JOURNAL OF
PETROLOGY
Journal of Petrology, 2015, Vol. 56, No. 3, 527–562
doi: 10.1093/petrology/egv008
Advance Access Publication Date: 29 March 2015
Original Article
Magmatic Response to Slab Tearing:
Constraints from the Afyon Alkaline Volcanic
Complex, Western Turkey
Dejan Prelević1,2*, Cüneyt Akal3, Rolf L. Romer4, Regina Mertz-Kraus1
and Cahit Helvacı3
1
Institute for Geosciences, University of Mainz, Becherweg 21, D-55099 Mainz, Germany, 2Faculty of Mining and
Geology, University of Belgrade, Djušina 7, 11000 Belgrade, Serbia, 3Dokuz Eylül Üniversitesi, Mühendislik
Fakültesi, Jeoloji Mühendisliği Bölümü, TR-35397 Buca, Izmir, Turkey and 4Helmholtz-Zentrum Potsdam Deutsches
GeoForschungsZentrum GFZ, Telegrafenberg, D-14473 Potsdam, Germany
*Corresponding author. E-mail: [email protected]
Received July 16, 2014; Accepted February 9, 2015
ABSTRACT
The Middle Miocene Afyon alkaline volcanic complex (western Anatolia) erupted lavas of highly
variable geochemistry, ranging from silica-undersaturated to silica-oversaturated and from ultrapotassic to Na-alkaline compositions. There are two major volcanic groups showing substantial
differences in K-enrichment and different Sr, Nd and Pb isotopic compositions: plagioclase–
amphibole-bearing lavas and sanidine- and/or leucite-bearing lavas. The most remarkable feature
of Afyon volcanism is the close relationship in time and space of these two lava types. There
is clear stratigraphic evidence for a switch from early Si-oversaturated sanidine- and/or leucitebearing lavas, towards Si-undersaturated sanidine- and/or leucite-bearing lavas, which eventually
change to slightly Si-undersaturated to -saturated plagioclase–amphibole-bearing lavas that make
up the youngest formations. This change in composition is coupled with a decrease in 87Sr/86Sr
(whole-rock and in situ apatite, perovskite, melilite and clinopyroxene), 207Pb/204Pb, Zr/Nb and Th/
Nb, and an increase in 143Nd/144Nd, 206Pb/204Pb, 208Pb/204Pb and Ce/Pb, thus delineating a systematic change from orogenic (crust-like) to anorogenic (within-plate) signatures. Magma genesis in
the Afyon volcanic complex has been controlled by roll-back of a subducted lithospheric slab since
the Early Tertiary and post-collisional extensional events in Miocene times. It is associated with the
upwelling of asthenospheric mantle through a gap in the subducted slab under western Anatolia.
Magmatism is concurrent with the collapse of the orogenic belt and the development of extensionrelated horst and graben structures. We interpret the geochemical transition from orogenic to
anorogenic affinity as being due to the increasing role of lithosphere–asthenosphere interaction
that is most strongly reflected in the geochemistry of the Afyon lavas. Melting of peridotite in the
convecting mantle (asthenosphere) may be a viable model for the origin of the plagioclase–amphibole-bearing lavas. Their ubiquitous high K2O contents, orogenic trace element signatures and isotopic compositions imply that the asthenosphere-derived primary melts were contaminated by
melts derived from lithospheric mantle containing an orogenic chemical signature. Conversely, the
ultrapotassic sanidine- and/or leucite-bearing lavas are derived from at least two types of metasomatized lithospheric mantle. The dominant source is a phlogopite–pyroxene-rich metasome,
which was generated by recycling of continental sediments during previous subduction episodes.
This is responsible for the orogenic geochemical signature dominantly seen in lamproites
and shoshonites. On the other hand, melting of recently generated phlogopite-wehrlite metasomes
resulted in the parental melts of melilite-leucitites, which should be of proto-kamafugitic composition. The wehrlitic metasomes were generated when convecting mantle-derived precursor
C The Author 2015. Published by Oxford University Press. All rights reserved. For Permissions, please e-mail: [email protected]
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melts reacted with lithospheric mantle peridotite along the solidus ledge in the system
lherzolite þ CO2 (<22 kbar).
Key words: active continental margin; alkali basalt; melilitite; asthenosphere; clinopyroxene; continental lithosphere; leucite; geothermobarometry; lamproite; laser ablation
INTRODUCTION
Tearing of the subducting slab occurs in many orogenic
regions worldwide (Davaille & Lees, 2004; Ferrari, 2004;
Dimalanta & Yumul, 2008). It was first reported on the
basis of seismic data as a major gap or hole in the slab of
lithosphere plunging beneath the New Hebrides island
arc (Chatelain et al., 1992). More recently, a number of
tomographic models have indicated segmentation and
disruption of oceanic slabs in several regions of the
Mediterranean, characterized by slow upper mantle
P-wave velocities (Mason et al., 1998; Maury et al., 2000;
Rosenbaum & Lister, 2004; Rosenbaum et al., 2008;
Pérez-Valera et al., 2013). Segmentation of subduction
zones through slab tearing is associated either with rollback of subducting slabs or with the arrival of nonsubductible continental lithosphere, oceanic plateaux or
seamounts at the trench (Rosenbaum & Mo, 2011).
A direct response to the disruption of the subducting
slab is the incursion of fresh asthenospheric mantle that
can initiate magmatism of different chemistry. The
magmatism may originate in both the convecting mantle and overlying recently metasomatized mantle lithosphere, resulting in a general transition of the chemical
signature from arc type to ocean island basalt (OIB)
type with change of the magma source. It is, however,
unclear how specific magmatic rock types relate to particular aspects of slab disruption. In the traditional view,
the resulting magmatism will show a general transition
in its chemical signature from arc type to OIB type.
Lithospheric mantle heterogeneity combined both with
variable depth and degree of partial melting combined
with intracrustal differentiation of the primary magmas,
may, however, generate a wide range of magma
compositions that could deviate from this general transitional pattern. Preconditioning of the mantle lithosphere needs to be given special attention because of
the ‘memory’ effect of this source component and its
ability to keep the typical chemical fingerprint of a magmatic arc [high K and large ion lithophile element (LILE)/
high field strength element (HFSE) ‘spiky’ trace element
patterns] in heterogeneous ‘metasomes’ that formed
during preceding subduction events by infiltration of
fluids and melts liberated from the subducting oceanic
slab, including continent-derived sediments on top of
this slab (e.g. Prelević et al., 2013 and references
therein). Moreover, a later generation of metasomes
related to infiltration of partial melts of asthenospheric
mantle also may be present in the lithosphere. The
chemical composition of the resulting magmas generated by partial melting of the lithospheric mantle may
represent a blend of all these ingredients.
Circum-Mediterranean Cenozoic magmatism (sensu
Lustrino & Wilson, 2007) is represented by widespread
K- and Na-alkaline basaltic and more evolved lavas,
developed within European continental and Mediterranean regions. The magmatism is spatially and temporally associated with the Late Cretaceous–Cenozoic
convergence of Africa–Arabia with Eurasia that resulted
in the progressive closure of oceanic basins and ultimately in the formation of the Alpine collisional orogen at
the southern passive continental margin of Europe.
This young volcanism has been subdivided into two
chemically distinct groups, referred to collectively as
‘anorogenic’ and ‘orogenic’ (Wilson & Bianchini, 1999;
Lustrino & Wilson, 2007). The ‘anorogenic’ magmatism
is generally characterized by basalts with chemical
signatures similar to the sodic alkaline basalts (withinplate) typical for oceanic and continental plates worldwide. In contrast, the chemical signature of ‘orogenic’
magmatism has all the characteristics typical of arc volcanism, potassium enrichment being the most distinctive feature. Previous isotope studies imply that the
mantle source(s) of the Circum-Mediterranean Cenozoic
magmatism had been invaded by metasomatic agents.
The location of these reservoirs, their ultimate origin
and the trigger for magmatism, however, are still a matter of debate (Lustrino & Wilson, 2007; Lustrino et al.,
2011). In a highly simplified view, the anorogenic magmatism is assumed to be derived dominantly from the
convecting mantle with contributions from plume material, whereas the orogenic magmatism is derived
from the variably metasomatized lithospheric mantle or
supra-subduction zone mantle wedge. Traditionally, the
two types of magmatism have been treated separately
and only recently was it recognized that in several Mediterranean volcanic districts, where tomographic models
indicate gaps within the subducting slab, geochemical
transitions and/or interfingering of the two types occur,
as in Spain (Duggen et al., 2005; Prelević et al., 2008),
the Alps (Davies & von Blanckenburg, 1995), Italy (Gasperini et al., 2002; Rosenbaum et al., 2008; Conticelli
et al., 2013) and western Anatolia in Turkey (Prelević
et al., 2012).
The western Anatolian region together with the
entire Aegean region is considered as an arc–back-arc
region, dominated by slab roll-back (van Hinsbergen
et al., 2005; Jolivet & Brun, 2010) and extensive postMiocene extension that took place after thickening of
the lithosphere during a series of collisional events during the Alpine orogeny from the late Cretaceous to the
early Tertiary (S
engör et al., 1985; Yılmaz et al., 2000;
Rimmelé et al., 2003; Ring et al., 2003; Isık et al., 2004;
Journal of Petrology, 2015, Vol. 56, No. 3
Çemen et al., 2006; Westaway, 2006; Glodny & Hetzel,
2007). A thermal anomaly originating from a tear, identified by seismic tomography, in the subducted slab(s)
under western Anatolia (Fig. 1) is likely to play an additional role as a trigger for this magmatism and strongly
influences its chemical signature (Dilek & Altunkaynak,
2009; Prelević et al., 2012; Karaoğlu & Helvacı, 2014).
The chemistry of the western Anatolian volcanic rocks
shows a regional trend of southward increasing Siundersaturation, with the first occurrence of highly alkaline undersaturated volcanism in the Afyon alkaline
volcanic complex at 10–12 Ma (Akal et al., 2013). This regional change is coupled with a transition in a number
of chemical parameters that delineate a systematic
change from dominantly orogenic (crust-like) to increasingly anorogenic (convecting mantle-like) signatures (Francalanci et al., 2000; Çoban & Flower, 2007;
Dilek & Altunkaynak, 2009; Çoban et al., 2012; Prelević
et al., 2012; Semiz et al., 2012). Therefore, western
Anatolia—and more specifically the Afyon alkaline volcanic complex—represents a perfect natural laboratory
in which to study the response of the orogenic lithosphere to the specific incursion of melts derived from
the convecting mantle.
In this study we report a comprehensive set of
whole-rock and mineral chemical data for the lavas and
entrained cumulitic enclaves from the Afyon alkaline
volcanic complex. Combining major and trace element
and Sr, Nd and Pb isotope data for whole-rock samples
with in situ major and trace element and Sr isotope data
for minerals, we are able to trace the complex interplay
between the thick lithosphere in the Afyon region and a
mantle input that has a progressively higher contribution from the asthenosphere.
529
developed after the early–middle Miocene initiation of
oceanic lithosphere subduction (Meulenkamp et al.,
1988; Garfunkel, 1998). However, partly coeval OligoMiocene to early Pliocene dominantly mantle-derived
volcanism in western Anatolia farther east (Fig. 1) is interpreted not to be directly related to this active subduction (Prelević & Seghedi, 2013 and references therein).
There is a systematic younging trend from north to
south (Fig. 1): the oldest magmatic rocks occur in the
northernmost part of western Anatolia and yield
Eocene and Oligo-Miocene ages (i.e. they postdate the
Eocene continental collision); farther to the south, a
change from regional compression to widespread extension occurred in the Middle Miocene (Yılmaz, 1989;
Yılmaz et al., 2000, 2001) and alkaline volcanism
occurred at this time. During this period, most of the
Menderes Massif was exhumed, resulting in the development of horst–graben structures (Yılmaz et al., 2000;
Westaway, 2006; Glodny & Hetzel, 2007). Magmatism
follows in space and time the southward lithospheric
slab roll-back (Jolivet & Brun, 2010), and is controlled
by a combination of three major processes: (1) post-collisional extension initiated after major lithospheric
thickening within the Menderes Massif (Akal, 2003;
Prelević et al., 2012; Ersoy & Palmer, 2013; Seghedi
et al., 2013); (2) subsequent collapse of the orogenic
belt (Gessner et al., 2013); coupled with (3) the initiation
and migration of a thermal anomaly (Prelević et al.,
2012) originating from a gap in the subducted slab(s)
under western Anatolia (Spakman et al., 1988; Wortel &
Spakman, 1992, 2000; Piromallo & Morelli, 2003; Biryol
et al., 2011). The relative significance of these processes, however, remains a matter of continuing debate.
GENERAL GEOLOGICAL BACKGROUND
THE AFYON VOLCANIC COMPLEX AND ITS
LOWER MIOCENE STRATIGRAPHY
The Aegean region and western Anatolia represent an
arc and back-arc basin system in which at least 2400 km
of continental or oceanic lithosphere has been underthrust or subducted (Faccenna et al., 2003) since the
Middle Jurassic, interpreted as either a single continuous subduction zone (van Hinsbergen et al., 2005) or
multiple successive and diachronous subduction systems (Dercourt et al., 1986). Tomographic studies indicate the presence of two subducted slabs beneath
Anatolia: the South Aegean slab in the west (Hellenic
arc in Fig. 1) and the Cyprus slab in the east (Amaru,
2007; Özacar et al., 2010). In the single-subduction zone
model, the present Aegean subduction system formed
after roll-back of the South Neotethyian slab at 65 Ma
(Jolivet & Brun, 2010, and references therein).
Tomographic models indicate that segments of the
Aegean and Cyprus slabs are separated by a left-lateral
tear (Özacar et al., 2010) or a slab window (Amaru,
2007), which is characterized by slow upper mantle
P-wave velocities. The lateral extent of this window
spans nearly the entire width of western Anatolia
(Fig. 1). The active South Aegean volcanic arc (Fig. 1)
Afyon volcanism (Fig. 2a) is traditionally considered as
part of the southward-trending Kırka–Afyon–Isparta volcanic zone (Fig. 1). It occurs between the Köroğlu caldera (Figs 1 and 2a) that collapsed at 18–16 Ma (Aydar
et al., 1998; Prelević et al., 2012) in the north and the
Isparta–Gölcük volcanic area in the south, where volcanism began with lamproites at around 4 Ma and terminated with trachytic rocks at the Gölcük volcano at
around 10 ka (Lefevre et al., 1983; Floyd et al., 1998;
Prelević et al., 2012). The volcanic successions of the
Afyon volcanic complex were deposited on strongly deformed sedimentary formations of the Western Tauride
belt, which are dominantly represented by carbonate
platform sequences. Two major volcanic systems occur
in the studied area to the south of the city of Afyon
(Fig. 2a, b): the ‘early stage formations’ (Akal et al.,
2013) representing an older volcanic system of domes,
lava flows and volcaniclastic successions with trachytic
compositions are hereafter referred to as the ‘Early
Afyon volcanic formation’; the ‘late stage formations’
(Akal et al., 2013) that partially overlie the older volcanic
system to the south include volcanoclastic successions
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Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 1. (a) Simplified tectonic map of the eastern Mediterranean. (b) Map of western Anatolia (Turkey) modified from Prelević et al.
(2012) and references therein, showing the major volcanic fields (pink) and their ages. Distribution of the volcanic rocks is modified
from the MTA 1:500 000 scale geological map of Turkey, combined with published age data (Borsi et al., 1972; Besang et al., 1977;
Pis
kin, 1980; Paton, 1992a, 1992b; Ercan et al., 1996; Seyitoglu et al., 1997; Aldanmaz et al., 2000; Helvacı & Alonso, 2000; Robert &
Montigny, 2001; Erkül et al., 2005; Innocenti et al., 2005; Westaway et al., 2005; Ersoy et al., 2008; Helvacı et al., 2009; Prelević et al.,
2012). (c) Geodynamic model for the post-collisional Tertiary tectonic development of southwestern Anatolia. The distribution of
Neogene to Quaternary volcanic rocks within western Anatolia is superimposed on a tomographic model showing the tearing of the
Aegean and Cyprus slabs (Biryol et al., 2011; Prelević et al., 2012; Akal et al., 2013). The red rectangle delineates the area shown in
Fig. 2a. The black rectangle delineates the profile presented in Fig. 19. The three-dimensional schematic illustration shows the actual
position of the tear through which the counterflow of convective mantle is channelled into the lithospheric mantle (Faccenna & Becker,
2010). Hot asthenosphere upwelling triggers volcanism by melting of either the lithosphere because of the thermal instability of phlogopite or the adiabatically decompressed asthenosphere. Tectonic zones of western Anatolia according to Okay & Tüysüz (1999).
(continued)
Journal of Petrology, 2015, Vol. 56, No. 3
531
Fig. 1. Continued.
and late-stage lavas, dykes and domes, and will be
referred to below as the ‘Late Afyon volcanic complex’.
The older volcanism of the Early Afyon volcanic formation produced Si-oversaturated lavas with relatively
similar petrology, whereas the magmas of the younger
Late Afyon volcanic complex, which erupted through a
spatially overlapping magma plumbing system, were
both Si-oversaturated and Si-undersaturated, variably
potassic and ultrapotassic types with very different petrological characteristics. In this study, we concentrate
on the products of the Late Afyon volcanic complex.
The volcanic and volcaniclastic products of the Late
Afyon volcanic complex cover 200 km2 in the studied
area (Fig. 2c). The volcanic rocks exhibit a wide range of
compositions and volcanic facies. As a detailed description of the stratigraphy and volcanic successions of the
volcanic complex was published previously (Akal et al.,
2013) they are only briefly summarized below.
Three episodes of volcanic activity are observed in
the Late Afyon volcanic complex (Fig. 2b and c).
Lamproites, shoshonites, (melilite-) leucitites and tephriphonolites and volumetrically small trachyandesite
lava flows are the oldest products of volcanic activity.
They cover and intrude the sedimentary rocks of the
Western Tauride Belt and the volcanoclastic products of
the Early Afyon volcanic formation. Products belonging
to the second phase of volcanic activity are lamproites,
trachyandesitic lavas and associated pyroclastic rocks.
Lamproites occur as lava flows and domes. The thick
and widespread trachyandesitic pyroclastic succession
is composed of multistage pyroclastic fall and flow deposits, and lava blocks. Several variably sized trachyandesitic feeders in the form of lava domes, plugs and
subvolcanic stocks, and laterally discontinuous stubby
lava flows, cut the pyroclastic succession. Lacustrine
sedimentary rocks, consisting of limestone, claystone,
sandstone and pebblestone alternations, indicate an
interval of unknown duration between the second and
the third volcanic phase. The third episode represents
the youngest volcanic activity in the study area and is
characterized by phonotephritic, phonolitic, basaltic trachyandesitic, and nosean-bearing trachyandesitic lava
domes, dykes and lava flows.
ANALYTICAL METHODS
Our sample collection comprises more than 100 samples of lavas and xenoliths from the Afyon alkaline volcanic sequence. We selected 67 samples for whole-rock
major and trace element analyses by X-ray fluorescence
(XRF) and inductively coupled plasma mass spectrometry (ICP-MS), and 20 representative samples for Sr, Nd
532
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 2. (a) Geological map of the region south of Afyon city. (b) Generalized stratigraphic section of the southern part of the Late Afyon
volcanic complex, showing stratigraphic correlations in the volcanic succession, which overlies the sedimentary formations of the
Tauride Belt. (c) Detailed geological map of the study area showing the main rock types of the Late Afyon volcanic complex.
Representative stratigraphic sections are indicated. Map is modified from the MTA 1:500 000 scale geological map of Turkey.
(continued)
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 2. Continued.
533
534
and Pb isotope analyses by thermal ionization mass
spectrometry (TIMS). The full dataset is given in
Supplementary Data File 1 (supplementary data are
available for downloading at http://www.petrology.
oxfordjournals.org).
Whole-rock major elements and Pb, Ni and Cr were
determined by XRF using a Philips MagiXPRO spectrometer on fused discs and pressed pellets at the
University of Mainz. The rest of the trace elements were
analyzed by laser ablation (LA)-ICP-MS using an Agilent
7500ce ICP-MS system coupled with a New Wave
UP-213 LA system at the University of Mainz. For that
purpose, rock powders were melted to form homogeneous glass beads without any fluxing agent on an
iridium strip heater in an argon atmosphere (Stoll et al.
2008). The glass beads were subsequently analyzed by
LA-ICP-MS. Details on measurement conditions, accuracy and reproducibility of the analyses have been reported by Prelević et al. (2012).
Isotope compositions were determined on a split of
the powders used for major and trace element analyses.
Samples were digested in concentrated HF in Savillex
beakers on a hotplate for 4 days and then evaporated to
near dryness. The samples were taken up in 2 N HNO3 to
convert fluorides to nitrates and slowly dried again. The
samples were then redissolved in 6 N HCl and once a
clear solution was obtained it was aliquoted for Sr–Nd
and Pb isotope analysis. Strontium, Nd, and Pb were separated using standard procedures (e.g. Romer et al.,
2001, 2005, and references therein). Special care was
taken to remove Ba from the rare earth element (REE)
fraction (using an additional washing step with 25 N
HNO3) to avoid interferences of BaO on mass 146Nd.
Strontium and Nd isotope compositions were measured
by TIMS on a Triton system operated in dynamic multicollection mode, using Ta single filaments and Re double
filaments, respectively, at Helmholtz-Zentrum Potsdam
Deutsches GeoForschungsZentrum GFZ (GFZ Potsdam,
Germany). Strontium and Nd isotope ratios were normalized to 86Sr/88Sr ¼ 01194 and 146Nd/144Nd ¼ 07219, respectively. During the measurement period, NBS-987 Sr
reference material and the LaJolla Nd standard gave
average 87Sr/86Sr and 143Nd/144Nd values of
0710249 6 12 (2SD of 20 measurements) and
0511850 6 7 (2SD of 11 measurements), respectively.
The Pb isotopic composition was measured by TIMS on a
Triton system on single Re filaments using static multicollection (GFZ Potsdam, Germany). Instrumental fractionation was corrected by 01% per a.m.u. as determined
from the long-term reproducibility of Pb reference material NBS-981. The accuracy and precision of the reported
Pb isotope ratios is better than 01% at the 2r level of uncertainty. The initial Sr, Nd and Pb isotopic compositions
of the Afyon lavas were calculated for the published ages,
using the following decay constants: 87Rb 142 10–11 a–1
(Steiger & Jager, 1977); 147Sm 654 10–12 a–1 (Lugmair &
Marti, 1978); 232Th 49475 10–11 a–1; 235U 9848 10–10 a–1; 238U 155125 10–10 a–1 (Jaffey et al., 1971;
Steiger & Jager, 1977).
Journal of Petrology, 2015, Vol. 56, No. 3
Mineral compositions (including major and trace
elements) were analysed by electron microprobe and
LA-ICP-MS. Clinopyroxene, apatite, melilite and perovskite in selected samples of cumulate xenoliths were
analysed in situ for their 87Sr/86Sr ratios. The data are
given in Supplementary Data Files 2 and 3.
The major element compositions of all minerals
were determined by electron microprobe (JEOL
JXA 8900RL) at the University of Mainz, and with the
MS-46 CAMECA electron microprobe at the Université
Pierre et Marie Curie (Laboratoire de Pétrologie
Minéralogique), Paris. Operating conditions were 20 kV
accelerating voltage, 20 nA beam current and 2 mm
beam diameter. Synthetic and natural minerals were
used for standards.
Trace element analyses of clinopyroxene, phlogopite, apatite, melanite, melilite and perovskite were performed by LA-ICP-MS at the University of Mainz using
an ArF excimer laser (193 nm wavelength, NWR193
system by esi/NewWave) coupled to an Agilent 7500ce
ICP-MS system. The laser was operated at a repetition
rate of 10 Hz with laser energy at the sample site of
5–6 J cm–2, allowing data collection from single grains
in polished thick sections (up to 80 mm thickness) for at
least 40 s. Helium was used as carrier gas with a flow
rate of 07 l min–1. Analyses were carried out with spot
diameters of 50 and 80 mm. Backgrounds were measured for 20 s prior to each ablation and 29Si (clinopyroxene, phlogopite, melanite and melilite) and 43Ca
(perovskite and apatite) were used as the internal standards, applying the Si and Ca concentrations previously
determined by electron microprobe. For calibration
NIST SRM 612 was analysed at the beginning and after
every 30 measurements on the unknown samples. The
time-resolved signal was processed using the program
GLITTER 4.4.1 (www.glitter-gemoc.com; Macquarie
University, Sydney, Australia), applying the preferred
values for NIST SRM 612 reported in the GeoReM database (http://georem.mpch-mainz.gwdg.de/) (Jochum
et al., 2005, 2011) as the ‘true’ concentrations to calculate the element concentrations in the samples.
Analytical uncertainty (1r) for one spot analysis or line
scan was less than 10%. During each run the basaltic
USGS BCR-2 G reference glass and RP91-17 clinopyroxene crystals (Mason et al., 1999) were analysed as unknowns to monitor the accuracy and reproducibility of
the analyses (Supplementary Data File 3).
Laser ablation 87Sr/86Sr in situ analyses of clinopyroxene, apatite, melilite and perovskite crystals were obtained using an Nd:YAG UP213 nm laser system (New
Wave Research) coupled to a NuPlasma MC-ICP-MS
system at the University of Alberta, Radiogenic Isotope
Facility. The analytical protocol for in situ Sr isotope
analyses follows the method outlined by Schmidberger
et al. (2003). The minerals were ablated in a He atmosphere (10 l min–1) using the following parameters: 60 s
ablation time; 160 mm spot size and raster; 20 Hz repetition rate; 15 J cm–2 energy density. The sample-out line
from the laser ablation cell was ‘y’-connected to the
SiO2
TiO2
Al2O3
Fe2O3T
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Sum
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
Rb
Sr
Y
Zr
Nb
Cs
Ba
La
Ce
Pr
Nd
Sm
Rock
type:
Sample:
Nosean phonolite
436
11
154
7
01
27
102
22
97
04
6
985
6
194
3
29
19
101
106
21
535
2950
24
733
44
n.a
7760
133
223
25
94
15
28a*
Phonolite
564
11
199
58
01
23
51
15
61
05
11
998
14
200
15
38
22
32
131
120
250
3250
42
961
48
71
5830
134
250
28
104
17
66b
Latite
578
1
162
59
01
16
36
2
96
03
13
992
89
108
07
23
17
17
51
98
233
2520
26
595
34
42
4860
745
159
17
64
11
97
Latite
462
1
18
65
01
16
98
2
112
03
26
99
82
152
14
35
41
46
53
111
294
3080
23
517
39
10
4580
103
183
21
77
12
50*
Latite
456
13
165
71
01
29
94
18
95
05
4
985
65
174
37
50
12
72
104
162
492
2150
27
633
38
13
8820
122
208
23
85
14
60
Phonotephrite
528
09
168
59
01
38
65
34
54
05
21
982
13
131
56
33
38
30
51
56
114
1680
23
446
35
52
2370
125
213
22
80
12
29
Phonotephrite
589
01
211
21
01
02
12
97
57
0
12
1002
22
15
51
07
12
35
90
29
339
58
79
1018
57
23
28
137
168
12
26
2
27-c*
Shoshonite
606
11
218
43
0
08
29
31
39
06
09
100
23
204
200
53
33
36
54
71
240
2090
23
190
41
21
2900
105
186
20
81
14
84
Latite
547
1
178
57
01
34
67
47
32
05
17
995
13
117
44
24
33
18
37
53
417
2150
21
338
30
30
2200
125
212
22
80
12
28b
Phonolite
487
1
146
71
01
63
109
29
4
07
13
975
19
155
203
39
48
32
33
53
203
2850
26
309
29
34
2390
174
305
34
125
18
31*
Shoshonite
484
07
208
59
01
71
105
15
37
05
04
996
24
177
167
67
40
50
82
72
232
3650
36
575
59
86
2830
244
397
41
147
22
19-a*
Mel.leucitite
562
15
147
73
01
38
58
27
66
1
05
100
26
179
312
34
79
38
44
39
190
1050
23
460
27
41
1500
81
151
19
73
12
1-a*
Mel.leucitite
521
11
202
57
01
57
78
29
35
05
05
999
25
161
90
38
43
26
47
53
148
1780
35
275
38
85
1880
139
237
25
96
15
23-a*
Tephriphonolite
593
471
01
13
212
22
21
67
01
01
01
6
09
8
93
29
58
43
0
11
11
06
100
999
23
19
11
188
047
21
33
42
04
35
3
26
72
79
32
72
374
207
31
3780
74
49
983
330
57
100
27
96
563 2830
146
331
176
583
13
64
28
236
21
35
42
Tephriphonolite
563
07
197
46
01
16
51
28
63
03
21
997
48
100
55
89
29
68
59
56
132
2510
24
393
42
22
2330
153
241
25
86
12
06 AF03
Leucitite
53a*
†
40C
EAVF
533
1
166
78
01
5
78
29
4
07
07
999
21
145
133
38
39
32
46
41
128
1480
23
238
22
68
1620
87
158
18
67
10
(continued)
497
54
1
13
9
109
67
65
01
02
77
5
101
69
14
14
72
83
14
17
2
59
1021
964
25
20
163
115
170
617
n.a.
43
51
291
n.a.
50
n.a.
75
n.a.
12
201
200
1250
1350
26
22
449
325
29
22
n.a
n.a
4150
3580
64
29
119
51
27
61
65
25
21
52
05 BH
01
Mel.leucitite
549
13
109
62
01
51
72
17
82
18
18
991
27
163
170
n.a.
41
n.a.
n.a.
n.a.
173
1360
28
469
29
n.a
3630
48
93
12
51
98
05 IL
01†
Lamproite
465
08
18
59
01
21
69
18
123
02
4
986
4
139
3
15
8
65
90
22
605
2190
17
511
37
n.a
7410
86
149
17
66
11
05 BH
05†
05 IL
02†
Lamproite
SLB lavas
Lamproite
PAB lavas
Latite
Table 1: Selected whole-rock major (wt %) and trace element (mg g–1) analyses for Afyon lavas
Journal of Petrology, 2015, Vol. 56, No. 3
535
43
13
16
84
15
41
06
39
06
26
32
65
22
138
36
10
12
64
112
29
042
27
037
17
27
65
21
139
Nosean phonolite
3
81
1
57
1
27
037
23
035
16
27
41
23
87
Phonolite
32
85
094
5
086
22
032
2
03
12
21
57
19
30
66b
Latite
34
91
11
57
099
25
036
22
032
15
23
56
17
121
97
Latite
29
77
089
45
081
22
032
21
031
95
21
41
11
54
50*
Latite
02
12
013
077
016
061
014
12
024
18
12
136
28
71
60
Phonotephrite
34
95
11
56
093
23
028
19
026
65
29
37
15
51
29
Phonotephrite
29
79
087
44
078
2
028
18
026
79
18
37
11
37
27-c*
Shoshonite
43
11
11
55
092
23
03
18
027
75
18
35
78
10
84
Latite
53
13
14
71
122
33
048
34
051
13
33
79
25
37
28b
Phonolite
27
8
094
49
086
22
031
21
03
12
18
26
10
16
31*
Shoshonite
35
10
12
68
122
32
046
31
044
75
27
47
13
24
19-a*
Mel.leucitite
83
22
24
1107
182
43
053
32
046
93
64
74
21
49
1-a*
Mel.leucitite
28
025
73
11
085
012
44
077
081
015
22
06
033
012
23
11
033
022
82
17
25
11
45
137
69
47
30
72
06 AF03†
42
23-a*
Tephriphonolite
28a*
79
17
n.a.
118
4
73
n.a.
65
15
29
65
n.a.
n.a.
n.a.
13
44
057
34
06
15
022
14
02
65
078
13
41
40
05 BH
01
Mel.leucitite
23
79
1
58
103
26
037
23
035
12
16
22
45
45
05 IL
01†
Lamproite
26
7
084
46
077
2
029
19
028
12
2
52
15
92
05 BH
05†
05 IL
02†
Lamproite
SLB lavas
Tephriphonolite
53a*
Lamproite
PAB lavas
Leucitite
EAVF, Early Afyon volcanic formation. n.a., not analysed.
*Akal et al. (2013).
†
Prelević et al. (2012).
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Hf
Ta
Th
U
Pb
Rock
type:
Sample:
40C
EAVF
24
71
085
48
082
22
03
21
03
65
14
27
64
20
Latite
Table 1. Continued
536
Journal of Petrology, 2015, Vol. 56, No. 3
Journal of Petrology, 2015, Vol. 56, No. 3
537
Table 2: Whole-rock Rb, Sr, Nd, Sm concentration (mg g–1) and Sr and Nd isotope data for Afyon lavas, recalculated for the given
ages
Sample
PAB lavas
53a
66b
97
50
60
29
27-c
84
28b
31
SLB lavas
19-a
1-a
23-a
42
06 AF03*
05 BH 01*
05 IL 01*
05 IL 02*
05 BH 05*
EAVF
40C
Rock type
Sr
87
87
Sr/
Sr
Sr/
Sri
Nd
Age
(Ma)
Rb
Nosean phonolite
Latite
Latite
Latite
Phonotephrite
Phonotephrite
Shoshonite
Latite
Phonolite
Shoshonite
114‡
114‡
114‡
114‡
114‡
114‡
114‡
114‡
114‡
114‡
132
207
148
190
232
203
417
240
339
114
2507
3778
1780
1049
3647
2849
2144
2086
58
1672
0704769 6 7
0705670 6 6
0705445 6 7
0706008 6 6
0706048 6 8
0704349 6 8
0705590 6 7
0705378 6 7
0707758 6 8
0705559 6 4
070474
070564
070541
070592
070602
070432
07055
070532
070504
070553
86
236
96
73
147
125
80
81
26
80
Mel.-leucitite
Mel.-leucitite
Tephriphonolite
Tephriphonolite
Leucitite
Mel.-leucitite
Lamproite
Lamproite
Lamproite
114‡
114‡
114‡
114‡
114‡
114‡
119
119
119
492
294
233
250
535
605
173
201
200
2154
3082
2516
3247
2946
2193
1360
1253
1351
0706049 6 13
0705666 6 10
0706310 6 10
0706346 6 7
0706031 6 7
0705806 6 6
0707418 6 5
0707372 6 8
0708012 6 5
070594
070562
070627
070631
070595
070568
070742
07073
070794
85
77
64
104
94
66
51
65
25
Latite
144‡
128
1481
0706210 6 4
070617
67
86
86
143
143
144
144
Nd/
Ndi
eNd(T)†
12
35
15
12
22
18
12
14
2
12
0512664 6 5
0512613 6 5
0512602 6 5
0512561 6 5
0512693 6 5
0512697 6 5
0512633 6 4
0512583 6 4
0512645 6 6
0512610 6 5
051266
051261
051259
051255
051269
051269
051263
051257
051264
05126
07
–03
–06
–14
12
13
01
–09
04
–04
14
12
11
17
15
11
10
21
52
0512484 6 5
0512534 6 5
0512500 6 4
0512492 6 4
0512515 6 5
0512516 6 5
0512487 6 5
0512491 6 5
0512392 6 5
051248
051253
051249
051248
051251
051251
051248
051248
051238
–29
–19
–26
–27
–23
–22
–28
–29
–47
10
0512435 6 6
051243
–38
Sm
Nd/
Nd
EAVF, Early Afyon volcanic formation.
*Prelević et al. (2012).
†
eNd(T) is calculated using k147Sm ¼ 654E – 12 a–1, (147Sm/144Nd)0 CHUR ¼ 01967, (143Nd/144Nd)0 CHUR ¼ 0512638 and the concentration data given in the table.
‡
Akal et al. (2013).
sample-out line from the desolvating nebulizing introduction system (DSN-100 from Nu Instruments) to allow
for simultaneous aspiration of a dilute Tl solution for
mass bias correction. The accuracy of the analytical
protocol was evaluated with repeated analysis (n ¼ 21)
of Durango apatite and Ice River perovskite (see
Supplementary Data File 4). The average 87Sr/86Sr ratio
for the Durango apatite is 07060 6 00001 (2SD). This
value is slightly lower than the TIMS determined value of
0706327 (Horstwood et al., 2008). The average 87Sr/86Sr
ratio for the Ice River perovskite is 07032 6 00001 (2SD).
This value is slightly higher than the TIMS determined
value of 0702838 6 51 (Tappe & Simonetti, 2012).
CLASSIFICATION, PETROGRAPHY AND MINERAL
CHEMISTRY OF THE ROCK TYPES
Representative whole-rock analyses of Afyon lavas
(major and trace elements, and Sr, Nd and Pb isotope
ratios) are given in Tables 1 and 2; the full dataset,
including mineral composition data, is provided in
Supplementary Data Files 1–3. Representative photomicrographs of the Afyon alkaline volcanic rocks and detailed petrographic descriptions of the mafic cumulates
they host are reported in Supplementary Data File 2.
The lavas of the Late Afyon volcanic complex include
a broad range of mineralogically and petrographically
different rocks. We classified them according to the
IUGS classification scheme (Le Maitre, 2002) and the
systematics for potassic rocks and lamproites (Mitchell
& Bergman, 1991; Woolley et al., 1996). Furthermore,
we tentatively arranged them into two major groups,
according to their K2O contents and K2O/Na2O ratio,
and the presence or the absence of plagioclase and
amphibole (hornblende) on the one hand, and leucite or
sanidine on the other: plagioclase–amphibole-bearing
(PAB) lavas and sanidine- and/or leucite-bearing (SLB)
lavas. This approach is similar to the subdivision used
for Italian K-rich volcanic rocks (Conticelli et al., 2002,
2009; Perini et al., 2004), which are grouped into leucitebearing and leucite-free rocks. Here, we extend the mineralogical criteria to include more minerals, to avoid the
effects of heteromorphism that may result in the absence of leucite in nominally undersaturated ultrapotassic lavas owing to suppression of leucite crystallization
from hydrous magmas at high pressure (Yoder, 1986).
PAB lavas
Plagioclase–amphibole-bearing (PAB) lavas are potassium rich, but with K2O/Na2O ratios never exceeding
two. They are represented by potassic trachybasalts,
trachyandesites, trachytes and nosean-bearing phonolites (Fig. 3a) and very rare plagioleucitite, plotting
within the shoshonite series of the Peccerillo & Taylor
(1976) K2O–SiO2 diagram (Fig. 3b). The PAB lavas are
characterized by the presence of plagioclase 6 sanidine,
both in the groundmass and as phenocrysts in variable amounts. Olivine, clinopyroxene, phlogopite and
amphibole (hornblende) occur as phenocrysts. Fe–Ti
oxides and apatite are major accessories. No mafic
538
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 3. (a) Na2O þ K2O vs SiO2 diagram after Le Bas et al. (1986) for the volcanic rocks from the Late Afyon volcanic complex. The
grey line approximates the transition between alkaline and subalkaline magma series after Irvine & Baragar (1971). (b) K2O vs SiO2
(wt %) variation in the volcanic rocks (Peccerillo & Taylor, 1976). The red dashed line approximates the transition between ultrapotassic and shoshonitic lavas (see text for more explanation). PAB, plagioclase–amphibole-bearing lavas; SLB, sanidine- and/or
leucite-bearing lavas.
cumulate xenoliths have been observed. Most of the
clinopyroxene phenocrysts and microphenocrysts have
diopsidic compositions with variable MgO contents
(Fig. 4a). The amphibole is ferroan pargasitic hornblende to ferroan pargasite, distinguished by Mg/
(Mg þ Fetot) ¼ 047–068, relatively high K contents (up
to 042 atoms per 23 oxygens) and high AlVI (028–039
atoms per 23 oxygens). Plagioclase phenocrysts display
a wide diversity of morphologies from equant to elongate. Plagioclase composition dominantly falls in the
range An35–45 but cores show a broader range of
An30–55 (Supplementary Data File 2).
SLB lavas
Sanidine- and/or leucite-bearing (SLB) lavas are extremely potassium-enriched rocks with K2O/Na2O ratios
exceeding two. They include lamproites and meliliteleucitites. In these lavas, plagioclase and amphibole
(hornblende) are ubiquitously absent, and the most
abundant tectosilicate is either alkali feldspar or leucite.
Journal of Petrology, 2015, Vol. 56, No. 3
539
melanite, calcite, apatite, and opaque minerals (Akal,
2003).
The characteristic feature of the melilite–leucitite volcaniclastic deposits is the presence of mafic cumulate
xenoliths (enclaves) that have the following mineralogy:
clinopyroxene, phlogopite, melanite, melilite, leucite,
perovskite, ilmenite, spinel and apatite. (Representative
photomicrographs and detailed petrographic descriptions of these enclaves are given in Supplementary Data
File 2.) The cumulates range from 2 to 20 cm in diameter
and are randomly distributed throughout the volcaniclastic succession. They are subangular, but the smallest
enclaves exhibit angular outlines. Most samples are medium- to coarse-grained, showing equigranular, holocrystalline, panidiomorphic to hypidiomorphic and
cumulate textures. Two major groups of cumulates are
recognized: Group 1—clinopyroxene þ phlogopite þ
apatite 6 K-feldspar (sanidine); Group 2—clinopyroxene þ phlogopite þ melanite
garnet þ leucite þ K-feldspar þ apatite þ titanite þ perovskite þ spinel.
MINERAL CHEMISTRY
Major and trace elements in clinopyroxene
Fig. 4. Compositions of clinopyroxene from the Late Afyon volcanic complex. All data are included in Supplementary Data
File 2. (a) MgO vs Al2O3 (wt %); (b) Al vs Ti (a.p.f.u.). The field
for lamproite clinopyroxenes is from Mitchell & Bergman
(1991); fields for clinopyroxene from the Italian volcanic province: high-potassium series (HKS), potassium series (KS) and
transitional rocks are after Perini & Conticelli (2002) and references therein. The field for kamafugitic clinopyroxene is from
Prelević et al. (2005).
Lamproites plot in the field of trachyandesite and trachyte in a total alkalis–silica (TAS) diagram, whereas the
melilite-leucitite lavas plot in the foidite field (Fig. 3a).
All SLB lava samples plot above the shoshonite field in
the Peccerillo & Taylor (1976) diagram (Fig. 3b). We tentatively draw a line to mark the boundary between the
two lava groups from the Late Afyon volcanic complex,
and we add a new field named ‘Ultrapotassic series’, in
which the SLB lavas plot (Fig. 3b). In terms of mineralogy, the major phenocrysts in the lamproites are
Mg-rich phlogopite, Fo-rich olivine and Mg-rich, low-Al
clinopyroxene, set in a sanidine, K-richterite, clinopyroxene groundmass (Akal, 2008; Prelević et al., 2012).
Apatite, Cr-spinel and Fe–Ti oxides occur as major
accessories (Akal, 2008). The melilite-leucitite lavas are
composed of leucite and diopside phenocrysts set in a
groundmass of nepheline, barium feldspar, melilite,
Clinopyroxenes from the two subgroups of lavas (PAB
and SLB) and in the mafic cumulates clearly define two
broad arrays that show negative correlations in an
Al2O3 vs MgO diagram (Fig. 4a). The array for SLB lavas
and mafic cumulates is characterized by a flatter slope
in the Al2O3 vs MgO variation than the array for PAB
lavas. Most clinopyroxene crystals from Group 1 cumulates (Phl-clinopyroxenites) and lamproites plot together exhibiting high MgO concentrations, whereas
clinopyroxenes from Group 2 cumulates (leucite- and
melanite þ perovskite-bearing) are considerably more
evolved and have lower MgO and generally low Al2O3
and TiO2 concentrations. Clinopyroxenes from PAB
lavas plot in an array with a steeper slope in the Al2O3
vs MgO diagram (Fig. 4a). Further distinction between
clinopyroxenes can be made using an Al vs Ti plot
(Fig. 4b), showing that the clinopyroxenes from the SLB
and PAB lavas occupy distinct fields, whereas those
from the cumulates either show transitional compositions or fall in the field of PAB clinopyroxenes (Fig. 4b).
The trace element compositions furthermore demonstrate substantial differences between the clinopyroxenes from PAB and SLB lavas (Fig. 5a). The patterns for
several grains from PAB lavas are subparallel with
a high degree of light to heavy rare earth element
(LREE/HREE) fractionation (Fig. 5a). On the other hand,
clinopyroxenes from SLB lavas and mafic cumulates
show several different REE patterns, suggesting that the
clinopyroxenes within this group are genetically heterogeneous. The clinopyroxenes from Group 1 mafic cumulates and melilite-leucitites show strong LREE/HREE
fractionation, illustrated by high and relatively variable
La/Yb and Dy/Yb ratios. The patterns are subparallel,
showing a similar extent of REE fractionation at different bulk REE abundances (Fig. 5a), with clinopyroxene
540
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 5. (a, b) Chondrite-normalized rare earth element patterns for representative clinopyroxene, melanite garnet and melilite
phenocrysts and crystals in mafic cumulates from the Late Afyon volcanic complex; (c) primitive mantle-normalized incompatible
trace element patterns for representative apatite crystals from mafic enclaves; (d) chondrite-normalized rare earth element patterns
for representative perovskite crystals from mafic cumulates. Normalizing values after Sun & McDonough (1989).
from melilite-leucitites having the highest REE contents.
In contrast, the REE patterns of clinopyroxene from
lamproites are flatter (Fig. 5a), resulting in systematically lower La/Yb and Dy/Yb ratios. The highly evolved
clinopyroxenes from Group 2 (melanite-bearing) mafic
cumulates exhibit extremely fractionated and HREEdepleted patterns, showing that this clinopyroxene
crystallized after or simultaneously with melanite
garnet, which incorporated most of these elements
(Fig. 5a).
Trace elements in melanite garnet, apatite,
perovskite and melilite
Melanite garnet (Fig. 5b) shows a REE pattern with considerable LREE enrichment (>100 chondrite), as well
as HREE enrichment (>50 chondrite). This pattern differs substantially from the typical composition of garnet
from peridotites and eclogites, which shows ubiquitous
depletion in the LREE (<1 chondrite) and HREE enrichment (around 10 chondrite). This REE pattern suggests that melanite garnet in the Afyon cumulates has a
highly evolved character. Melanite shows enrichment
in most incompatible trace elements, with extreme
Th and U abundances (Supplementary Data File 3).
Strontium shows a very large trough in primitive mantle normalized patterns, suggesting considerable
fractionation of some Sr-rich coexisting phase (e.g. apatite) (Supplementary Data File 3). Coexisting melilite
also shows a strongly fractionated REE pattern (Fig. 5b).
Apatite from cumulate xenoliths shows extreme
LREE enrichment of up to 10 000 chondrite and HREE
as low as 2 chondrite, resulting in particularly high
LREE/HREE ratios (Fig. 5c). Important differences are
demonstrated by apatite from different groups of cumulate xenoliths with respect to the extent of HREE depletion and Th–U enrichment. Apatite crystals from Group
1 cumulates show less intense HREE depletion and
Th–U enrichment than those from Group 2 cumulates,
which are extremely depleted in HREE and, on the other
hand, show a huge Th–U peak (Fig. 5c).
Perovskite from Group 2 cumulates shows even
more extreme fractionation than apatite, with LREE enrichment up to >30 000 chondrite and HREE as low as
20 chondrite, resulting in particularly high LREE/HREE
ratios.
WHOLE-ROCK CHEMISTRY
Major and trace element compositions
The magmatic rocks of the Late Afyon volcanic complex
exhibit a broad range of compositions and include both
silica-undersaturated and silica-oversaturated lava
types (Akal, 2003, 2008; Akal et al., 2013). The PAB and
SLB subgroups define different compositional arrays in
most Harker diagrams, especially in plots of MgO vs
Al2O3, K2O, Na2O, P2O5 and K2O/Na2O (Fig. 6). The PAB
lavas are characterized by variable MgO (7–<1 wt %)
and relatively high concentrations of Al2O3 (up to 23%
Journal of Petrology, 2015, Vol. 56, No. 3
541
Fig. 6. (a)–(h) MgO vs SiO2, Al2O3, CaO, Na2O, K2O, TiO2, P2O5 and K2O/NaO2 (wt %) for lavas of the Late Afyon volcanic
complex. The shaded field in (h) is after Foley et al. (1987). Data for Eifel, Germany, are from Lustrino & Wilson (2007) and
references therein.
542
wt %) and Na2O (3–5 wt %). Both PAB and SLB lavas
have relatively low TiO2 (below 15%). PAB lavas have
systematically lower K2O contents and K2O/Na2O ratios
than SLB lavas (Fig. 6e and h). Interestingly, the major
element variation of PAB lavas shows considerable
resemblance to that of Pliocene alkaline lavas from the
Eifel Province (Germany). The Eifel volcanic district represents typical basaltic intracontinental volcanism characterized by an anorogenic chemical signature and a
spectrum of alkaline lava types from nepheline basalts
and basanites to phonolites (Mertes & Schmincke,
1985; Schmincke et al., 1985). In contrast, lamproites
and melilite-leucitites from the SLB subgroup show
marked differences in their major element composition,
especially in terms of MgO abundances: lamproites are
more MgO-rich with typically low Al2O3 and high SiO2
and P2O5, showing a positive correlation between MgO
and CaO (Fig. 6c), whereas the melilite-leucitites are
MgO-poor lavas with low SiO2, high Al2O3 and CaO and
very low P2O5 concentrations (Fig. 6). Both SLB lava
types are extremely potassium rich (7 to >12 wt %) and
have relatively low TiO2.
The lavas from both the PAB and SLB groups are
enriched in LILE and LREE. At first glance, they demonstrate the characteristic features of orogenic volcanic
rocks in primitive-mantle normalized trace element diagrams, including high LILE/HFSE, peaks at Pb and
troughs at Nb, Ta and Ti (Fig. 7). Although there is a significant overlap in the patterns, the orogenic signature
is more pronounced in the SLB lavas (e.g. higher Rb
and Ba contents), whereas the chemical fingerprint of
anorogenic magmas is more prominent in the PAB
lavas (e.g. higher Nb and HFSE contents). Within the
two groups, the most evolved haüyne-bearing phonolites of the PAB lavas show excessive depletion in Ba,
Sr, P, middle REE (MREE) and Ti, suggesting extreme
fractionation of apatite and amphibole. The meliliteleucitites are also depleted in P owing to apatite fractionation (see yellow arrow in Fig. 7b).
Isotopic compositions
The initial Sr and Nd isotopic compositions of the PAB
lavas fall in the range 143Nd/144Nd ¼ 051277–051255
and 87Sr/86Sr ¼ 070492–07078 (Fig. 8a; Table 2). SLB
lavas have more enriched isotopic signatures
with 143Nd/144Nd ¼ 051252–051238 and 87Sr/86Sr ¼
07052–07085. These ranges are similar to those of
K-rich lavas from the northern parts of western Anatolia
(Prelević et al., 2012). Less obvious differences between
the two Afyon lava groups are recognized in Pb–Pb
isotopic space (Fig. 8b and c). The PAB laves have
206
Pb/204Pb and 208Pb/204Pb ratios between 189 and
191 and 387 and 392, respectively, with highly variable 207Pb/204Pb ratios, whereas the SLB lavas are characterized by 206Pb/204Pb around 190 and 208Pb/204Pb of
390 (Fig. 8b and c).
To fingerprint the origin of the different types of
mafic cumulates found in the SLB lavas, we measured
in situ 87Sr/86Sr ratios in clinopyroxene, apatite,
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 7. (a, b) Primitive mantle-normalized trace element patterns for representative samples of volcanic rocks from the
Late Afyon volcanic complex. Primitive mantle composition
from Sun & McDonough (1989). Highly evolved samples 28a
and 28b of PAB lavas show extremely fractionated patterns.
perovskite and melilite (Supplementary Data File 4). All
grains were characterized by microprobe to select relatively homogeneous crystals and to avoid those with
abrupt compositional changes. The in situ 87Sr/86Sr
ratios of different minerals from five mafic cumulates
are similar to those of their host melilite-leucitite lavas
(Fig. 9). There are no systematic differences, either between mineral types or between different crystals from
different mafic cumulates (Group 1 vs Group 2). This indicates that the minerals from the cumulitic mafic enclaves are isotopically well equilibrated, both mutually
and with their host lava. Most importantly, the minerals
do not define a trend pointing toward the composition
of limestones from the Tauride belt, which are a potential crustal contaminant (see below).
DISCUSSION
The igneous rock suites from the Late Afyon volcanic
complex include different lithologies and have a considerable compositional range, which is controlled by multiple petrogenetic processes. A number of Afyon lava
samples that have sufficiently high MgO, Cr and Ni contents to be in equilibrium with mantle assemblages still
Journal of Petrology, 2015, Vol. 56, No. 3
543
Fig. 9. In situ analyses of 87Sr/86Sr in clinopyroxene (Cpx), apatite (Ap), perovskite (Per) and melanite garnet (Mel) from
Group 1 and Group 2 mafic cumulates from the Late Afyon volcanic complex. Whole-rock data for the host leucitites and
limestones from the Tauride belt (Sarı et al., 2004) are shown
for comparison.
complex, concluding that at least three types of parental
melts (one for PAB and two for SLB lavas) are involved,
and (2) the relationship of the Afyon volcanism with regional mantle dynamics and the mantle constraints on
the broadly synchronous association of contrasting alkaline magma compositions.
Low-pressure evolution and the origin of the
most extreme magma compositions
Issue of crustal contamination—metamorphic
basement rocks vs limestone
Fig. 8. Isotopic variations in volcanic rocks from the Late Afyon
volcanic complex and Early Afyon volcanic complex (EAVC):
(a) 87Sr/86Sr vs 143Nd/144Nd; (b) and (c) 206Pb/204Pb vs
207
Pb/204Pb and 208Pb/204Pb. Sources of data: Turkish high-Mg
mantle-derived ultrapotassic rocks of lamproitic affinity from
Prelević et al. (2012); Menderes metamorphic basement from
Prelević et al. (2012) and references therein; data for Afyon
basement rocks from D. Prelević & C. Akal (unpublished data);
NHRL, Northern Hemisphere Reference Line from Hart (1984).
demonstrate considerable isotopic and trace element
variations, pointing to the involvement of heterogeneous mantle source(s) and mixing of different mantle
components. Several lava types, most specifically those
with very low Mg# and compatible element contents,
cannot represent primitive magma compositions and,
therefore, must have experienced fractional crystallization, possibly in combination with low-pressure crustal
contamination. In the following sections we discuss (1)
the evolution of the lavas from the Afyon volcanic
The composition of primary melts may change significantly by crustal contamination and magmatic differentiation. In SW Anatolia, the basement largely consists of
Palaeozoic metamorphic rocks of the Menderes Massif
and the Afyon zone and of thick limestone formations of
the Tauride Belt. Assimilation of these two lithologies
would have fundamentally different effects on the geochemistry of the contaminated melts.
The Afyon metamorphic basement dominantly consists of Si-rich rocks and their assimilation would lead
to considerable Si-oversaturation. Therefore, the assimilation of metamorphic basement has a relatively restricted potential to affect the composition of the lavas
from the Late Afyon volcanic complex, especially the
most Si-undersaturated ones. Moreover, the extent of
trace element enrichment in the lavas is commonly several orders of magnitude higher than in the basement
rocks. For example, typical gneisses, mica schists and
phyllites from the Afyon metamorphic belt have
8–300 lg g–1 Sr and 6–40 lg g–1 Nd (D. Prelević et al., unpublished data) and their contamination potential is
small for those lavas that have more than 3000 lg g–1 Sr
and 70 lg g–1 Nd. Such contrasting abundances of these
544
Fig. 10. SiO2 vs Nb/U and Th/Nb to illustrate the role of lowpressure crustal contamination in the petrogenesis of volcanic
rocks from the Late Afyon volcanic complex; data for Kula volcanic rocks are from Grützner et al. (2013) and references
therein; data for Afyon basement rocks are from D. Prelevic &
C. Akal (unpublished data); data for OIB are from the GEOROC
database (http://georoc.mpch-mainz.gwdg.de/georoc).
elements between the lavas and the crust place stringent limits on the amount of upper crustal involvement
in their petrogenesis. The variation of SiO2, 87Sr/86Sr
and trace element ratios sensitive to contamination
(e.g. Nb/U, Th/Nb) (Fig. 10) can be used to qualitatively
evaluate the role of crustal contamination. The fact that
these parameters vary only little with SiO2 content
implies that assimilation of old metamorphic basement
rocks cannot be responsible for the compositional variations and the specific crustal signature observed in the
Afyon volcanic rocks. Similar arguments apply to enriched alkaline lavas such as lamproites and kamafugites from elsewhere in the Mediterranean region with
similarly high MgO, Ni and Cr contents and Mg-number
values, and containing highly forsteritic olivine phenocrysts and mantle xenoliths in some lavas (Peccerillo
et al., 1988; Conticelli & Peccerillo, 1989; Prelević &
Foley, 2007). Assimilation of crust would dilute the high
concentrations of trace elements and weaken the distinctive mantle-derived crustal chemical signature of
these rocks (Conticelli, 1998; Murphy et al., 2002;
Prelević et al., 2013), but not cause them.
Journal of Petrology, 2015, Vol. 56, No. 3
Assimilation of limestones by ultrapotassic basaltic
melts can produce contaminated lavas with Si-undersaturated compositions. Carbonate syntexis, first proposed for the origin of alkaline lavas (Daly, 1918;
Rittmann, 1933) and recently studied experimentally,
drives the crystallization of Ca-rich pyroxene, which results in desilication of the melt and an increase of Siundersaturation (Gaeta et al., 2006; Iacono-Marziano
et al., 2007, 2008, 2009; Conte et al., 2009; Mollo &
Vona, 2014). If the primary melt is of alkaline composition and is initially strongly enriched in Sr with a high
87
Sr/86Sr ratio, 5–10% of limestone syntexis (Gaeta
et al., 2006) will be cryptic with no significant effect on
87
Sr/86Sr, 143Nd/144Nd and LREE/HREE ratios (Peccerillo,
1985; Conticelli & Peccerillo, 1992; Di Battistini et al.,
1998; Conticelli et al., 2002; Perini et al., 2004). There
will be, however, an increase in the oxygen isotope
ratios (Gaeta et al., 2006). This contrasting behaviour reflects simple mass balance and depends on the contents (and isotopic compositions) of the relevant
elements in the different mixing end-members: limestones typically have high d18O values (25–30%; Turi,
1970) and oxygen is a major component in both the alkaline melt and the limestone. On the other hand,
87
Sr/86Sr, 143Nd/144Nd and LREE/HREE ratios do not
show a corresponding effect, either because the isotopic contrast is too small (87Sr/86Sr in limestone ranges
around 0707–0708) or the content of REE and Sr in the
limestone is too low (e.g. Turekian & Wedepohl, 1961).
Nonetheless, there are several trace-element parameters that are inconsistent with simple limestone syntexis and that have been used to question the model
(Savelli, 1967; Avanzinelli et al., 2009; Boari et al.,
2009a, 2009b; Martin et al., 2012). For instance, in the
Roman Province of Italy, the Si-undersaturated lavas
have higher incompatible trace element abundances
than both limestone and alleged primary melts of trachytic composition (Savelli, 1967). Several explanations
have been advanced to explain this; the most recent
treatment of the subject given by advocates of the limestone syntexis model is that a cryptic process of fluid
phase mobilization involving CO2 þ H2O þ HF, HCl and
H3PO4 is able to concentrate these trace elements in the
lavas and surrounding skarns (Wendlandt & Harrison,
1979; Lentz, 1996; Federico & Peccerillo, 2002).
Alternatively, this enrichment may be induced by
mantle metasomatism. Avanzinelli et al. (2009) proposed that carbonated mantle may yield highly
Si-undersaturated, ultra-enriched primary melts similar
to those found in the Roman Province if this mantle
is affected by carbonate-rich fluids derived from
subducted sedimentary components. Although the
above studies have fuelled a continuing debate over
the chemical, petrological and isotopic evidence and its
significance for limestone syntexis, the issue is no longer whether the above processes were active in the
Roman Province volcanic systems, but merely to what
extent mantle vs intra-crustal processes were playing a
role.
Journal of Petrology, 2015, Vol. 56, No. 3
545
Fig. 11. Variation diagrams for SiO2, Al2O3, CaO and K2O vs MgO (wt %) for the lavas from the Late Afyon volcanic complex.
Fractional crystallization vectors (the pink line for PAB lavas, the red line for leucite melilitites) for different mineral assemblages
from hypothetical parental melts are shown; small symbols denote 10% fractionation. A multi-step least-squares major element
mass-balance model of fractional crystallization for the PAB lavas is given in Supplementary Data File 5. 40c the dominant fractionating assemblage for the PAB lavas includes olivine, augite, amphibole, biotite and feldspars. In contrast, the parental melts of the
melilite-leucitites should have a proto-kamafugitic composition (KAM). Given their high K2O contents and K2O/Na2O ratios, dominant fractionation of clinopyroxenes will induce silica-undersaturation and amplify the K2O enrichment seen in the most evolved
samples. The compositions of partial melts of anhydrous and carbonated fertile lherzolite from experimental studies (Takahashi &
Kushiro, 1983; Hirose & Kushiro, 1993; Hirose & Kawamoto, 1995; Falloon et al., 1997; Hirose, 1997a, 1997b) are shown as progressive melting trends: HK66 and KLB1 are websterite and spinel-lherzolite, respectively, whereas KG-1 and KG-2 are spinel-lherzolites.
There are no substantial compositional differences between the melts in the different experimental studies.
Among the Afyon Si-undersaturated lavas, meliliteleucitites have high CaO contents and Si-undersaturated
compositions that may be, at first glance, interpreted as
due to syntexis of limestone. Some clinopyroxenes
from Group 1 enclaves have a ubiquitously high
Ca-Tschermak’s component that is in accordance with
experimental results showing that in a CaCO3-bearing
system (e.g. after limestone syntexis), Ca-Tschermak’s
and hedenbergite components increase as a function of
CaCO3 addition (Mollo & Vona, 2014). The overall evidence is far from clear-cut, however, as clinopyroxenes
from Group 2 enclaves and melilite-leucitites mostly
have low-Ca compositions. There are two mechanisms
alternative to carbonate syntexis that may yield similar
Si-undersaturation and Cpx chemistry (see discussion in
the following sections).
The existence of different parental melts and their
fractional crystallization
The most reasonable parental melt for PAB lavas resembles potassic trachybasaltic sample 40c (Supplementary Data File 1). These lavas evolve from high
MgO towards high SiO2, similar to most intracontinental basalts from the Circum-Mediterranean region and
consistent with fractionation of olivine and amphibole
(þ augite) from the more primitive magmas followed
by saturation with biotite, plagioclase and sanidine
(Figs 11 and 12). Such an evolutionary trend is also
supported by quantitative modeling of fractional crystallization (Supplementary Data File 5). An additional
test of the fractionation trends estimated by our
major element modelling is based on trace element
modelling, applying the same fractionating assemblages and using whole-rock chemical data (Fig. 13).
Details of the composition of the fractionating minerals
and their proportions are given in the caption of Fig. 11
and Supplementary Data File 5. This modelling demonstrates that the most evolved lavas may be produced
from the most primitive magmas by extreme fractionation of olivine and amphibole (þ augite, biotite and
feldspars).
In contrast, compared with the few lamproitic
rocks that plot together with the PAB lavas on a plot
of MgO vs SiO2, most of the evolved SLB lavas of
546
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 12. CaO vs MgO variation in clinopyroxene from Afyon lavas and enclaves. The results of MELTS simulations (Ghiorso & Sack,
1995; Asimow & Ghiorso, 1998) are illustrated using hypothetical parental melts of lamproitic (MELTS lamproite) and kamafugitic
(MELTS Kam.) composition at different crystallization pressures for fixed amounts of water (1 and 2 wt %). Data for carbonatites
and aillikites are from Tappe et al. (2004, 2006, 2008, 2009); data for kamafugites are from Prelević et al. (2005).
Fig. 13. (a) Zr vs Zr/Sc and (b) Nb vs Nb/Sc diagrams to illustrate the role of fractional crystallization in the petrogenesis of volcanic
rocks from the Late Afyon volcanic complex. Modelling shows that clinopyroxene-dominated fractionation can explain the variation
observed in SLB lavas. In contrast, the variation observed in the PAB lavas requires (Ol) þ Cpx þ Amph þ (Pl) fractionation, as also
inferred from the major element variations. Modelling details and mineral Kd data are given in Supplementary Data File 5; (c) Zr vs
Rb and (d) Zr vs Ba diagrams to illustrate the role of leucite and phlogopite accumulation in SLB lavas from the Late Afyon volcanic
complex.
Journal of Petrology, 2015, Vol. 56, No. 3
melilite-leucitite composition cluster in a low-SiO2 array
down to 43% (Fig. 6a). These lavas are among the most
Si-undersaturated rocks in the Afyon complex.
Estimation of the composition of the parental melt for
this group is not straightforward as there is considerable isotopic contrast between the most primitive and
the most evolved lavas of the group (Fig. 8; Table 2).
Furthermore, the evolved members of the group (melilite-leucitites) are considerably more Si-undersaturated
than the most primitive members (lamproites and
shoshonites). The isotopic composition of these lavas
has two important implications, as follows. (1) Meliliteleucitites (Fig. 8) and the cumulates they host (Fig. 9)
have less radiogenic 87Sr/86Sr and more radiogenic
143
Nd/144Nd than the limestones of the Tauride
belt (Sarı et al., 2004). This invalidates the assumption
that carbonate syntexis by lamproitic lavas accounts
for the isotopic variation seen in the melilite-leucitites,
simply because the lamproites have more radiogenic
87
Sr/86Sr and less radiogenic 143Nd/144Nd than the
limestones (Figs 8 and 9). (2) By ruling out the involvement of limestone assimilation, it becomes likely that
mantle source heterogeneity is responsible for their
extreme Si-undersaturation, which indirectly implies
the existence of different parental melts for lamproites
and melilite-leucitites. Importantly, both parental melts
should have ultrapotassic character, low Al2O3 contents and similar trace element abundances. The major
difference, however, should be in terms of SiO2 and
CaO contents, with the parental melt of the meliliteleucitites being more enriched in CaO and poorer in
SiO2 (Fig. 6).
This interpretation is further supported by the major
element characteristics of clinopyroxenes from the SLB
lavas and their cumulate enclaves (Figs 4 and 12). There
is a systematic variation in clinopyroxene compositions,
which occupy distinct areas on a plot of Al vs Ti (Fig. 4b)
in which the compositional arrays are defined empirically by the clinopyroxene compositions in K-rich orogenic lavas in the Mediterranean region (Perini &
Conticelli, 2002). Clinopyroxenes from lamproites and
melilite leucitites plot within the lamproite field,
whereas those from Group 1 enclaves partly resemble
kamafugitic clinopyroxenes. Together with lamproites,
kamafugites represent silicate lavas with the most extreme compositions erupted within the Mediterranean
area. These two lava types are considered to be compositional end-members of the ultrapotassic rock family
(Foley et al., 1987; Prelević et al., 2005), with lamproites
being Si-saturated to -oversaturated, with low Al2O3
and CaO (<10% and <6%, respectively), whereas
kamafugites are Si-undersaturated (SiO2 40–45%), also
having low Al2O3 but very high CaO (<10% and >12%,
respectively). Both groups of ultrapotassic lavas share
geochemical characteristics with magmas generated in
subduction-related (orogenic) tectonic settings.
In the CaO vs MgO diagram (Fig. 12), clinopyroxenes
from Group 1 cumulates resemble kamafugitic
547
clinopyroxenes, but also rare xenocrystic clinopyroxene
from melilite-leucitites and Group 2 cumulates and
clinopyroxene from carbonatites and ultramafic
lamprophyres (Tappe et al., 2004, 2006, 2008, 2009),
having similarly high CaO and an increased CaTschermak’s component. The highest CaO contents are
observed for clinopyroxene from Group 1 enclaves, and
a few xenocrystic clinopyroxene grains from meliliteleucitites and Group 2 cumulates. On the other hand,
clinopyroxenes from the other lava types plot away
from the high-CaO array (Fig. 12). To test whether early
clinopyroxene crystallizing at different pressures from
different melt compositions will have comparably high
CaO contents, we conducted MELTS simulations
(Ghiorso & Sack, 1995; Asimow & Ghiorso, 1998) using
parental melts of lamproitic and kamafugitic composition (Fig. 12). We explored the variation of crystallization trajectories, fractionated mineral assemblages and
Cpx compositions as a function of melt composition
and pressure, to constrain the conditions for which the
observed composition of lavas and their constituent
clinopyroxene, including cumulitic enclaves, could be
generated. Our modelling results, together with data for
other CaO-enriched lavas, clearly demonstrate that only
Ca-enriched parental melts such as kamafugites and
ultramafic lamprophyres can crystallize high-CaO clinopyroxene (Fig. 12).
Based on the above modelling, we infer that the parental magma of the melilite-leucitite lavas should have
a proto-kamafugitic composition. Given their high K2O
contents and K2O/Na2O ratios, dominant fractionation
of clinopyroxene induces silica-undersaturation and
amplifies K2O enrichment, as seen in the most evolved
samples (Figs 6e, h and 11). However, it is clear from
Fig. 11 that another process is needed to achieve very
high K2O and K2O/N2O as well as an increase in Al2O3,
as observed in the most evolved melilite-leucitites. We
propose that the second stage of magma evolution is
dominated by accumulation of leucite owing to its lower
density than the host melt, resulting in its flotation
(O’Brien et al., 1991). This results in high K2O, K2O/N2O
and Al2O3 for invariantly low MgO concentrations
(<2%). Petrographically, melilite-leucitites are almost
bi-mineralic lavas that in some cases resemble leucite
cumulates with intercumulus melilite microcrysts
(Supplementary Data File 2; Akal, 2003). Major element
modelling and fractionation paths modelled using trace
elements, applying the same fractionating assemblages
and using whole-rock chemical data together with
mineral compositions of cumulitic enclaves exhumed
by the melilite-leucitites, are shown in Figs 11 and 13.
Details of the composition of the fractionating minerals and their proportions are given in the figure
captions. Modelling demonstrates that the most
evolved melilite-leucitites could be produced from a
kamafugitic parental melt (KAM in Figs 11 and 13) by
extreme clinopyroxene fractionation combined with
leucite accumulation.
548
Fig. 14. Pressure (kbar) vs T ( C) diagram based on P–T calculations for lava phenocrysts and mafic enclaves from the Late
Afyon volcanic complex, using the clinopyroxene–liquid thermobarometer of Putirka (2005) and the amphibole–liquid thermobarometer of Ridolfi & Renzulli (2012). Full dataset is
presented in Supplementary Data File 6.
Thermobarometric limitations and a summary of
intra-crustal evolution
Pressure constraints for the crystallization of plagioclase–amphibole-bearing lavas can be estimated using
the clinopyroxene thermobarometers of Putirka (2008)
applying equation 32d for the temperature and equation 32c (Al distribution between lava and Cpx) for the
pressure, and the Microsoft Excel spreadsheet
RiM69_Ch03_cpx_PT. Values for the liquid composition
were taken from the whole-rock analyses of the lava
samples. The input parameters and estimated P and T
data are summarized in Supplementary Data File 6.
The estimated crystallization pressures of clinopyroxene from the PAB lavas range from 9 to 12 kbar at
temperatures ranging from 1040 to 1150 C. The amphibole-based thermobarometer applicable to alkaline
melt compositions (Ridolfi & Renzulli, 2012) yields pressure estimates of 5–10 kbar at temperatures ranging
from 1000 to 1150 C (see Supplementary Data File 6).
The results, summarized in Fig. 14, imply that the entire
spectrum of variably fractionated PAB lavas is generated by polybaric fractional crystallization between 10
and 45 km depth.
The same approach can be applied to constrain the
temperature and pressure of crystallization of the SLB
lavas and associated mafic cumulates. Values for the liquid composition were taken from the whole-rock analyses of the lava samples. In the case of clinopyroxene
from the mafic cumulates the assumption was made
that the equilibrium melt was of kamafugitic (Prelević
et al., 2005) and melilite–leucititic composition for
Group 1 and Group 2 cumulates, respectively. The input
parameters and estimated P–T results are summarized
in Supplementary Data File 6. The estimated crystallization pressures of clinopyroxene from lamproites, melilite-leucitites and Group 2 cumulates indicate shallow
Journal of Petrology, 2015, Vol. 56, No. 3
crystallization depths, ranging from 7 kbar down to zero
at temperatures ranging from 1050 to 1250 C. In contrast, Group 1 cumulates show considerably and ubiquitously higher crystallization depths ranging from 31 to
>50 km (from 11 kbar to 17 kbar) at temperatures ranging from 1190 to 1150 C (Supplementary Data File 6;
Fig. 14).
The above P–T estimates suggest that the PAB
lavas most probably underwent polybaric differentiation,
beginning at a depth of 45 km. In contrast, SLB lavas
crystallized along two different pressure paths: Group 1
(hereafter referred to as high-P) cumulates represent
early high-pressure crystallization products from a maximum depth of 50 km, derived from a melt that has a
kamafugitic composition, whereas Group 2 (hereafter
referred to as low-P) cumulates, melilite-leucitites and
lamproites are shallow crystallization products from a
depth of not more than 20 km. The contrasting depth of
enclave equilibration indicates that the high-P cumulates
formed before the low-P cumulates.
The idea that the two groups of cumulates from SLB
lavas represent early and late crystallization stages is
further supported by the abundances of certain trace
elements and their ratios in cumulate minerals (Fig. 15).
Early crystallization of high-P cumulates is suggested
by Sc and Cr enrichment in clinopyroxene and complementary depletion in the clinopyroxene from the low-P
cumulates (Fig. 15a), confirming that clinopyroxene
was a major phase in the early fractionates. Similar conclusions may be drawn from the phlogopite composition, which shows enrichment in Cr and relative
depletion of Ba in high-P phlogopite, and balancing depletion of Cr but enrichment of Ba in phlogopite from
the low-P cumulates (Fig. 15b), confirming the incompatible behaviour of this element owing to the dominance of Cpx over phlogopite in the early fractionating
assemblage. Apatite composition differs considerably
in the two types of cumulates (Fig. 15c). Early crystallization in the high-P cumulates may explain the Th and
La depletion in apatite, whereas complementary enrichment is observed in apatite from the low-P cumulates.
This implies that apatite fractionation took place dominantly during the second stage (low-pressure) fractionation resulting in the P2O5 depletion in the leucititic
lavas (Fig. 6g).
To summarize the above discussion and modelling results, we suggest that the two subgroups of Afyon lavas
evolved as discrete volcanic lineages that are not related
by fractional crystallization to a common parental
magma. The role of assimilation of crustal rocks is of limited importance, meaning that much of the geochemical
diversity of the lavas results from mantle processes. The
Afyon magmas evolved along multiple differentiation
pathways from different parental magmas.
Mantle source dynamics and the character of the
parental magmas
The composition of the western Anatolian volcanic
rocks shows a systematic large-scale variation in the
Journal of Petrology, 2015, Vol. 56, No. 3
549
Fig. 16. (a) Si-saturation index Q [normative q – (ne þ lc þ kls)
vs age (Ma) of western Anatolian volcanism. (b) Si-saturation
index vs K2O/Na2O of lavas from the Late Afyon volcanic
complex. Data are from Akal et al. (2013) and references
therein.
Fig. 15. In situ trace element (mg g–1) bivariate diagrams
for clinopyroxene phlogopite (Phl), (Cpx) and apatite
(Ap) from mafic cumulates from the Late Afyon volcanic
complex.
parameter Q [normative q – (ne þ lc þ kls)], which is
used as a measure of Si-saturation. This parameter
clearly decreases from north to south with time
(Fig. 16). Importantly, in a Q vs age of eruption diagram,
Q changes sign in the time interval from 12 to 10 Ma,
which is the time when the Late Afyon volcanic complex
developed (Fig. 16a) (Akal et al., 2013). This implies that
the Afyon lavas represent an integral part of this more
regional trend, reflecting the response of the composition of the magmas to regional mantle dynamics, as
also illustrated by a transition in their radiogenic isotope compositions (Fig. 17). It is evident from these data
that mixing between different mantle source components may represent a first-order process in the petrogenesis of the magmas. In the following sections we
combine available geochemical data for the whole volcanic province with the data from this study and discuss
in detail the geodynamic setting of the Afyon alkaline
volcanism.
Fingerprinting different mantle components
within metasomatically modified mantle
We can broadly distinguish between components
derived from the convecting mantle (asthenosphere)
and those from the lithospheric mantle, contributing to
the final magma compositions.
(1) An asthenosphere-derived ‘anorogenic’ component.
Regionally, an anorogenic component is most prominent in the southernmost late Miocene–Pliocene
Bucak and Denizli volcanic provinces (Prelević et al.,
2012), as well as in the Quaternary Golcuk volcano
(Platevoet et al., 2008, 2014). This component has
unradiogenic 87Sr/86Sr, and radiogenic 143Nd/144Nd,
176
Hf/177Hf and 206Pb/204Pb, similar to the FOZO
(FOcal ZOne) (Hart et al., 1992; Stracke et al., 2005)
or HIMU (high-l) components recognized in OIB.
Coupled with variable, but ubiquitously high, Nb/U
and Ce/Pb ratios (Fig. 17), this geochemical
signature has affinities to melts derived from the
convecting mantle, similar to the EAR (European
550
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 17. Identification of mantle components contributing to the Afyon alkaline magmatism based on characteristic trace element
and radiogenic isotope ratios: (a) 206Pb/204Pb vs 87Sr/86Sr; (b) 87Sr/86Sr vs Nb/U; (c) 143Nd/144Nd vs Ce/Pb; (d) 87Sr/86Sr vs Th/Nb.
Yellow lines delineate hypothetical mixing hyperbolae (Prelević et al., 2012). Reference data: CiMACI (Circum-Mediterranean
Anorogenic Cenozoic Igneous Province) from Lustrino & Wilson (2007); upper crust from Rudnick & Gao (2003); data for Turkish
high-Mg mantle-derived ultrapotassic rocks of lamproitic affinity from Prelević et al. (2012); data for OIB from the GEOROC database (http://georoc.mpch-mainz.gwdg.de/georoc).
Asthenosphere Reservoir), and overlaps with that of
intra-plate basalts from the Circum-Mediterranean
Anorogenic Cenozoic Igneous Province (CIMACI)
and OIB. In the Late Afyon volcanic complex the
presence of the anorogenic component is mostly
evident in the PAB lavas, most distinctly in their
Nd–Sr isotopic composition (Fig. 17). This component also contributes to the parental melts of the
melilite leucitites (see below). It shows a clear southward increase with time (Prelević et al., 2012).
Importantly, none of the samples within the studied
volcanic suites fingerprint a ‘pure’ anorogenic component; their geochemical characteristics always
represent a mixure with an orogenic component.
Thus, the presence of an anorogenic component in
the Afyon lavas is largely inferred from the regional
geodynamic setting.
Melting of peridotite from the convecting mantle, at
first glance, is a viable model for the origin of this anorogenic component, based upon the results of high-pressure experimental melting studies of peridotite (e.g.
Jaques & Green, 1980; Green & Falloon, 1998; Herzberg
& O’Hara, 1998). The compositions of partial melts of
anhydrous and carbonated fertile lherzolite from a
range of experimental studies (Takahashi & Kushiro,
1983; Hirose & Kushiro, 1993; Hirose & Kawamoto,
1995; Falloon et al., 1997; Hirose, 1997a, 1997b) are
summarized and highlighted as progressive melting
trends in Fig. 11. According to these experiments, PAB
parental magmas may be produced by melting of anhydrous and/or hydrous peridotite, although the limited
number of PAB lavas with MgO > 6% restricts the possibility for a detailed comparison. Nevertheless, the major
element compositions of the high-MgO PAB samples
are mostly consistent with the melts produced by relevant peridotite melting experiments.
Figure 18 shows the Dy/Yb versus La/Yb variation in
calculated melts in equilibrium with clinopyroxene in
the Afyon lavas, compared with partial melts of garnetand spinel-peridotite and whole-rock data from the PAB
and SLB lavas and enclaves. The data are indicative of
the presence of garnet in the mantle source and roughly
constrain the depth of partial melting. They indicate
mixing between melt batches derived from the garnet
and spinel stability fields, reflecting polybaric melting of
the mantle source. It is highly plausible that the PAB
lavas are produced by melting predominantly in the
Journal of Petrology, 2015, Vol. 56, No. 3
Fig. 18. La/Yb vs Dy/Yb variations of melts calculated to be in
equilibrium with clinopyroxene phenocrysts from Afyon lavas
and mafic cumulates. The whole-rock data for the volcanic
rocks are also plotted. Melt compositions were calculated
using the method of Wood & Blundy (1997). Fields of melt
compositions for melting in the garnet and spinel peridotite
stability fields are from Prelević et al. (2012),
garnet stability field, whereas the role of garnet diminishes in the SLB lavas.
However, there is a substantial discrepancy between
the chemistry of the PAB lavas and peridotite melting
models in terms of K2O and Na2O contents (Fig. 11),
trace element signatures (Figs 7 and 10) and isotopic
composition (Fig. 8). All these parameters are inconsistent with the exclusive contribution of melts from ambient convecting mantle. It is more likely that
asthenosphere-derived melts were contaminated with
melts derived from the lithospheric mantle containing
an orogenic chemical signature. This contamination
could account for the chemical discrepancy between
major element compositions and trace element and isotopic compositions of the PAB lavas (see below).
(2) Lithospheric mantle component(s) of different origin.
The involvement of two types of metasome is recognized. Regionally dominant metasome (M 1), which
is ultimately derived from the continental crust, is responsible for the enriched isotopic composition of
these lavas with radiogenic 87Sr/86Sr (up to 0712)
and 207Pb/204Pb (up to 157), and unradiogenic
143
Nd/144Nd (as low as 05120) and 176Hf/177Hf (down
to 028245) (Prelević et al., 2012). This isotopic enrichment complements trace element enrichment
indicated by high LILE/HFSE and high canonical
trace element ratios (Hofmann & White, 1982) such
as Th/Nb, Hf/Sm, Th/La and Sm/La, and low Ce/Pb
and Nb/U. In the Late Afyon volcanic complex this
component is most evident in the composition of the
SLB lavas that are of lamproitic and shoshonitic
composition (Fig. 17), although it is present to a variable extent in all Afyon volcanic rocks. Melting of
peridotite cannot produce melts of such composition, in line with experimental studies showing
that even extremely small melt fractions will not
have ultrapotassic and ultra-enriched compositions
551
similar to those of the Late Afyon volcanic complex
(Foley, 1992). Moreover, the smallest amount of extractable melt (2–3%), which is controlled by the permeability of an interconnected melt network at low
melting degrees, represents a threshold barrier for
any ultra-small melt fraction to reach the surface (Faul
& Fitz, 1999; Faul, 2001). A better explanation is that
this component is stored in phlogopite–pyroxene-rich
metasomes within the lithospheric mantle (Prelević
et al., 2012). These metasomes formed during earlier
subduction episodes when sediments were subducted and melted within the mantle wedge, producing high-Si melts (Sekine & Wyllie, 1982) that
interacted with the overlying lithospheric mantle peridotite (Frost, 2006) to yield hydrous, phlogopite-rich
assemblages (Prelević et al., 2013).
Conversely, it is the addition of an ‘anorogenic’
metasomatic component (M 2) that induced the transitional geochemical character (i.e. between the most enriched SLB and PAB lavas; Figs 8 and 17) observed in
the melilite-leucitites. The parental magma of the melilite-leucitites should be of proto-kamafugitic composition with strong Si-undersaturation, implying that this
component is not produced by melting of normal mantle peridotite. Liquidus experiments using different
kamafugitic compositions (Edgar et al., 1976, 1980;
Arima & Edgar, 1983), melilitites and nephelinites
(Bultitude & Green, 1968, 1970, 1971; Brey, 1978; Gee &
Sack, 1988) have demonstrated that multiple saturation
with the four peridotite minerals (olivine, orthopyroxene, clinopyroxene and garnet or spinel) does not occur
anywhere close to the liquidus (Foley et al., 2012). The
common occurrence of clinopyroxene, olivine and
phlogopite in these experiments favours the interpretation that these minerals play a more important role in
the source region than would be the case for melting
of peridotite alone. Experiments suggest the presence
of non-peridotitic, ultramafic assemblages in the source
of kamafugites (Foley et al., 2012), which most probably
resemble phlogopite-wehrlite (Edgar et al., 1976, 1980;
Arima & Edgar, 1983; Thibault et al., 1992), resulting
from the interaction between carbonatitic melts and
mantle peridotite (Yaxley et al., 1991, 1998). There are
two mechanisms that can produce wehrlitic metasomes, with reference to the mantle dynamics of the
Mediterranean region. (1) Subduction recycling of
CaCO3-rich subducted marly sediments (e.g. Avanzinelli
et al., 2008; Bianchini et al., 2008; Frezzotti et al., 2009;
Grassi & Schmidt, 2011; Martin et al., 2012) and high
XCO2 during melting (McNeil & Edgar, 1987) is able to
induce Si-undersaturation of the resulting melts, which
will inevitably react with the surrounding mantle peridotite producing metasomes of wehrlitic composition.
The composition of the metasomes is controlled by the
nature and composition of the recycled continental
crustal material in the mantle wedge above the subducting lithosphere (Conticelli & Peccerillo, 1992;
Avanzinelli et al., 2008, 2009; Conticelli et al., 2013). (2)
552
The convecting mantle provides CO2 during metasomatism of the continental lithosphere (Thibault et al., 1992
and references therein). This explanation is preferred
for Afyon, as the trace element and radiogenic isotope
compositions of Afyon lavas show a transition from an
orogenic signature dominated by radiogenic 87Sr/86Sr
and unradiogenic 143Nd/144Nd, as seen in the lamproites, to an anorogenic signature with unradiogenic
87
Sr/86Sr and radiogenic 143Nd/144Nd as observed in the
melilite-leucitites. Sediment recycling would also result
in a transition, but in the opposite direction; that is, towards more radiogenic 87Sr/86Sr and less radiogenic
143
Nd/144Nd in the melilite-leucitites. The high XCO2 and
CaO may have been provided by the deep asthenospheric component, as in many intracontinental alkaline
mafic magmatic provinces (Lustrino & Wilson, 2007). Its
origin may be due to lithospheric mantle metasomatism by melting derived from the convecting mantle,
acting as a precursor to the dominantly anorogenic
Na-alkaline magmatism (PAB lavas). These asthenosphere-derived precursor melts probably formed in the
dolomite–garnet peridotite stability field, and reacted
with mantle peridotite along the solidus ledge in the
system lherzolite þ CO2 (<22 kbar), producing phlogopite-bearing wehrlitic metasomes (Green & Wallace,
1988; Thibault et al., 1992). Carbonatite-like metasomatic melts reacted with mantle minerals at pressures as
high as 22 kbar to produce a olivine þ clinopyroxene þ
phlogopite 6 chromite metasomes at the expense of
orthopyroxene. This second metasome type occurs within
the lithosphere: it is just as enriched in potassium as the
orogenic metasome type, meaning that it originated recently (Frost, 2006), owing to the necessity that phlogopite
is one of the major metasomatic minerals that convey this
geochemical signature (Prelević et al., 2012). In a further
step, the phlogopite-wehrlite assemblage melts incongruently (Edgar et al., 1976, 1980; Arima & Edgar, 1983)
because of its low solidus temperature, producing silicate
melt of proto-kamafugitic composition.
Figure 18 demonstrates that these two metasome
types identified in the composition of the SLB lavas
occur within the lithospheric mantle. The La/Yb vs
Dy/Yb diagram (Fig. 18) shows that melt production is
dominantly in the spinel stability field with a diminishing contribution from the garnet stability field.
Three major inferences are compelling from the
above discussion, as follows. (1) The orogenic component is dominantly present in the lithospheric mantle
under western Anatolia, and most specifically in the
Afyon region. (2) The anorogenic component ultimately
originated from the convecting mantle, and was conveyed into and reacted with the lithospheric mantle in
the form of a precursor melt that crystallized phlogopite-wehrlite metasomes, which produced protokamafugitic melts in later melting events. Importantly,
this second type of metasome crystallized in an already
metasomatized lithospheric mantle. This explains the
ubiquity of the orogenic signal in SLB lavas. (3) The
same convecting mantle, in a more advanced stage
Journal of Petrology, 2015, Vol. 56, No. 3
of melting, will result in the dominantly anorogenic
Na-alkaline magmatism that has some transitional
characteristics owing to interaction with the crustally
contaminated lithospheric mantle, which explains
the partial decoupling of the isotopic and trace element signatures. In this view, the transitional compositions found in PAB and some SLB lavas are generated
by instantaneous mixing during melting and interaction of the asthenosphere- and lithosphere-derived
components.
Implications for a viable geodynamic model
The geology of the Aegean region and western
Anatolia was dramatically shaped during the
Alpine–Himalayan orogeny, in which several continental crustal blocks intercalated with ophiolitic terranes
of various size and age collided, resulting in a complex collage of terranes. The major process that
accommodated most of the Africa–Europe convergence since Cretaceous times was nappe stacking and
lithospheric slab underthrusting. Since the subduction
of oceanic lithosphere started at 23 Ma (Jolivet &
Brun, 2010; Jolivet et al., 2013) and initiated the volcanism within the Aegean arc (Fytikas et al., 1984), the
region underwent intense extension, slab roll-back
and trench retreat (van Hinsbergen et al., 2005; Jolivet
& Brun, 2010), which is proposed to provide the general driving force for the post-collisional magmatism.
Magma genesis is controlled by a combination of the
roll-back of the underthrust lithospheric slab (van
Hinsbergen et al., 2005; Jolivet & Brun, 2010) that initiated post-collisional extension and collapse of the orogenic belt, coupled with the initiation and progression
of a slab tear (Fig. 1). Mantle circulation is additionally
induced by fragmentation of the lithospheric slab in
the East Mediterranean mobile zones that generated
vigorous upwellings (Faccenna & Becker, 2010). Such
complex geodynamic settings can activate different
mantle and crustal sources (Prelević & Seghedi, 2013):
the suction of hot convecting mantle into the mantle
wedge generated by slab roll-back (Aldanmaz et al.,
2000) in combination with incursion of fresh mantle
through a slab tear (Dilek & Altunkaynak, 2009;
Prelević et al., 2012) has the potential to activate previously enriched domains within the lithospheric mantle, resulting in the generation of post-collisional mafic
magmas (Prelević et al., 2012).
It is widely accepted that the lithospheric mantle
under western Anatolia is heavily metasomatized, dominantly by melts derived from subducted continental
crust-derived sediments, and that melting of this mantle
is responsible for the widespread presence of the orogenic fingerprint observed in the Tertiary lavas of western Anatolia (Prelević et al., 2010, 2012; Ersoy et al.,
2011; Ersoy & Palmer, 2013; Karaoğlu & Helvacı, 2014).
The chemical transition observed in the lavas of the
Late Afyon volcanic complex reflects the increasing role
of a newly formed metasomatic component, which ultimately originated from the convecting mantle, in the
Journal of Petrology, 2015, Vol. 56, No. 3
553
Fig. 19. Sketch cross-section of the melting region in the mantle beneath the Late Afyon volcanic complex; location shown in Fig. 1.
This illustrates the geodynamic setting of the magmatism during the Tertiary post-collisional tectonic development of western
Anatolia, based on tomographic models indicating tearing of the Aegean and Cyprus slabs combined with chemical and petrological data for the volcanic rocks. (a)–(d) show the texture of cumulate enclaves and two types of metasome within the lithospheric
mantle. It should be noted that the texture of Group 1 and 2 enclaves (a) and (b) is based on the observations made on the available
samples (Supplementary Data File 2), whereas the textures of the metasomes (c) and (d) are interpretative (see text for details).
554
Journal of Petrology, 2015, Vol. 56, No. 3
Table 3: Whole-rock Pb, U and Th concentration (mg g–1) and Pb isotope data for Afyon lavas, recalculated for the given ages
Sample
PAB lavas
53a
28a
66b
97
50
60
27-c
84
28b
31
SLB lavas
19-a
1-a
23-a
42
06 AF03*
05 BH 01*
05 IL 01*
05 IL 02*
05 BH 05*
EAVF
40C
Rock type
Age
(Ma)
Pb
U
Th
206
207
204
204
Pb/
Pb
Pb/
Pb
208
206
207
208
204
204
204
204
Pb/
Pb
Pb/
Pb
Pb/
Pb
Pb/
Pb
Nosean phonolite
Phonolite
Latite
Latite
Latite
Phonotephrite
Shoshonite
Latite
Phonolite
Shoshonite
114†
114†
114†
114†
114†
114†
114†
114†
114†
114†
30
72
49
24
16
37
37
51
71
54
69
47
21
13
10
25
11
15
28
11
45
137
74
47
26
79
37
37
136
41
19037
19112
18988
18958
18974
19138
1887
18967
19108
19047
1566
15621
1568
15673
15683
15672
15688
15694
15737
15719
38918
3915
38898
38902
38904
39081
38854
3897
39226
39179
1903
1907
1881
1884
1894
1905
1885
1891
1909
1903
1566
1562
1567
1567
1568
1567
1569
1569
1574
1572
3889
3911
3869
3877
3888
3899
3883
3893
3919
3916
Mel.-leucitite
Mel.-leucitite
Tephriphonolite
Tephriphonolite
Leucitite
Mel.-leucitite
Lamproite
Lamproite
Lamproite
114†
114†
114†
114†
114†
114†
119
119
119
121
30
87
138
139
92
45
48
40
17
19
23
22
21
15
45
4
41
56
57
41
65
65
52
22
23
13
18866
19121
1869
18687
18853
1887
1887
19101
18986
15719
15738
15696
15698
15704
15688
15705
15691
15701
38941
39266
38737
38745
389
38879
38953
39002
38981
1885
191
1869
1868
1885
1885
1886
1909
1898
1572
1574
157
157
157
1569
157
1569
157
3892
3924
3872
3871
389
3886
3893
3898
3897
Latite
144†
20
64
27
18943
15715
39021
189
1571
3896
EAVF, Early Afyon volcanic formation.
*Prelević et al. (2012).
Akal et al. (2013).
†
mantle source. This mantle component provided the
high XCO2, the enrichment in HFSE that induced lower
LILE/HFSE, and an isotopic signature typical of the convecting mantle. The convecting mantle is considered to
be the source for primary carbonatitic melts generated
from the dolomite–garnet peridotite stability field,
which reacted with mantle peridotite along the solidus
ledge, ultimately resulting in the crystallization of
phlogopite-wehrlite assemblages. Further melting of
these assemblages may contribute to all of the melts
present in the Late Afyon volcanic complex (both PAB
and SLB lavas), but it was least cryptic in the production
of the melilite leucitites.
Figure 19 illustrates the structure of the mantle and
melt evolution pathways beneath the Late Afyon volcanic complex. Volcanism is strongly controlled by the
lithospheric mantle, which must be internally heterogeneous at scales similar to those of melting and magma
extraction (i.e. metres to kilometres). Two metasome
compositions residing in the lithospheric mantle play a
role in the origin of the primary melts of the SLB lavas
(Prelević et al., 2012). Both of these are phlogopite-bearing metasomatic veins. On a regional scale, it is the
interplay of the distribution of these metasomes and
the shape of asthenospheric upwelling that controls the
spatial distribution of volcanism. Melts from these
metasomes account for the isotopic composition and
chemical fingerprints indicative of different geochemical reservoirs. The increasing extent of interaction and
mutual hybridization of the melts derived from such
metasomes occurred simultaneously with magmatic
differentiation.
The volcanism within the Late Afyon volcanic complex started with the most mafic representatives of the
SLB lavas (Fig. 2) and is directly related to the position
of the slab tear, as heat transfer from the upwelling asthenosphere triggered the magmatism. Heat supply
from the adiabatically ascending asthenosphere
induced partial melting in the most fertile parts of the
lithospheric mantle (Fig. 19). The base of the melting
zone was controlled by the solidus of mantle peridotite,
whereas its top should be situated in colder mantle regions and should be controlled by the stability of the
metasomatic assemblages (essentially phlogopite) in
the lithosphere. Therefore, initial melting will take place
in the most enriched portions of the lithospheric mantle,
which are composed of phlogopite-pyroxenite metasomes (M 1, Fig. 19c), representing the source of the
early Si-saturated SLB lavas of lamproitic composition
(Fig. 2). These melts will dominantly have an orogenic
chemical signature. Simultaneously with the upwelling,
low-degree partial melting in the asthenospheric mantle
at pressures >20 kbar produced dolomite carbonatitelike alkaline melts. At a depth 75 km, these precursor
alkaline melts start to interact with the mantle lithosphere, most specifically eliminating orthopyroxene
and producing a new generation of metasomes of
wehrlitic composition (M 2, Fig. 19d). The depth of this
interaction may play a substantial role in determining
the mineral composition of the metasomes. With continued asthenosphere upwelling, these metasomes
may melt again, possibly lowering SiO2 contents and
increasing CaO abundances in the parental melts of the
melilite-leucitites (Fig. 19). Melilite-leucitite magmas are
Journal of Petrology, 2015, Vol. 56, No. 3
generated by intense clinopyroxene fractionation, first
in a deep magma chamber (Fig. 19a) and later in a system of shallower magma chambers (Fig. 19b). The
mafic cumulates may have formed in different systems
of magma chambers and may have been included in
stages following the initial alkaline volcanism in the
Afyon area (Fig. 2). They were entrained in the rising
magma just prior to eruption. The fact that both types
of enclave (high-P and low-P) show general isotopic
equilibration with their melilite-leucitite host lavas
(Fig. 9) indicates homogenization of the melts, which
are derived from a heterogeneous mantle source, already in the deep chambers. Further replenishment of
the magma conduit system within the Late Afyon volcanic complex occurred at a variety of depths
(50–15 km). Thus, there were a series of shallow reservoirs connected with a recharging system at greater
depth (Fig. 19). Major element data indicate that a significant component of the PAB magmas could have
formed by low-degree partial melting of (carbonated)
peridotite at conditions prevailing in the convecting
upper mantle. Thus, the same asthenosphere upwelling
that produced the precursor melts for the Si-undersaturated SLB lavas produced the primary PAB lavas by
more extensive melting. Trace element signatures and
radiogenic isotopes show that these melts also interacted with the lithosphere (note the M 1 metasomes in
the path of the PAB primary melts in Fig. 19), acquiring
some of the hallmark chemical features of orogenic
magmatism, such as pronounced Ti–Nb–Ta depletion,
high LREE/HREE ratios and Pb enrichment (Fig. 7).
The chemical transition in the Afyon volcanic rocks
may reflect the interplay of thicker lithosphere in the
Afyon region and an increasing contribution of a deep
asthenospheric component owing to a slab tear within
the down-going slab. A similar asthenospheric component has also been identified for volcanoes located to
the south of the studied area, such as the Denizli (6 Ma)
and Isparta regions (<6 Ma), including the Quaternary
Gölcük volcano (Çoban & Flower, 2006, 2007; Prelević
et al., 2012; Platevoet et al., 2014). We interpret the
chemical transition in the rocks of the Late Afyon volcanic complex as the interplay of two processes, the
asthenospherization of the sub-Afyon lithospheric mantle and the ‘cleansing’ of previously metasomatized
contaminated lithospheric mantle. The asthenospherization is related to upwelling convecting mantle and is
spatially related to a slab tear. The ‘cleansing’ refers to
the preferred melting of metasomes in the lithospheric
mantle, which implies that this component is effectively
removed by early processes and is not subsequently
available for interaction with later melts. This ‘cleansing’ process is observed locally in many regions, such
as the Selendi (Ersoy et al., 2008) and Usak–Güre basins
(Karaoğlu et al., 2010) (Fig. 1). In these localities, igneous activity started with Early Miocene ultrapotassic
and high-K calc-alkaline products (Selendi andesites
and basaltic andesites and Güre lamproites), but shows
a more intense contribution of an Na-alkaline
555
component from the Middle Miocene onwards (Ersoy
et al., 2008). Eventually, volcanism was reactivated
(17 Ma to recent) as Na-alkaline volcanism in Kula (Alıcı
et al., 2002) with an OIB-like geochemical signature
resembling magmas derived from the convecting mantle; these are principally nepheline-bearing alkaline basalts and basanites, which clearly imply derivation from
a mantle free of orogenic influence.
THE ROLE OF SLAB TEARING IN THE ORIGIN OF
POST-COLLISIONAL MAGMATISM
The post-collisional geodynamic setting represents one
of the most complex environments for volcanism. This
complexity is reflected in the geochemistry of the
erupted magmas, which represents the end result of
multistage and multicomponent processes. Several
major stages are involved: processes related to the
hybridization between fluids or melts derived from a
subducting slab and the overlying mantle wedge during
subduction; tectonic imbrication of the lithosphere during collision; and, finally, post-collisional triggering of
the volcanism. The post-collisional magmatism represents the net effect of all these processes. Although volcanism in many post-collisional geodynamic settings
reflects the same sequence of events, local variations in
the extent and nature of interaction and the local character of the metasomatized mantle may result in very
heterogeneous volcanic products.
The tectonic trigger for widespread K-rich postcollisional volcanism in several orogenic belts, including the Tertiary Alpine–Himalayan belt and the Variscan
orogenic belt (e.g. Abdelfadil et al., 2014, and references
therein) is particularly controversial. In general, two
major types of tectonic trigger have been proposed for
this magmatism: lithosphere delamination associated
with orogenic collapse (Dewey, 1988) and tearing, segmentation and disruption of the subducted slab
(Rosenbaum & Mo, 2011).
We propose three major stages during the interaction between asthenosphere and lithosphere that
result from the tearing of the slab.
1. In the first stage, magma generation will be triggered
by heat from the upwelling asthenosphere, depending on the chemical state of the lithospheric mantle.
The most fertile parts of the lithospheric mantle will
be activated, resulting in extremely enriched melts
such as lamproites and lamprophyres. This process
also represents the beginning of the ‘cleansing’ of the
lithosphere and the elimination of ultra-fertile metasomes. Geochemically similar lavas would be produced by delaminating the lithospheric mantle and
orogenic collapse; therefore, this type of volcanism
cannot be diagnostic for the presence of a slab tear.
2. The second stage will reflect an increasing contribution of a deep asthenospheric component that becomes available because of a slab tear, resulting in
asthenosphere–lithosphere interaction. The role of
556
the asthenosphere may be twofold: its partial melting will result in Na-alkaline basaltic magmatism
and it may also serve as a source of precursor melts
of extreme composition. The depth at which this
happens will strongly affect the extent of melting,
the composition of the precursor melts and the nature of the asthenosphere–lithosphere interaction. In
other words, it is the lithosphere thickness that plays
a significant role in the character of these precursor
melts: in the case of lithosphere thickness greater
than 75 km, the asthenospheric melts may include a
dolomitic carbonatite component. Infilltration of the
lithosphere by such melts leads to wehrlitization and
to refertilization of the lithosphere, either by introduction of material from the asthenosphere or by
melting of the remaining metasomes in the lithosphere (‘cleansing’). This process largely resembles
the model of lithospheric mantle refertilization proposed by Thibault et al. (1992) and this type of volcanism should be diagnostic for the presence of a
slab tear.
3. In the third stage, volcanism eventually may be reactivated owing to regional extension. This volcanism
has a within plate chemical signature, resembling
magmas derived from the convecting mantle.
In conclusion, if we compare the above mechanism
that involves the tearing of the slab with other types of
tectonic trigger in post-collisional tectonic settings such
as orogenic collapse (Dewey, 1988; Elkins-Tanton,
2007), we may consider that the incursion of the asthenospheric mantle will be regionally recognized
though the clear geochemical signature typical for the
convecting mantle. A similar geochemical signal is not
expected when delamination of the lithospheric mantle
is the only process that represents the response to orogenic collapse. Moreover, fragmentation of the subducted lithospheric slab will induce mantle circulation
that can generate vigorous mantle upwelling, which
plays a more active role in the preconditioning of the
lithospheric mantle, producing freshly metasomatized
mantle domains that play an important role in the origin
of Si-undersaturated post-collisional lavas.
ACKNOWLEDGEMENTS
The paper benefited greatly from reviews by two anonymous reviewers; their considerable efforts have
been very much appreciated. We are also grateful to
Professor Marjorie Wilson for editorial handling of this
paper. D.P. acknowledges Endy DuFrane (University of
Alberta, Canada) for help with in situ Sr isotope measurements, and Sebastian Tappe for providing us with
IceRiver perovskite grain mount standard.
FUNDING
The study was supported by the Deutsche
Forschungsgemeinschaft (Grant PR 1072/1-1) and
Mainz University Research Fund to D.P.
Journal of Petrology, 2015, Vol. 56, No. 3
SUPPLEMENTARY DATA
Supplementary data for this paper are available at
Journal of Petrology online.
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