions were determined at Deutsches Geo-
ForschungsZentrum (GFZ), Potsdam, Germany. Powders from represen-
tative unaltered samples of volcanic and subvolcanic rocks were
dissolved in concentrated HF for four days on a hot plate at 160 C.
Samples were dried, re-dissolved in 2 N HNO
3
and dried slowly at 80 C
overnight to convert uorides to nitrates. Finally, the samples were taken
up in 6 N HCl to convert nitrates to chlorides. Lead was separated in
columns using ion exchange resin Bio Rad AG-1-X8. Procedures for the
separation and purication of Pb are described in detail by Romer et al.
(2005). The ion exchange procedure was repeated to purify Pb elutes.
Lead was loaded together with H
3
PO
4
and silica-gel emitter on single
Re-laments. The Pb isotope ratios were measured using a Thermo-
Finnigan Scientic TRITON TIMS multi-collector mass-spectrometer
operated in static multi-collection mode. The obtained Pb isotope ratios
were corrected for instrumental fractionation of 0.1%/a.m.u. as deter-
mined from repeated measurement of Pb reference material NBS 981.
Total procedural blank is between 15 and 30 pg Pb, thus, negligible.
Accuracy of the determined Pb data is better than 0.1% at the 2
σ
level.
4. Results
4.1. Field and petrographic descriptions
The studied volcanic rocks were mainly taken from exposures and
quarries (Fig. 2A and B), whereas lamprophyres and related dykes were
mostly collected as loose angular blocks in the eld (Fig. 2C). Melilitolite
sample (OC-12) originates from a 271 m deep Holi
cský vrch borehole
near Ose
cn
a(Table 1).
Late Cretaceous ultramac lamprophyres, such as polzenite and
aln
oite dykes and the melilitolite of the Ose
cn
a intrusion, were sampled
in the RPVC. Polzenites and aln
oites generally have a ne-grained
greyish groundmass and microporphyritic textures. They contain abun-
dant partially serpentinised olivine phenocrysts/xenocrysts in a ne-
grained poikilitic groundmass containing microphenocrysts of phlogo-
pite, which may be variably chloritised, ow-oriented melilite laths,
feldspathoids, apatite and abundant Ti-rich magnetite (Fig. 2D). Aln
oites
contain, in contrast to polzenites, also phenocrysts of clinopyroxene.
Tertiary basaltic volcanic rocks predominantly have microporphyritic
textures with a ne-grained groundmass. Basanite samples are generally
dominated by clusters of slightly serpentinised olivine along with elon-
gated brownish, rarely corroded clinopyroxene crystals with oscillatory
and/or hour-glass zoning. Twinned clinopyroxene forms isolated crystals
or larger glomerophyres. The basanite ne-grained groundmass contains
tiny plagioclase laths irregularly distributed together with smaller cli-
nopyroxene columns (Fig. 2E). Most of the tephrite samples are charac-
terised by the presence of both euhedral and partially corroded kaersutite
with apatite inclusions. Kaersutite is surrounded by a very ne-grained
hypocrystalline groundmass.
Tertiary alkaline lamprophyres, such as camptonite, monchiquite and
more alkaline leucocratic microsyenite, are especially common in the
Roztoky Intrusive Complex of the CSVC. A special form of camptonite to
monchiquite dykes occurs in the RPVC. The alkaline lamprophyres
typically contain macroscopic amphibole phenocrysts in a ne-grained
groundmass. The amphibole phenocrysts typically show corrosion rims
comprising Ti-rich magnetite. Clinopyroxene in camptonites forms
oscillatory and hour-glass zoned phenocrysts and smaller tabular crystals
in a ne-grained plagioclase-rich groundmass. A typical accessory min-
eral is apatite, which is commonly associated with amphibole and/or
clinopyroxene, titanomagnetite and feldspathoids. Zeolites are rare and
form vesicular llings (Fig. 2F). Camptonite to monchiquite samples from
RPVC are characterised by amphibole that is rimmed by distinct dark
mica, strongly serpentinised olivine and rare oscillatory-zoned clino-
pyroxene in a glassy matrix. Leucocratic microsyenite samples contain
glomerophyres of amphibole and alkaline feldspar in a ne-grained
feldspar-rich matrix (Fig. 2G). Amphiboles have inclusions of apatite
and ilmenite. The groundmass consists primarily of glass, plagioclase and
rare sodalite. Vesicles are lled by hydrothermal calcite. Very ne-
grained chilled margins of the leucocratic microsyenite dyke show
ow-aligned alkali feldspar microphenocrysts (Fig. 2A).
Additionally, two areas with nephelinte to melilitite occurrences are
included in this study. Cretaceous intrusive melilitite from the Great
Devils Wall in the RPVC is spatially related to ultramac lamprophyres
and forms a spectacular wall with horizontal columnar jointing (Fig. 2H).
It contains mildly corroded olivine often forming clusters, and strongly
zoned brown Ti-rich clinopyroxene commonly displaying characteristic
twinning. Opaque minerals and ow-oriented thin laths of clinopyroxene
occur in a melilite-bearing groundmass. In contrast, the Quaternary
melilitite effusion from the youngest volcanic eld in the Bohemian
Massif, i.e., Cheb Basin, is characterised by the mineral association
olivine (with thick, corroded rims) and large, kink-banded phlogopite
akes distributed in a vesicular glassy groundmass. For detailed petro-
graphical characteristic of individual studied samples see Supplement B.
4.2. Mineral chemistry of amphibole, clinopyroxene and dark mica
Chemical compositions of amphibole, clinopyroxene and dark mica
from camptonite/monchiquite, leucocratic microsyenite and polzenite
samples are listed in Supplement C available in the electronic appendix.
Analysed minerals have relatively homogenous compositions without
signicant variation between cores and rims.
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
928
Amphibole phenocrysts from both camptonite/monchiquite and
leucocratic microsyenite samples have relatively uniform contents of
CaO (11.612.8 wt.%) and Na
2
O (1.92.5 wt.%). Amphibole from leu-
cocratic microsyenite has higher concentrations of FeO
tot
(13.615.3
wt.%) and TiO
2
(4.65 wt.%) and lower contents of MgO (1011.2 wt.%)
and Al
2
O
3
(1213.4 wt.%) than those from camptonite/monchiquite
samples that have concentrations of MgO ¼1315.2 wt.%, Al
2
O
3
¼
13.614.6 wt.% and TiO
2
¼3.64.1 wt.%, except for three analyses that
yielded very high TiO
2
contents of 66.5 wt.%. Concentrations of FeO
tot
in amphibole from camptonite/monchiquite samples are between 7.4
wt.% and 10.4 wt.%. All analysed amphibole crystals have high Mg/(Mg
þFe) atomic ratios between 0.54 and 0.82 and have ~6 Si atoms per
formula unit (apfu) and ~2 Ca apfu, respectively (Supplement C.1).
Analysed amphiboles correspond both to pargasite and kaersutite
(Fig. 3A).
Clinopyroxene phenocrysts from camptonite/monchiquite and leu-
cocratic microsyenite samples have similar compositions (Supplement
C.2). They are Ca-rich with a relatively narrow range of CaO contents
(22.424.8 wt.%; ~1 Ca apfu) along with variable concentrations of MgO
(7.814.2 wt.%), Al
2
O
3
(4.813.1 wt.%), FeO
tot
(5.610.4 wt.%) and
TiO
2
(1.45.1 wt.%). They have only minor MnO (up to 0.4 wt.%) and
Na
2
O (below 1 wt.%) contents. Their Fe/(Fe þMg) atomic ratios range
between 0.19 and 0.43. All analysed clinopyroxene phenocrysts fall in
the diopside eld (Fig. 3B).
Dark mica from the polzenite sample shows a broad variation in MgO
(16.822.2 wt.%; ~1.82.4 Mg apfu) and Al
2
O
3
(1116.3 wt.%), and
plots in two compositional elds (see Fig. 3C). In contrast, dark mica
from camptonite/monchiquite is relatively uniform, having MgO and
Al
2
O
3
contents in the ranges of 16.317.3 wt.% and 15.916.4 wt.%,
respectively. The K
2
O contents are more variable in mica from polzenite
Fig. 2. Field appearances and petrographic features
of intrusive (lamprophyric) and extrusive volcanic
rocks from the
Cesk
eSt
redoho
rí (CSVC) and Ralsk
a
Pahorkatina (RPVC) volcanic complexes. (A) Contact
of a leucocratic microsyenite dyke in basanite host,
characterised by a very ne-grained chilled margin
(CSVC). (B) Columnar jointing of sandstone (see
Vel
azquez et al., 2008) at the contact to a camptonite
dyke (RPVC). (C) Polzenite with characteristic warty
surface (RPVC). (D) Partly serpentinised olivine (Ol)
within ne-grained groundmass with slightly chlori-
tised phlogopite (Phl), melilite (Mll) and Ti-rich
magnetite (Mag) in polzenite (PPL; RPVC). (E) Clus-
ter of olivine (Ol) and clinopyroxene (Cpx) display-
ing hour-glass zoning in groundmass composed of
smaller plagioclase (Plg) laths and tiny clinopyroxene
in basanite (PPL; RPVC). (F) Large amphibole (Amp)
phenocryst surrounded by corrosion rim with Ti-rich
magnetite (Mag) which encloses rounded vesicles
lled with zeolite (Zeo) in camptonite (PPL; RPVC).
(G) Glomeroporphyritic amphibole (Amp) and
plagioclase (Plg) in aphanitic groundmass in leuco-
cratic microsyenite (XPL; CSVC). (H) The Great
Devils Wall melilitite dyke displaying horizontal
columnar jointing (RPVC).
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
929
(~8.610.8 wt.%; ~0.81.0 K apfu) than in mica from camptonite/
monchiquite (8.89.4 wt.%; ~0.80.9 K apfu). The latter also contains
mica that is relatively poor in BaO (max. 1.8 wt.% BaO). FeO
tot
con-
centrations range between 6 wt.% and 10.2 wt.%, except for few dark
mica analyses from polzenite that show higher contents of 14.115.6
wt.%. Tetrahedral Al ranges between 0.94 and 1.41 apfu along with Mg/
(Mg þFe) atomic ratio of 0.660.88. Dark mica from polzenite falls in the
phlogopite and Mg-biotite elds, whereas dark mica from camptonite/
monchiquite samples falls in the phlogopite eld (Fig. 3C). Moreover,
phlogopite from polzenite with very high Mg/(Mg þFe) shows an
evolutionary trend towards tetra-ferriphlogopite, which is attested by
decient tetrahedral Al in recalculated analyses (Supplement C.3).
4.3. Whole-rock geochemistry
4.3.1. Major elements
The SiO
2
and MgO contents of a representative set of 36 samples were
recalculated on a volatile-free basis and range from 33 wt.% to 61 wt.%
and 0.1 wt.% to 18 wt.%, respectively (Fig. 4,Table 2). The most ultra-
basic rocks are Cretaceous ultramac lamprophyres (SiO
2
~36 wt.%)
followed by associated melilitic rocks (SiO
2
~40 wt.%). Tertiary volcanic
rocks compositionally range from ultrabasic nephelinite and basic teph-
rite/basanite (majority of the samples with SiO
2
~43 wt.%) to more
acidic rocks of phonolite composition. Tertiary alkaline lamprophyres are
basic (SiO
2
~46 wt.%) with highly variable contents of alkalies. One
sample of leucocratic microsyenite plots in the tephriphonolite eld
(Fig. 4).
The various samples dene a coherent trend in binary variation dia-
grams, using SiO
2
as differentiation index, that may be interpreted as (i)
fractionation trend or (ii) mixing/assimilation trend or (iii) superposition
of both (Fig. 5). Generally, Al
2
O
3
contents correlate positively, whereas
CaO and MgO are negatively correlated. In contrast, TiO
2
contents
behave differently. TiO
2
contents in Cretaceous ultramac lamprophyres
and melilitic rocks are in the range of 23 wt.% and correlate positively
with SiO
2
. Contrary to that, TiO
2
contents in the most primitive members
of the Tertiary volcanic rocks and alkaline lamprophyres reach ~4 wt.%
and correlate negatively with increasing SiO
2
. FeO
tot
behaves similarly as
TiO
2
(Fig. 5). K
2
O contents are ~2 wt.% in the majority of ultramac
lamprophyres and associated melilitic rocks. In the most primitive
members of Tertiary volcanic samples, K
2
O contents are below 1.5 wt.%
and correlate positively with increasing SiO
2
. An extremely high K
2
O
content of 8.3 wt.% was found in a leucocratic microsyenite (sample
KK7A).
4.3.2. Trace elements
The samples generally have highly variable transition metal contents
and a slightly variable enrichment of large-ion lithophile elements (LILE)
and light rare earth elements (LREE) relative to high-eld strength ele-
ments (HFSE) and heavy rare earth elements (HREE). This is clearly
visible both in binary variation diagrams and in primitive mantle nor-
malised trace and chondrite normalised rare earth element abundance
plots (Figs. 5 and 6AF).
The highest Cr concentrations (up to 800 ppm) occur in Cretaceous
ultramac lamprophyres and associated melilitic rocks. Chromium cor-
relates negatively with increasing SiO
2
, reaching levels as low as 2 ppm
Cr in differentiated trachyte (Fig. 5,Table 2). Cretaceous ultramac
lamprophyres show the highest degree of LREE enrichment. La, Ce and
Ce/Yb show a negative correlation with increasing SiO
2
for ultramac
Table 1
List of mantle-derived rocks of the Bohemian Massif selected for Pb isotope determination, sample location, petrographic types and eld characteristics.
No. Sample Locality Area Rock type Age (Ma) Latitude (N) Longitude (E) Outcrop characteristic
1 1_1367 V
sechlapy CSVC basanite 41.9 5037.10001347.7670Quarry
2 4_New Radobýl CSVC basanite 30 5031.85001405.7170Abandoned quarry
3 5_New Sout
esky CSVC basanite 30 5044.88301416.0170Quarry
4 8_1299
Zandov CSVC basanite 25.4 5042.48301423.8170Quarry
5 12_1295 Rade
sín CSVC trachybasalt 26.8 5041.88301403.5670Abandoned quarry
6 14_1358 Valke
rice CSVC trachybasalt 24.7 5042.11701419.7170Rock exposure
7 15_1359 Chlum CSVC tephrite 26.6 5044.81701413.5830Abandoned quarry
8 17_BM-8 Rýde
c CSVC phonolite 25.8 5036.433 01409.3000Abandoned quarry
9 18_BM-16
Strbický vrch CSVC phonolite 30 5033.46701350.5170Rock exposure
10 19_BM-48 Bo
re
n CSVC phonolite 30 5031.61701345.7170Rock exposure
11 20_BM-13 Lhenice CSVC trachyte 30 5034.50001352.1830Rock exposure
12 21_BM-51 Mile
sovský Kloc CSVC trachyte 30 5032.35001355.0670Rock exposure
13 23_BM-4 Kalich CSVC trachyandesite 31 5036.18301412.5670Abandoned quarry
14 25_BM-60 Bore
c CSVC trachyandesite 30 5030.85001359.2670Rock exposure
15 CS-30 Dobkovice I CSVC monchiquite 30 5042.63301411.5670Dyke in old quarry
16 CS-31 Dobkovice II CSVC camptonite 30 5042.64001411.5490Dyke in old quarry
17 CS-43 Le
stina CSVC camptonite 30 5039.36701412.3000Abandoned quarry
18 KK1 Komorní Hůrka CHB melilitite 1 5006.02001220.1660Rock exposure
19 KK2
Ríp CSVC tephrite 25.6 5023.23001417.3200Angular blocks
20 KK3 Pansk
ask
ala CSVC basanite 29.7 5046.14601429.1020Abandoned quarry
21 KK4 Tlustec RPVC basanite 30 5043.53501444.6490Angular blocks
22 KK5 St
ríbrník RPVC tephrite 30 5043.93601450.9250Rock exposure
23 KK6 JanůvDůl RPVC camptonite/monchiq uite 28.7 5042.10201457.3790Abandoned quarry
24 KK7A P
rední Lhota I CSVC leucocratic microsyenite 30 5042.52901412.0310Dyke in quarry - centre
25 KK8 P
rední Lhota II CSVC camptonite 30 5042.55701412.0180Dyke in quarry
26 KK9 Sv
arov RPVC polzenite 70 5042.31301453.4500Angular blocks
27 KK10 Velk
a
Certova ze
d RPVC melilitite 70 5040.41401456.7250Dyke exposure
28 KK11 Hamerský
Spi
c
ak I RPVC polzenite 70 5041.33601450.9890Angular blocks
29 KK12 Pta
cí vr
sek RPVC basanite 70 5040.50901440.7820Dyke in old quarry
30 ME-3/13 Krkav
cí sk
ala I CSVC nephelinite 27 5035.15001404.7830Rock exposure
31 ME-4/13 Krkav
cí sk
ala II CSVC basanite 30 5035.11701404.7170Rock exposure
32 OC-1 Veselí RPVC camptonite/monchiquite 30 5038.28001438.4600Angular blocks
33 OC-2 Pelousek RPVC polzenite 70 5040.68001458.2000Abandoned quarry
34 OC-9 Vesec RPVC polzenite 68.4 5042.18001458.9800Angular blocks
35 OC-10 Nový Luhov RPVC aln
oite 70 5042.48001445.0000Angular blocks
36 OC-12 Holi
cský vrch RPVC melilitolite 70 5040.80001453.8200Borehole
CSVC
Cesk
eSt
redoho
rí Volcanic Complex; RPVC Ralsk
a Pahorkatina Volcanic Complex; CHB Cheb Basin.
Ages are taken from Ulrych et al. (1998, 2002, 2013, 2014, 2018), Skala et al. (2014), Ackerman et al. (2015), Dostal et al. (2017).
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
930
lamprophyres towards tephritic and basanitc rocks (Fig. 5).
The Cretaceous ultramac lamprophyres and associated melilitic
rocks are characterised by troughs for K and Pb in primitive mantle
normalised trace element plots (Fig. 6A). Tertiary alkaline lamprophyres
show slightly negative or even positive anomalies for K along with Pb
enrichment and P depletion. The leucocratic microsyenite has the most
pronounced anomalies among the analysed alkaline lamprophyres and
related rocks (Fig. 6C). Absolute trace element concentrations in Tertiary
volcanic samples differ from those of the basanite/tephrite samples and
the more evolved trachytic/phonolitic samples. However, they dene the
same trends in primitive mantle normalised trace element plots. Tertiary
volcanic samples show variable Rb depletion, K and Pb depletion or
enrichment, as well as troughs for P and Ti that markedly increase from
basanite/tephrite to trachyte/phonolite samples (Fig. 6E). There are two
geochemical types of phonolite in the representative sample set: type A
(Sr-rich) and type B (Sr-poor), rst characterised by Ackerman et al.
(2015). Whereas the samples of type A phonolite compositionally
resemble other types of evolved volcanic rocks, type B phonolite is
prominent by extreme depletion of Ba, Sr, P, and Ti along the most
pronounced Cs, U, and Pb enrichments among all Tertiary volcanic rocks
(Fig. 6E).
The samples show variable REE contents and generally lack pro-
nounced Eu anomalies (Fig. 6B, D, F). Cretaceous ultramac lamp-
rophyres have the highest total REE contents (ΣREE ~500 ppm), whereas
Tertiary trachytic/phonolitic samples have the lowest total REE contents
οf ~270 ppm. Ultramac lamprophyres together with associated melilitic
samples are prominent by the highest enrichment in LREE over HREE
with Ce
N
/Yb
N
of ~30 and Lu
N
of ~7, whereas Tertiary volcanic samples
have contrasting enrichment trends in LREE/HREE (Ce
N
/Yb
N
~21 in
basanite/tephrite samples, and ~31 in more differentiated trachyte/
phonolite samples). Type B phonolite is prominent by its U-shaped REE
normalised pattern (Fig. 6F). Ultramac lamprophyres, associated meli-
litic rocks, the majority of alkaline lamprophyres and basanite/tephrite
volcanic rocks all have Lu
N
<10 indicating the presence of garnet in the
mantle source of their parent melts (see Wilson and Downes, 1991).
Fig. 3. Classication of amphibole, clinopyroxene and dark mica from camp-
tonite/monchiquite, leucocratic microsyenite, and polzenite. All chemical ele-
ments are given in apfu (atoms per formula unit). (A) Analyses of amphibole
phenocrysts from camptonite/monchiquite samples plot in the elds of both
pargasite and kaersutite, whereas amphibole in leucocratic microsyenite falls in
the kaersutite eld of the amphibole classication scheme of Leake (1997). (B)
Clinopyroxene from both camptonite/monchiquite and leucocratic microsyenite
samples is diopside (Morimoto, 1988), diagram adopted from Rapprich (2005).
(C) Dark mica from camptonite/monchiquite and polzenite corresponds to
phlogopite. Dark mica from camptonite/monchiquite is compositionally rather
homogeneous, whereas dark mica from polzenite falls in two groups ranging
from phlogopite to Mg-biotite. Classication diagram after Rieder et al. (1998).
Fig. 4. Total Alkali-Silica (TAS) diagram showing the chemical composition of
analysed samples (after Le Maitre, 2002).
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
931
Table 2
Major oxide (wt.%) and trace element compositions (ppm) of mantle-derived rocks of the Bohemian Massif.
No. 1
a
2
a
3
a
4
a
5
a
6
a
7
a
8
a
9
a
10
a
11
a
12
a
13
a
14
a
15
b
16
b
Sample 1_1367 4_New 5_New 8_1299 12_1295 14_1358 15_1359 17_BM-8 18_BM-16 19_BM-48 20_BM-13 21_BM-51 23_BM-4 25_BM-60 CS-30 CS-31
SiO
2
40.2 40.4 42.3 41.6 47.0 45.2 42.2 55.8 55.2 54.2 59.1 56.8 53.9 54.6 45.9 47.8
TiO
2
2.65 3.73 3.3 2.8 2.97 3.59 3.92 0.14 0.59 0.28 0.57 0.45 1.27 0.86 3.22 2.13
Al
2
O
3
11.1 12.1 14.7 13.5 15.2 15.5 13.6 21.4 20.3 21.7 19.8 20.2 18.9 19.7 15.0 16.2
Fe
2
O
3
tot
12.5 14.0 12.1 11.5 9.6 10.8 11.5 2.3 3.63 2.2 3.22 3.11 5.66 4.36 11.1 8.34
MnO 0.19 0.19 0.19 0.19 0.16 0.18 0.33 0.28 0.23 0.26 0.23 0.22 0.2 0.2 0.16 0.16
MgO 12.9 8.91 8.4 9.87 4.57 5.4 6.53 0.08 0.43 0.23 0.21 0.34 1.32 0.67 4.91 2.48
CaO 12.5 12.5 11.8 12.6 9.85 10.6 13.5 0.95 4.15 1.8 3.4 3.59 5.4 4.42 8.79 7.35
Na
2
O2.73 1.85 2.25 1.84 3.67 3.38 2.61 9.9 6.85 9.11 5.0 6.47 6.47 5.94 2.92 3.7
K
2
O0.66 1.47 1.63 1.31 2.31 1.96 1.38 5.46 5.05 4.73 5.69 5.39 3.84 5.06 4.09 4.07
P
2
O
5
0.79 0.68 0.54 0.54 0.49 0.51 0.65 0.02 0.1 0.03 0.08 0.07 0.35 0.15 0.47 0.52
LOI 2.61 3.2 2.69 2.99 3.62 3.27 4.31 1.79 2.16 5.52 2.47 3.22 2.77 4.51 3.44 6.86
Total 98.8 99.1 99.9 98.7 99.4 100.4 100.5 98.2 98.7 100.1 99.8 99.8 100.1 100.5 99.9 99.6
Mg# 67 56 58 63 49 50 53 6 19 17 11 18 32 23 47 37
Cr 350 210 140 320 50 40 70 10 10 2 2 10 20 10 27 16
Ni 280 170 120 190 10 70 90 10 10 10 10 10 20 10 22 15
Co 51 45 40 44 26 30 34 1 14 4 5 7 14 7 31 17
Sc 26 29 30 32 20 24 30 1 1 1 0 0 4 1 24 10
V281 336 345 317 336 378 436 23 71 18 60 57 108 100 367 212
Cu 70 60 70 90 40 60 180 5 10 5 5 5 10 5 ––
Zn 110 110 100 100 90 110 120 150 130 175 118 110 130 120 ––
Rb 10 36 33 29 144 84 44 371 123 219 184 153 98 121 81 117
Cs 0.5 0.5 0.0 1.7 1.7 1.4 1.2 9.9 1.4 4.78 5.45 2.0 1.70 2.5 0.9 1.5
Ba 643 600 494 700 992 776 752 17 1880 620 1870 1730 1400 1420 781 937
Sr 981 1260 1290 1000 937 904 1070 21 2070 433 1830 1560 1570 1410 641 1000
Ga 18 19 19 18 20 24 24 53 27 25 41 27 29 26 ––
Ta 4.6 7.1 4.5 5.7 4.4 5.5 5.7 2.02 5.5 7.2 8.56 4.1 6.9 5.0 4.84 6.13
Nb 68 70 66 79 70 78 76 138 140 447 198 113 127 110 94 123
Hf 5.0 6.2 8.0 5.3 6.7 8.1 7.9 21.2 10.2 22.2 11.1 10.0 13.1 10.4 9.2 12.4
Zr 228 232 211 218 324 329 323 1200 543 1650 634 568 615 562 359 576
Y23 27 22 22 22 22 24 16 24 24 29 20 27 22 21 25
Pb 4 4 3 4 7 7 7 34 9 12 25 11 15 11 6 7
Th 7.5 4.2 5.0 6.5 7.9 8.9 9.0 49.7 14.8 28.8 15.4 15.3 17.9 14.6 8.9 11.8
U1.9 1.1 1.4 1.6 2.3 3.1 5.7 18.6 3.8 16.6 4.0 3.9 4.1 3.4 2.0 3.4
La 68 46 44 60 58 51 69 139 124 115 148 112 108 110 49 77
Ce 124 96.7 87.7 112 112 105 133 146 210 160 216 178 186 181 101 147
Pr 13.7 12.3 10.3 12.3 12.9 12.4 15.4 8.5 19.5 22.0 13.0 15.9 18.7 17.6 11.8 16.4
Nd 50.2 49.2 39.4 43.8 48.2 47.4 57.6 17.3 56.1 45.8 67.3 44.0 60.9 54.4 45.8 59.3
Sm 9.2 10.1 7.5 7.9 8.6 8.5 10.4 1.61 6.8 6.72 8.2 5.6 9.3 7.5 8.58 9.84
Eu 2.92 3.21 2.35 2.5 2.49 2.53 3.08 0.32 2.01 1.34 2.06 1.56 2.71 2.18 2.48 2.89
Gd 7.7 8.4 6.3 6.5 6.5 6.5 8.2 1.13 4.9 3.9 4.4 3.7 6.9 5.4 8.21 9.88
Tb 1.0 1.2 0.9 0.9 0.9 0.9 1.1 0.22 0.7 0.69 0.8 0.5 0.9 0.7 1.03 1.23
Dy 5.1 5.8 4.7 4.8 4.6 4.6 5.3 1.56 3.7 5.3 5.7 3.0 5.0 3.9 4.84 5.62
Ho 0.9 1.0 0.9 0.9 0.8 0.8 0.9 0.41 0.7 0.9 0.9 0.6 0.9 0.8 0.85 1.0
Er 2.3 2.6 2.2 2.3 2.2 2.2 2.3 1.6 2.4 2.8 2.8 2.0 2.7 2.2 2.35 2.88
Tm 0.3 0.34 0.29 0.29 0.29 0.3 0.32 0.34 0.36 0.45 0.52 0.32 0.39 0.33 0.28 0.36
Yb 1.8 1.9 1.9 1.8 1.9 1.8 1.8 2.98 2.4 3.94 2.88 2.3 2.6 2.2 1.81 2.37
Lu 0.25 0.26 0.28 0.27 0.28 0.26 0.26 0.52 0.4 0.61 0.48 0.37 0.41 0.35 0.27 0.34
No. 17
b
18 19 20 21 22 23 24 25 26 27 28 29 30
c
31
c
32
d
33
d
34
d
35
d
36
d
Sample CS-43 KK1 KK2 KK3 KK4 KK5 KK6 KK7A KK8 KK9 KK10 KK11 KK12 ME-3/13 ME-4/13 OC-1 OC-2 OC-9 OC-10 OC-12
SiO
2
42.7 38.2 39.2 42.4 41.7 40.6 44.9 52.8 45.2 33.8 39.9 33.1 41.3 36.8 39.9 38.9 34.0 30.6 34.1 31.5
TiO
2
2.97 3.1 3.4 3.17 3.73 3.22 2.62 1.45 3.25 2.62 2.96 2.47 3.08 3.07 2.63 3.43 2.57 2.28 2.1 2.57
Al
2
O
3
13.8 11.5 11.8 13.6 14.5 14.7 13.0 17.8 15.8 7.88 9.99 7.74 14.5 12.6 12.2 13.7 8.96 7.38 8.24 8.65
(continued on next page)
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
932
Table 2 (continued )
No. 17
b
18 19 20 21 22 23 24 25 26 27 28 29 30
c
31
c
32
d
33
d
34
d
35
d
36
d
Sample CS-43 KK1 KK2 KK3 KK4 KK5 KK6 KK7A KK8 KK9 KK10 KK11 KK12 ME-3/13 ME-4/13 OC-1 OC-2 OC-9 OC-10 OC-12
Fe
2
O
3
tot
11.6 13.2 15.6 12.4 12.5 13.3 11.9 5.58 10.4 11.7 11.5 10.8 13.0 13.6 12.9 12.3 12.1 11.6 13.7 10.7
MnO 0.19 0.23 0.33 0.18 0.2 0.22 0.23 0.15 0.17 0.2 0.18 0.19 0.21 0.22 0.18 0.16 0.2 0.2 0.2 0.17
MgO 7.88 12.1 7.15 9.15 7.07 8.36 5.94 1.79 4.71 15.3 15.7 16.9 7.90 11.4 13.6 11.0 15.1 17.0 16.5 13.4
CaO 10.5 12.4 12.0 12.0 11.8 12.4 11.8 4.28 8.07 16.1 13.0 18.6 12.7 14.7 12.1 13.2 14.1 21.1 16.4 17.9
Na
2
O3.03 3.71 4.39 3.13 3.03 2.92 2.62 3.79 4.03 1.02 2.31 1.59 3.3 3.19 2.02 0.89 2.16 0.52 2.01 0.22
K
2
O2.15 2.15 1.1 1.57 1.75 1.53 1.67 7.97 4.17 1.72 1.41 1.25 1.33 1.5 0.96 1.9 1.85 0.75 1.66 1.98
P
2
O
5
0.6 0.92 1.41 0.73 0.77 1.09 1.23 0.32 0.52 1.24 0.75 1.30 0.87 1.5 0.63 0.53 1.07 1.12 1.09 0.95
LOI 4.57 1.8 3.0 1.1 2.3 1.2 3.6 3.7 3.2 7.6 1.4 5.2 1.4 1.8 3.5 5.18 7.4 6.3 4.9 12.0
Total 100.0 99.4 99.3 99.4 99.4 99.5 99.5 99.6 99.5 99.2 99.2 99.2 99.5 100.4 100.6 101.1 99.6 98.8 100.4 100.0
Mg# 57 65 48 59 53 56 50 39 47 72 73 76 55 63 68 64 71 74 71 71
Cr 70 431 27 294 144 130 55 14 34 602 842 746 192 197 486 262 749 810 623 700
Ni 75 147 25 98 41 54 27 5 12 317 339 336 56 132 239 134 327 330 253 216
Co 37 56 37 46 40 43 31 9 28 53 59 60 43 55 60 51 56 59 61 47
Sc 30 29 21 27 26 27 17 5 20 22 30 26 24 27 31 45 33 23 30 22
V345 304 275 289 348 336 269 141 341 252 308 210 315 378 334 ––252 ––
Cu 48 38 70 61 47 38 9 46 52 76 48 56 –––30 ––
Zn 63 150 90 100 88 91 77 82 86 72 65 84 –––70 ––
Rb 38 54 21 35 64 37 48 222 118 49 51 47 32 29 35 154 64 26 48 65
Cs 1.3 0.4 1.0 0.4 0.6 0.2 1.2 1.80 1.2 0.9 1.0 0.8 0.4 0.8 0.5 1.89 2.9 0.5 0.93 3.03
Ba 654 965 1030 717 791 591 541 1070 868 737 807 968 726 1470 651 1610 1980 717 1380 754
Sr 920 807 1580 937 1700 940 1230 757 890 866 873 1420 1140 1770 963 854 2000 1800 1780 1280
Ga 21 25 19 22 20 17 21 21 13 14 12 19 –––18 ––
Ta 4.9 6.8 10.1 4.5 4.8 5.7 5.7 5.2 5.1 8.5 5.1 6.9 5.3 5.8 4.1 3.49 5.5 10.7 7.80 7.73
Nb 95 122 168 87 78 88 96 107 82 129 84 134 90 141 96 79 188 151 183 125
Hf 8.1 6.7 14.7 5.8 8.2 6.3 7.0 9.7 8.2 5.5 5.7 4.6 6.1 8.5 7.1 10.2 7.5 5.5 6.3 7.8
Zr 313 299 616 233 351 256 290 461 340 251 224 197 243 333 259 301 357 247 292 345
Y24 27 42 25 29 28 30 23 23 30 19 24 27 34 21 21 26 29 28 29
Pb 1725 64 24 15853 4 2 10 3 4 6 4 6 4
Th 7.6 11.3 13.8 6.9 6.6 7.1 8.0 18.4 10.0 16.1 6.4 12.7 6.9 17.0 7.0 12.3 15.0 23.0 17.0 10.2
U2.1 3.1 3.9 1.5 1.8 1.6 2.0 5.2 2.5 3.9 1.6 2.9 1.5 3.7 1.7 1.45 3.9 4.83 4.8 2.49
La 62 89 125 58 63 67 73 80 57 124 52 106 62 165 64 51 103 163 130 84
Ce 128 168 249 110 132 136 144 145 115 248 104 201 125 289 117 104 186 304 235 177
Pr 14.6 18.9 28.7 12.1 15.4 15.4 16.9 14.8 13.0 28.2 11.9 22.2 14.7 28.0 13.0 12.5 21.0 33.8 25.0 19.9
Nd 56.9 69.3 110 46.1 61.5 59.6 61.5 50.4 50.4 106 46.1 84.3 56.9 102 48.0 50.6 79.0 120 90.0 77.2
Sm 10.3 11.4 18.0 8.5 11.0 10.7 10.4 7.69 8.4 17.1 8.61 13.9 9.67 16.0 8.4 9.27 13.0 18.4 14.0 13.4
Eu 3.02 3.33 5.01 2.52 3.3 3.11 3.13 2.06 2.42 4.73 2.56 3.92 3.07 4.5 2.5 2.74 3.6 5.13 3.9 3.85
Gd 9.74 9.34 14.2 7.51 9.39 8.98 9.05 6.11 6.97 12.7 7.06 10.6 8.29 15.0 8.1 8.61 10.0 12.9 13.0 13.9
Tb 1.22 1.16 1.75 0.98 1.2 1.11 1.16 0.83 0.9 1.52 0.86 1.23 1.08 1.7 1.0 1.03 1.30 1.53 1.5 1.62
Dy 5.51 5.88 8.98 5.13 6.31 6.35 6.50 4.58 4.93 7.69 4.28 6.03 5.88 7.6 4.8 4.77 6.6 7.08 6.3 6.28
Ho 0.96 0.95 1.52 0.88 1.02 1.04 1.07 0.75 0.84 1.12 0.67 0.91 0.94 1.3 0.85 0.82 1.10 1.11 1.0 0.97
Er 2.72 2.46 3.93 2.29 2.83 2.86 2.83 2.31 2.28 2.66 1.67 2.18 2.53 3.6 2.3 2.19 2.6 2.61 2.6 2.87
Tm 0.33 0.31 0.56 0.28 0.35 0.38 0.37 0.32 0.31 0.32 0.19 0.25 0.33 0.42 0.27 0.26 0.33 0.3 0.3 0.29
Yb 2.05 1.89 3.33 1.89 2.16 2.15 2.46 2.30 1.94 1.84 1.26 1.49 1.92 2.6 1.7 1.92 1.9 1.63 1.8 1.82
Lu 0.29 0.26 0.49 0.25 0.33 0.34 0.35 0.33 0.3 0.25 0.16 0.19 0.29 0.36 0.23 0.22 0.25 0.21 0.23 0.25
LOI - loss on ignition; Mg# ¼100 Mg/(Mg þFe
tot
).
a
Analyses adopted from Dostal et al. (2017).
b
Analyses adopted from Sk
ala et al. (2014).
c
Analyses adopted from Sk
ala et al. (2015).
d
Analyses adopted from Ulrych et al. (2014).
S. Krmí
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933
4.4. Lead isotope composition
The initial lead isotope composition of Cretaceous ultramac lamp-
rophyres, Tertiary basaltic volcanic samples as well as alkaline lamp-
rophyres, related leucocratic microsyenite and Cretaceous and
Quaternary melilitite rocks denes a scattered linear trend between two
end-members (schematically marked by red asterisks A and B in Figs. 7, 9
and 10). In the
207
Pb/
204
Pb vs.
206
Pb/
204
Pb diagram, most data fall above
the Northern Hemisphere Reference Line (NHRL; Hart, 1984), whereas in
the
208
Pb/
204
Pb vs.
206
Pb/
204
Pb diagram, the Pb trend intersects the
NHRL, as it is slightly atter than the NHRL (Fig. 7).
The Cretaceous ultramac lamprophyres and melilitolite sample of
the RPVC sampled a more radiogenic mantle component (
206
Pb/
204
Pb ¼
19.6220.01,
207
Pb/
204
Pb ¼15.6115.64,
208
Pb/
204
Pb ¼39.2939.77)
than the Tertiary alkaline lamprophyres and associated leucocratic
microsyenite (
206
Pb/
204
Pb ¼19.0519.35,
207
Pb/
204
Pb ¼15.6015.61,
208
Pb/
204
Pb ¼38.9139.14; Table 3). The Cenozoic volcanic samples
have a very heterogeneous Pb isotope composition (
206
Pb/
204
Pb ¼
18.9819.96,
207
Pb/
204
Pb ¼15.5915.65,
208
Pb/
204
Pb ¼38.8039.62).
Data points plot along a linear array extending from the least radiogenic
Pb isotope compositions of the alkaline lamprophyres to the most
radiogenic compositions of the ultramac lamprophyres (Fig. 7).
Fig. 5. SiO
2
vs. selected major and trace element contents of analysed samples.
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
934
Intrusive and effusive melilitite samples, which are of Cretaceous and
Quaternary age, respectively, have essentially identical Pb isotope com-
positions (
206
Pb/
204
Pb ¼19.3019.36,
207
Pb/
204
Pb ¼15.60,
208
Pb/
204
Pb ¼39.0939.10).
Among the ultramac lamprophyres, aln
oite OC-10 with a
206
Pb/
204
Pb ratio of 20.01 is the most radiogenic sample and melili-
tolite OC-12 with
206
Pb/
204
Pb ratio of 19.62 is the least radiogenic
sample. The Pb isotope compositions of polzenite samples fall between
these two values (Fig. 7). The monchiquites from the CSVC as well as
the camptonite to monchiquite sample OC-1 are among the least
radiogenic samples from the Bohemian Massif. They have
206
Pb/
204
Pb
ratios of 19.0519.10. Similarly, the Pb isotope signature of the leu-
cocratic microsyenite is very unradiogenic and points to derivation
from the same mantle source as the less fractionated camptonites and
monchiquites.
5. Discussion
5.1. Chemically contrasting mantle sources in the Bohemian Massif
Common K-bearing silicate phases such as phlogopite or K-amphibole
are not stable in the asthenospheric (convective) mantle (Class and
Goldstein, 1997). Negative K anomalies observed in intraplate
mantle-derived rocks indicate that these phases were stable during the
petrogenesis of these rocks, and, thus, are considered as geochemical
evidence for the involvement of lithospheric mantle sources (e.g., Wilson
and Downes, 1992; Melluso et al., 2011; Pf
ander et al., 2018).
Variations in the whole-rock composition of Late Cretaceous to
Quaternary undifferentiated magmatic rocks from the Bohemian Massif
point to contributions from two geochemically distinct mantle sources.
The Cretaceous ultramac lamprophyres show strong LREE over HREE
enrichment (Fig. 6B) in combination with high K
2
O and low TiO
2
at given
SiO
2
(Fig. 5) and signicant troughs in K and Pb in primitive mantle
Fig. 6. Abundances of trace elements and rare earth elements (REE) in the analysed samples normalised to primitive mantle values (McDonough and Sun, 1995) and
chondritic abundances (Boynton, 1984), respectively.
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
935
normalised trace element plots (Fig. 6A), indicating that they originated
by low degrees of partial melting of a metasomatised lithospheric mantle
source containing a stable K-bearing residual phase (e.g., Ulrych et al.,
2008, 2014; Sk
ala et al., 2015). Such a lithospheric mantle formed during
Variscan subduction of continental crust. This process led to stabilization
of metasomatised parts containing inclusions of granitic composition
within the mantle rocks beneath the Saxothuringian and Moldanubian
zones of the Bohemian Massif (Borghini et al., 2018; Ferrero et al., 2018
and references therein). This is in line with the composition of phlogopite
in ultramac lamprophyres that shows high MgO and elevated K
2
O and
BaO contents, coupled with a compositional afnity towards the
tetra-ferriphlogopite end-member (cf. Sk
ala et al., 2015). Moreover,
K-rich sulphides of rasvumite (KFe
2
S
3
) type, identied by Ulrych et al.
(2008), occur locally in association with phlogopite in melilitolites. Such
a K-rich sulphide may be a primary phase that crystallised from meta-
somatised lithospheric mantle-derived ultramac melts (e.g., Sharygin
et al., 2008).
Compared with the source of Cretaceous ultramac lamprophyres,
the second mantle source is characterised by signicant TiO
2
enrichment
and K
2
O depletion. This resulted in relatively at OIB-like primitive
mantle normalised trace element patterns for the majority of the Tertiary
alkaline lamprophyres (characterised by the presence of TiO
2
-rich
amphibole kaersutite with up to 0.74 apfu Ti) and basanite/tephrite
volcanic rocks (Fig. 6C, E). Their mantle source was signicantly inu-
enced by interaction with asthenospheric (convective) mantle depleted
in large-ion lithophile elements (Rock, 1991). This is in line with
rift-related passive asthenospheric upwelling that resulted in the gener-
ation of large volumes of Tertiary mantle-derived magmas in the Eger
Rift (e.g., Ulrych et al., 2011).
Both mantle sources are distinguishable: undifferentiated Cretaceous
ultramac lamprophyres show a trend with relatively large variations in
Nb/Ta and narrow variation in Lu/Hf, whereas Tertiary alkaline lamp-
rophyres and basanite/tephrite volcanic rocks show a narrow range of
Nb/Ta values as also known from OIBs(Fig. 8A). The higher and more
variable Nb/Ta ratios in the Cretaceous ultramac lamprophyres likely
indicate a higher contribution from the lithospheric mantle and a more
heterogeneous source than for the Tertiary alkaline lamprophyres and
the basanite/tephrite volcanic rocks (cf. Pf
ander et al., 2012). The
Cretaceous ultramac lamprophyres and undifferentiated Cenozoic vol-
canic rocks show a comparable range in Nb/La ratios at different
(Ce/Yb)
N
(Fig. 8B). The Nb/La ratios fall between the values for depleted
MORB mantle source and for the lithospheric mantle source meta-
somatised by continental crust (Fig. 8B). The narrow range of Nb/La
ratios may be indicative for a similarly enriched mantle source, whereas
different (Ce/Yb)
N
may reect contrasting degrees of partial melting (cf.
Haase and Renno, 2008). The Cretaceous ultramac lamprophyres and
undifferentiated Cenozoic volcanic rocks show different trends in the
Ba/La vs. (Ce/Yb)
N
diagram (Fig. 8C), which may indicate involvement
of different mantle sources. For instance, the high and variable (Ce/Yb)
N
and relatively constant Ba/La ratios of Cretaceous ultramac lamp-
rophyres may reect different degrees of partial melting of highly met-
asomatised domains of the lithospheric mantle, producing such
SiO
2
-undersaturated melts (cf. Hegner et al., 1995; Blusztajn and Hegner,
2002). In contrast, Cenozoic volcanic rocks with relatively constant
(Ce/Yb)
N
and highly variable Ba/La may reect a carbonatite component
(Haase et al., 2017), which may be indicative for a mantle source that
was affected by interaction with the convective mantle. Signicant CO
2
release from the asthenosphere in a continental rifting environment is
also indicated by metasomatised mantle xenoliths from the lithospheric
mantle beneath the Bohemian Massif (Ackerman et al., 2013). Besides
these compositional differences related to different degrees of partial
melting of heterogeneously metasomatised mantle sources, fractional
crystallisation also affected the chemical composition of the rocks (Dostal
et al., 2017). Fractional crystallisation in particular affected the compo-
sition of Tertiary trachytic to phonolitic volcanic rocks and the leuco-
cratic microsyenite of the CSVC and is reected in linear enrichment
trends for SiO
2
,Al
2
O
3
and K
2
O and depletion trends for CaO, MgO,
FeO
tot
,Cr
2
O
3
,P
2
O
5
or TiO
2
(Figs. 5 and 6).
5.2. Lead isotope compositions of the samples asthenospheric vs.
lithospheric mantle components
The lead isotope data dene a relatively narrow two-component
mixing trend ranging from moderately high
206
Pb/
204
Pb ratios (~19.0)
to more radiogenic values (up to 20.0), suggesting a mixture between a
mantle source showing geochemical characteristics of modication by
interaction with convective mantle and metasomatically enriched litho-
spheric mantle (cf. Blusztajn and Hart, 1989; Stracke et al., 2005; Ulrych
et al., 2016 and section 5.1 in this paper). The suggested two-component
Fig. 7. Initial Pb isotope composition of individual samples from the Bohemian
Massif. Melilitite samples of different age (Cretaceous and Quaternary) are
plotted separately as their Pb isotope composition is relatively uniform but
different from other melilitic rocks (i.e., ultramac lamprophyres). NHRL:
Northern Hemisphere Reference Line (Hart, 1984). Abbreviations and refer-
ences for other volcanic elds from the Bohemian Massif: WB Western
Bohemia (Ulrych et al., 2016); LSB Lower Silesian Basin (Blusztajn and Hart,
1989); EG Erzgebirge, ESG Elbsandsteingebirge, LL Lower Lusatia (Haase
and Renno, 2008).
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
936
mixing trend crosses the NHRL of Hart (1984) and most samples have
higher
207
Pb/
204
Pb at given
206
Pb/
204
Pb than the NHRL (Fig. 7). There is
no correlation between SiO
2
and
207
Pb/
204
Pb. The absence of crustal
xenoliths excludes substantial crustal contamination of the
mantle-derived magmas during fractionation of the more acidic rocks en
route (cf. Blusztajn and Hart, 1989; Ulrych et al., 2016).
Late Cretaceous ultramac lamprophyres, including polzenite, aln
oite
and related melilitolite cluster at the radiogenic end of the Pb array
(Fig. 7), suggesting that their parent magma resulted from decompres-
sion melting of metasomatically enriched lithospheric mantle beneath
the Bohemian Massif (cf. Embey-Isztin et al., 1993a; Riley et al., 2003;
Kroner et al., 2010; Rudnick and Gao, 2014; Krmí
cek et al., 2016). Ter-
tiary alkaline lamprophyres, comprising camptonite, monchiquite and
related leucocratic microsyenite, which is a more felsic equivalent of
alkaline lamprophyre, fall near the low-
206
Pb/
204
Pb end-member of the
mixing trend (Fig. 7). This suggests that they were produced via subse-
quent interaction of the lithospheric mantle with asthenosphere-derived
material during the extensional phase of the Eger Rift opening (cf. Gal-
lagher and Hawkesworth, 1992; Wilson, 1993; Harangi and Lenkey,
2007; Aghazadeh et al., 2015; Dostal et al., 2019a). Other Tertiary vol-
canic samples of the Bohemian Massif are distributed between above two
end-members. The variations in
206
Pb/
204
Pb do not depend on petro-
graphic rock type or on the degree of fractionation, but instead, reect
various contributions from contrasting mantle sources during melting.
Cretaceous (pre-rift) intrusive and Quaternary (late-rift) effusive
melilitites have Pb isotope compositions very similar to those of alkaline
lamprophyres (Fig. 7). The Pb isotope data demonstrate that the Creta-
ceous melilitites (e.g., the Great Devils Wall) did not originate from the
same mantle source as the polzenites, despite emplacement in a similar
tectonic stress eld (Ulrych et al., 2014). The source of the melilitites was
more primitive, which may have been a direct consequence of higher
degree of melting (compared to the LREE enriched ultramac lamp-
rophyres), different melting depths, or a combination of both (cf. Hof-
mann, 1988; Sun and McDonough, 1989; Melluso et al., 2011).
5.3. Regionally contrasting Pb isotope signatures in mantle-derived rocks in
the Bohemian Massif
On the scale of the Bohemian Massif, most mantle-derived rocks
follow the same orogenic Pb growth curve in the
207
Pb/
204
Pb vs.
206
Pb/
204
Pb diagram (Figs. 7 and 9). As the metasomatised lithospheric
mantle has higher Pb contents than the asthenospheric mantle, small or
moderate contributions from other reservoirs than the lithospheric
mantle may not signicantly affect the Pb isotope composition of the
various melts and the character of the lithospheric mantle did not change
over time. The oldest mantle-derived dyke rocks emplaced during the
extensional phase after the nal collision of Laurussia, Gondwana and
Gondwana-derived terranes are orogenic lamprophyres and lamproites
(e.g., Krmí
cek, 2010; Krmí
cek et al., 2011). The orogenic Pb isotope
signature as sampled by mantle-derived rocks of the Bohemian Massif
since ca. 340 Myr predominantly originates from a metasomatised lith-
ospheric mantle source affected by subduction of oceanic and continental
crust during the Variscan orogeny (Krmí
cek et al., 2016; Dostal et al.,
2019a). The lower Pb isotope ratios of Variscan lamproites and lamp-
rophyres from the Bohemian Massif (Abdelfadil et al., 2013; Krmí
cek
et al., 2016) do not reect a different source than the Cretaceous to
Table 3
Whole-rock Pb isotope data of the mantle-derived rocks of the Bohemian Massif.
No. Sample Area Rock type Age (Ma)
206
Pb/
204
Pb
207
Pb/
204
Pb
208
Pb/
204
Pb
206
Pb/
204
Pb (T)
207
Pb/
204
Pb (T)
208
Pb/
204
Pb (T)
1 1_1367 CSVC basanite 41.9 19.752 15.620 39.489 19.55 15.61 39.23
2 4_New CSVC basanite 30 19.218 15.604 38.983 19.14 15.60 38.88
3 5_New CSVC basanite 30 20.001 15.647 39.619 19.86 15.64 39.45
4 8_1299 CSVC basanite 25.4 19.902 15.644 39.652 19.80 15.64 39.51
5 12_1295 CSVC trachybasalt 26.8 19.222 15.599 39.060 19.13 15.59 38.96
6 14_1358 CSVC trachybasalt 24.7 19.280 15.610 39.127 19.17 15.60 39.02
7 15_1359 CSVC tephrite 26.6 19.781 15.650 39.356 19.56 15.64 39.24
8 17_BM-8 CSVC phonolite 25.8 19.577 15.652 39.343 19.43 15.65 39.22
9 18_BM-16 CSVC phonolite 30 19.763 15.645 39.510 19.63 15.64 39.34
10 19_BM-48 CSVC phonolite 30 20.021 15.655 39.551 19.83 15.65 39.44
11 20_BM-13 CSVC trachyte 30 19. 426 15.623 39.270 19.36 15.62 39.18
12 21_BM-51 CSVC trachyte 30 19. 756 15.639 39.497 19.65 15.63 39.36
13 23_BM-4 CSVC trachyandesite 31 19.237 15.607 39.164 19.15 15.60 39.04
14 25_BM-60 CSVC trachyandesite 30 19.509 15.620 39.324 19.41 15.62 39.19
15 CS-30 CSVC monchiquite 30 19.201 15.601 39.187 19.10 15.60 39.04
16 CS-31 CSVC camptonite 30 19.345 15.618 39.216 19.20 15.61 39.05
17 CS-43 CSVC camptonite 30 19.389 15.600 39.190 19.35 15.60 39.14
18 KK1 CHB melilitite 1 19.323 15.605 39.119 19.30 15.60 39.10
19 KK2 CSVC tephrite 25.6 20.178 15.655 39.873 19.97 15.65 39.63
20 KK3 CSVC bazanite 29.7 19.057 15.604 38.908 18.98 15.60 38.79
21 KK4 RPVC bazanite 30 19.542 15.603 39.259 19.41 15.60 39.10
22 KK5 RPVC tephrite 30 20.041 15.625 39.665 19.79 15.61 39.31
23 KK6 RPVC camptonite 28.7 19.493 15.622 39.244 19.35 15.62 39.05
24 KK7A CSVC microsyenite 30 19.161 15.608 39.039 19.06 15.60 38.92
25 KK8 CSVC camptonite 30 19.193 15.603 39.088 19.10 15.60 38.97
26 KK9 RPVC polzenite 70 20.309 15.667 40.152 19.73 15.64 39.37
27 KK10 RPVC melilitite 70 19.756 15.622 39.613 19.36 15.60 39.09
28 KK11 RPVC polzenite 70 20.422 15.644 40.173 19.82 15.61 39.30
29 KK12 RPVC basanite 70 19.988 15.644 39.634 19.74 15.63 39.29
30 ME-3/13 CSVC nephelinite 27 19.720 15.625 39.536 19.62 15.62 39.38
31 ME-4/13 CSVC basanite 30 19.743 15.626 39.481 19.59 15.62 39.27
32 OC-1 RPVC camptonite 30 19.164 15.606 39.067 19.05 15.60 38.91
33 OC-2 RPVC polzenite 70 20.491 15.656 40.210 19.92 15.63 39.56
34 OC-9 RPVC polzenite 68.4 20.849 15.678 41.135 19.97 15.64 39.77
35 OC-10 RPVC aln
oite 70 20.567 15.655 40.233 20.01 15.63 39.55
36 OC-12 RPVC melilitolite 70 19.951 15.638 39.816 19.62 15.62 39.38
TLead isotope data recalculated for their published age using the constants recommended by IUGS (λ
232
Th ¼4.9475E11
y1
,λ
235
U¼9.8485E10
y1
,λ
238
U¼
1.55125E10
y1
), and Pb, U and Th concentrations listed in Table 2.
Ages are taken from Ulrych et al. (1998, 2002, 2013, 2014, 2018), Skala et al. (2014), Ackerman et al. (2015), Dostal et al. (2017).
S. Krmí
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a et al. Geoscience Frontiers 11 (2020) 925942
937
Quaternary mantle-derived rocks, but the older age of the Variscan rocks
(Fig. 9).
Generally, the Pb isotope signature of off-rift volcanic rocks shows
smaller contributions from the asthenospheric mantle than rocks from
the rift centre (cf. Haase and Renno, 2008). For instance, tephrite from
Ríp Hill (sample KK2) taken at the eastern margin of the Eger Rift, ca. 30
km away from the rift centre, has the most radiogenic Pb isotope
composition. Similarly, Tertiary basalts of Lower Silesia that are in an
off-rift position in the possible continuation of the Eger Rift (Fig. 1A)
have radiogenic Pb isotope compositions, overlapping the eld occupied
by ultramac lamprophyres from the Bohemian Massif (Blusztajn and
Hart, 1989). The shift towards higher
207
Pb/
204
Pb observed for some
Lower Silesia basalts (Fig. 7), however, is due to xenoliths of older crustal
rocks derived from the Proterozoic basement of the East European Craton
(Blusztajn and Hart, 1989).
Fig. 8. Nb/Ta vs. Lu/Hf (A), Nb/La vs. (Ce/Yb)
N
(B) and Ba/La vs. (Ce/Yb)
N
(C)
diagrams for Late Cretaceous to Cenozoic undifferentiated magmatic rocks of
the Bohemian Massif. Elevated and/or variable Nb/Ta, Nb/La and (Ce/Yb)
N
along relatively narrow variation in Lu/Hf and Ba/La in Late Cretaceous ultra-
mac lamprophyres reect low-degree partial melting of heterogeneous litho-
spheric mantle source (mantle component A in section 5.4). Tertiary alkaline
lamprophyres and basanite/tephrite volcanic rocks show relatively constant Nb/
Ta and variable Ba/La, which may indicate that their source interacted with
asthenosphere-derived melts (mantle component B in section 5.4). Average OIB
and chondritic Nb/Ta ratios are plotted according to Pf
ander et al. (2012).
Nb/La of average depleted MORB mantle and lithospheric mantle meta-
somatised by subducted continental crust (LMMCC) according to Workman and
Hart (2005) and Borghini et al. (2018), respectively.
Fig. 9. Initial Pb isotope composition of Variscan and Alpine (Cretaceous to
Quaternary) mantle-derived rocks from Central Europe: Bohemian Massif
(Blusztajn and Hart, 1989; Haase and Renno, 2008; Abdelfadil et al., 2013;
Krmí
cek et al., 2016; Ulrych et al., 2016), Germany, i.e., Vogelsberg (Wedepohl
and Baumann, 1999; Bogaard and W
orner, 2003; Jung et al., 2011), Siebenge-
birge (Wedepohl and Baumann, 1999; Jung et al., 2005, 2012; Kolb et al., 2012;
Schneider et al., 2016), Schw
abische Alb (Hegner et al., 1995; Wilson et al.,
2004), Westerwald (Haase et al., 2004), SiebengebirgeWesterwald transition
zone (Schubert et al., 2015), Kaiserstuhl (Schleicher et al., 1991), Rh
on (Meyer
et al., 2002; Jung et al., 2005, 2013; Mayer et al., 2013, 2014), Hegau (Wede-
pohl and Baumann, 1999), Heldburg (Pf
ander et al., 2018), Hocheifel (Fekia-
cova et al., 2007), Schwarzwald (Hegner et al., 1998), Halle (Romer et al.,
2001), Saar-Nahe Basin (Schmidberger and Hegner, 1999), Hessian Depression
(Wedepohl and Baumann, 1999; Wedepohl, 2000), and the
Carpathian-Pannonian region/Pannonian Basin (Salters et al., 1988;
Embey-Isztin et al., 1993a,b; Dobosi et al., 1995; Trua et al., 2006; Harangi and
Lenkey, 2007; Harangi et al., 2007). The in situ growth of Pb from Variscan to
Mesozoic and Cenozoic mantle-derived rocks is shown schematically by a thick
yellow arrow. On the contrary, the shift from more radiogenic Pb isotope
compositions in the Cretaceous ultramac lamprophyres (mantle component A)
to less radiogenic Cenozoic mantle-derived rocks (mantle component B) is
marked by a thin black arrow. Upper crust, orogenic and mantle growth curves
are taken from Zartman and Doe (1981). Some Cenozoic rocks of the Central
European Volcanic Province show conspicuously low
206
Pb/
204
Pb ratios
(<18.75) that broadly overlap with the eld for Variscan samples. This rela-
tively unradiogenic signature, accompanied by slightly enhanced
208
Pb/
204
Pb,
may reect the melting of ancient lower crustal material during magma stag-
nation within continental crust (e.g.,; Hegner et al., 1995; Wedepohl and Bau-
mann, 1999; Wedepohl, 2000; Borisova et al., 2001; Bogaard and W
orner, 2003;
Jung et al., 2011) or incongruent melting of Variscan metasomatised mantle
rocks with clinopyroxene hosting U and Th and having more radiogenic Pb
isotope composition than amphibole that is the major host of Pb (e.g., Wedepohl
and Baumann, 1999; Mayer et al., 2014).
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
938
5.4. Lead isotope evolution of the Central European upper mantle
The lithospheric mantle beneath Variscan Europe, by and large, did
not change its geochemical signature. Instead its Pb isotope signature
evolved by in situ growth, which depends on U/Pb and Th/Pb and time.
During the Alpine orogeny, however, some segments of Variscan Europe
were intensely reworked, and the mantle beneath these segments may
have been modied (i) by subducted crustal material during the Alpine
orogeny or (ii) by signicant input of asthenospheric mantle in exten-
sional zones.
In the
207
Pb/
204
Pb vs.
206
Pb/
204
Pb diagram, mantle-derived rocks of
Central Europe dene a broad eld that in part include rocks that have
higher or lower
207
Pb/
204
Pb at comparable
206
Pb/
204
Pb than mantle-
derived rocks from the Bohemian Massif (Fig. 9). Based on their Pb
isotope compositions, two major mantle components can be distin-
guished. The Pb isotope signature of Late Cretaceous pre-rift ultramac
lamprophyres has the most radiogenic
206
Pb/
204
Pb values, reecting the
lithospheric mantle metasomatised by Variscan subduction (mantle
component A, Fig. 9). This is supported by the chemical composition of
the rocks with strong LREE over HREE enrichment [(high (Ce/Yb)
N
]in
combination with high K
2
O, Nb/Ta, and low TiO
2
and Ba/La and distinct
troughs for K and Pb in primitive mantle normalised trace element plots
(Figs. 5, 6A and 8). The high-radiogenic mantle component A is also
prominent in the source of some off-rift volcanic areas (e.g., Blusztajn and
Hart, 1989; Haase and Renno, 2008). Alpine extension within the
Variscan massifs of Central Europe in the Tertiary eventually resulted in
the Cenozoic rift system with widespread and voluminous volcanism in
the Rhine, Ruhr Valley, and Leine grabens in Germany, and in the Eger
Rift in the Bohemian Massif (Ziegler, 1992). During these events,
Variscan metasomatised lithospheric mantle (component A) below the
rifts was refertilised by upwelling material from the asthenosphere
(Dostal et al., 2019a), which is reected by rocks with less radiogenic
206
Pb/
204
Pb values (Fig. 9). This mixed mantle (mantle component B) is
the major source for the majority of rift-related Tertiary and Quaternary
volcanic rocks of Central Europe, e.g., Vogelsberg, Siebengebirge,
Schw
abische Alb, Westerwald, Kaiserstuhl, Rh
on, Hegau, Heldburg as
well as for volcanic rocks from the Eger Rift.
Mantle-derived rocks from the Carpathian-Pannonian region have
207
Pb/
204
Pb values that resemble those of Siebengebirge, Kaiserstuhl,
Westerwald, SiebengebirgeWesterwald transition zone and Rh
on, and
are slightly higher than those of corresponding rocks from the Bohemian
Massif (Fig. 9). Additionally, several samples from the Carpathian-
Pannonian area are characterised by
207
Pb/
204
Pb values as high as
15.75 (Trua et al., 2006; Harangi and Lenkey, 2007). These high values
reect variable contributions from old crust.
Most volcanic rocks of CEVP fall on a two-component mixing trend in
143
Nd/
144
Nd vs.
206
Pb/
204
Pb diagram (Fig. 10). Although the Nd isotope
values more or less positively correlate with the Pb isotope signature, the
subtle differences in composition of the upper mantle sampled by
Mesozoic and Cenozoic rocks are not so clearly documented by Nd iso-
topes as by Pb isotopes. Dostal et al. (2019a) demonstrated the evolution
of Nd isotope composition of the upper mantle beneath Central and
Western Europe from very unradiogenic values induced by the Variscan
orogeny to more radiogenic values that developed during the Mesozoic
rifting of the Atlantic Basin. The Nd isotope composition of Central Eu-
ropean upper mantle was not further signicantly affected since ca. 100
Ma, which is supported by studied samples. The Mesozoic to Cenozoic
mantle-derived rocks are distinctly radiogenic (Ulrych et al., 2008; Sk
ala
et al., 2014; Dostal et al., 2017) compared to those of the Variscan age
(Krmí
cek et al., 2016).
Studied mantle-derived rocks are neither overlapping, nor trending to
the FOZO mantle component (Fig. 10) that isotopically overlaps with a
low-velocity component (cf. Hoernle et al., 1995). This excludes an
involvement of the mantle plume in their magma evolution, which is in
line with seismic data from the mantle beneath the Bohemian Massif
(Plomerov
a et al., 2007; Hrubcov
a et al., 2017).
6. Conclusions
Based on our Pb isotope study of mantle-derived rocks from the Bo-
hemian Massif and their comparison with other volcanic provinces in
Central Europe, we make the following conclusions:
(1) Upper Palaeozoic mantle-derived rocks demonstrate that the
lithospheric mantle beneath the Bohemian Massif was meta-
somatised during the Variscan orogeny and received its Pb iso-
topic signature from subducted crustal material. Cretaceous,
Tertiary and Quaternary mantle-derived rocks of the Bohemian
Massif have been extracted from this mantle and, therefore, share
the same crustal signature.
(2) The Pb isotope data dene a two-component mixing trend. The
Cretaceous ultramac lamprophyres represent a high radiogenic
end-member characterised by
206
Pb/
204
Pb ratios up to 20.0
(Variscan metasomatised lithospheric mantle), whereas the Ter-
tiary alkaline lamprophyres and Cretaceous and Quaternary
melilitites originated from the mantle with
206
Pb/
204
Pb ratios
below 19.4 (lithospheric mantle substantially modied via inter-
action with the convective mantle). The Pb isotope composition of
the Tertiary volcanic samples falls between these two
components.
Fig. 10. Initial
143
Nd/
144
Nd vs.
206
Pb/
204
Pb diagram of investigated samples
compared to major volcanic provinces of Central European Volcanic Province
(CEVP). Studied samples overlap with most of the CEVP mantle-derived rocks
except for the Heldburg dykes that have less radiogenic
206
Pb/
204
Pb (see
Pf
ander et al., 2018 for interpretation) and mac rocks from the Lower Silesian
Basin with more radiogenic
143
Nd/
144
Nd (Blusztajn and Hart, 1989). Late
Cretaceous ultramac lamprophyres show more radiogenic NdPb signature
(end-member A) trending towards high radiogenic HIMU-like mantle compo-
nent, whereas Tertiary alkaline lamprophyres have less radiogenic signature
(end-member B) trending towards OIB-like (non-HIMU type) mantle compo-
nent. Note, the mixing trend between A and B is dened by contributions from
two types of Pb-enriched lithospheric mantle. Deviations from this trend toward
higher
143
Nd/
144
Nd ratios may reect the presence of material from the
depleted mantle that has low Pb contents (and therefore does not affect the Pb
isotope composition of these rocks) and sufciently high Nd contents to shift the
Nd isotope composition to more radiogenic
143
Nd/
144
Nd ratios. Published
143
Nd/
144
Nd ratios for samples from the same localities as samples studied in
this paper are adopted from Ulrych et al. (2008), Skala et al. (2014) and Dostal
et al. (2017). Compositional elds for other lavas from the CEVP are taken from
Pf
ander et al. (2018). Mantle components and trends (OIB, HIMU, FOZO)
correspond to Stracke et al. (2005).
S. Krmí
ckov
a et al. Geoscience Frontiers 11 (2020) 925942
939
(3) Generally, off-rift volcanic/subvolcanic rocks are derived from a
mantle source with higher
206
Pb/
204
Pb than corresponding rocks
from axial parts of the rift, possibly indicating that the low
206
Pb/
204
Pb component is derived from the mantle source inu-
enced by asthenosphere upwelling.
(4) The majority of Cenozoic mantle-derived rocks of Central Europe
show similar Pb isotope variations as those of the Bohemian
Massif.
Acknowledgments
This research was nancially supported by the institutional project
RVO 67985831 of the Institute of Geology of the Czech Academy of
Sciences, as well as by the Brno University of Technology project LO1408
AdMaS UP Advanced Materials, Structures and Technologies, sup-
ported by the Ministry of Education, Youth and Sports of the Czech Re-
public under the National Sustainability Programme I. S.K., L.K. and
J.U. thank Jaroslav Dostal (Saint Marys University, Canada) and Martin
J. Timmerman (University of Potsdam, Germany) for discussion. The
authors greatly appreciate an anonymous reviewer and Dr. C. Spencer for
their very constructive and helpful comments and suggestions.
Appendix A. Supplementary data
Supplementary data to this article can be found online at https://doi.
org/10.1016/j.gsf.2019.09.009.
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