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Article

Petrography and Geochemistry of the Upper Cretaceous Volcaniclastic Deposits of the Haţeg Basin (Southern Carpathians): Inferences on Petrogenesis and Magma Origin

by
Violeta M. Vornicu
and
Ioan Seghedi
*
‘Sabba S. Ştefănescu’ Institute of Geodynamics, 19-23 Jean Louis Calderon, 020032 Bucharest, Romania
*
Author to whom correspondence should be addressed.
Minerals 2025, 15(2), 111; https://doi.org/10.3390/min15020111
Submission received: 23 December 2024 / Revised: 10 January 2025 / Accepted: 18 January 2025 / Published: 23 January 2025
(This article belongs to the Section Mineral Geochemistry and Geochronology)
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Abstract

:
Upper Cretaceous volcaniclastic deposits of the Haţeg Basin (VDHB) (Southern Carpathians, Romania) consist of relatively poorly exposed products of multiple phreatomagmatic volcanic eruptions of andesitic to rhyolitic composition and crop out around Densuş, Răchitova, Peşteniţa, and Ciula Mică localities. These deposits are commonly associated with the Late Cretaceous Neotethyan magmatic activity that developed in Central-Eastern Europe, forming the Apuseni–Banat–Timok–Srednogorie (ABTS) belt. Since the geochemistry of these deposits has been investigated very little so far, this study provides petrographic and whole-rock geochemical analysis for twenty new different volcaniclastic rock samples, out of which sixteen samples represent lava clasts and the other four are samples of pyroclastic flow deposits. According to our geochemical data, the VDHB have a calc-alkaline and high-K calc-alkaline character, similar to the majority of rock samples from all sectors of the ABTS belt. A comparison between the Haţeg rock samples and Banat and Apuseni samples reveals comparable major and trace element abundances and REE patterns, supporting the idea that they originate from similar magmas. Trace element patterns suggest that the parental magmas were mostly derived from the melting of a metasomatized lithospheric mantle source, previously modified by an earlier subduction event. A combination of crystal fractionation and variable degrees of crustal assimilation during storage at higher and lower pressures was the principal mechanism driving calc-alkaline differentiation. Our geochemical analyses indicate that the VDHB were produced by magmas generated during two different magmatic events. Older, silica-rich melts produced the Peştenita and Răchitova ignimbrite deposits, while the Densuş and Răchitova andesitic–dacitic–rhyolitic rock suite was generated by younger, intermediate magmas. The individual melt production episodes are evidenced by the emergence of two different crystal fractionation trends: an amphibole-controlled trend at mid-crustal levels and an upper-crust plagioclase-dominated trend. The hydrous, calc-alkaline magmas arguably occurred in a post-collisional setting, in agreement with the orogenic collapse model, among others, proposed for the origin of the ABTS magmatic activity.

1. Introduction

Upper Cretaceous volcaniclastic deposits are exposed in several locations along the western and north-western parts of the Haţeg Basin, Romania [1,2,3]. The most extended of these outcrops represent portions of former stratovolcano edifices made up of incomplete successions of volcaniclastic beds [4,5]. These volcaniclastic beds were produced by multiple explosive–effusive volcanic eruptions of mainly phreatomagmatic type during the Late Cretaceous [4,5,6].
The stratigraphic position of the volcaniclastic deposits with respect to other lithostratigraphic units and the lack of any volcanic vents anywhere in the Haţeg area or the surroundings suggest that the volcanic activity did not take place within the Haţeg Basin. Most probably, during the Late Cretaceous, the eruptive sequences have been detached from their original depositional area and transported to their current location through a series of tectonic movements [7].
Similar age, magma composition, tectonic history, and geographic proximity led to the hypothesis that the VDHB are related to the magmatism that manifested towards the end of the Late Cretaceous and led to the formation of the Apuseni–Banat–Timok–Srednogorie (ABTS) belt in Central–Eastern Europe [8], also known as the Banatitic Magmatic and Metallogenetic Belt (Figure 1 [9]).
The ABTS belt is approximately 1500 km long and 30–40 km wide, stretching from the Apuseni Mountains in north-western Romania to East Srednogorie in Bulgaria and further east to Turkey, Iran, and Afghanistan (Figure 1 [11,12]).
Based on compositional and tectonic criteria, the belt is divided into five major districts: Apuseni, Banat (Romania; Figure 2), Timok (Serbia), Panaghyuriste, and East Srednogorie (Bulgaria) (Figure 1) and consists of magmatic products both intrusive and extrusive, mostly large volcano–plutonic complexes, dykes, and volcanic formations of predominantly calc-alkaline and high potassium calc-alkaline to shoshonitic composition (Figure 2 [13,14,15,16]) and some mafic, more tholeiitic occurrences in the East Srednogorie segment [17,18]. All Upper Cretaceous magmatic rocks that constitute the ABTS belt are collectively known as “banatites” [19].
The general large-scale geodynamic framework responsible for the formation of the ABTS belt is related to the subduction of Neotethys or East Vardar oceanic basin under the European continental margin in the broader context of the collision between the European and African plates during the Late Cretaceous, forming the Vardar–Izmir–Ankara suture zone [22,23,24,25,26].
Although it is unanimously agreed that the convergence between Africa and Europe and the subduction of Mesozoic ocean basins during the Late Cretaceous set the initial conditions for the generation of the ABTS magmatic activity, there is still no consensus on what the exact mechanism that triggered melt formation was. In this regard, several geotectonic models have been developed in recent years, among which the most popular are orogenic collapse [9,27,28,29], slab rollback [15,30,31,32,33], and slab break-off [23,34].
Geochronological studies indicate that magmatic activity along the ABTS belt lasted roughly from 92 to 67 Ma [15,35,36]. In the Apuseni and Banat regions of Romania, the ABTS magmatism spanned for a period of more than 10 Ma, from 84 to around 71 Ma (Campanian). Recent zircon U-Pb and previous K-Ar dating of the VDHB yielded ages of about 80–82 Ma [6,37], which are consistent with the Late Cretaceous ABTS magmatism age range.
In spite of their frequent association with ABTS magmatism, the nature and source of the magmas that produced the Upper Cretaceous VDHB have been largely unexplored so far [7]. Another question that needs to be answered is how many magmatic events were responsible for the generation of these eruptive products.
Here, we present petrographic and whole-rock geochemical analyses for twenty new volcaniclastic rock samples of different compositions collected from Densuş, Răchitova, and Peşteniţa localities. From the available chemical data and the patterns emerging from these data, this study attempts to determine the origin and evolution of the parental melts, aspects that have not been investigated before. We also consider a possible correlation between the ABTS magmatism and the VDHB, and for this purpose, we compare our geochemical data obtained for the Haţeg rock samples with those for the Banat and Apuseni rock samples. Finally, we briefly discuss a potential geotectonic configuration that may have triggered the volcanism that produced the Upper Cretaceous eruptive products now emplaced within the Haţeg Basin.

2. Geological Background

Haţeg Basin is a syn-orogenic sedimentary depression in the Southern Carpathians of Western Romania (Figure 2 and Figure 3), bordered by several mountain chains such as the Retezat Mountains to the S and the Poiana Ruscă Mountains to the N and N-W. Its formation is thought to have been triggered by the collision between the Getic/Supragetic Units and the Danubian Unit after the consumption of the Severin Ocean, in the larger context of the convergence between the European and African plates during the Late Cretaceous [38].
For a significant period of its history, the area of the Southern Carpathians corresponding to the present-day Haţeg Basin was covered by the deep waters of the Tethys Ocean, as indicated by its sedimentary infill [29]. Marine sedimentation ceased towards the Late Maastrichtian when the extensional and/or trans-tensional tectonic regime determined by orogenic collapse [28] was replaced by uplift movements [39], ending with the installation of a continental environment [29].
The youngest marine deposits of the Haţeg Basin are represented by the Răchitova Formation that crops out in the W and N-W parts of the basin, with the stratotype in the locality of Răchitova (Figure 2 [40]). This unit comprises two individual lithological members—Lower and Upper [40,41]. The Lower Member, composed of mostly clay and marl deposits [40], was previously thought to represent the Late Campanian [41], but the latest palynostratigraphic studies revealed that it is younger, with an estimated age of early or early–late Campanian [42].
The marine deposits of the Răchitova Formation are overlain by Late Cretaceous continental sediments that crop out in multiple locations across the Haţeg Basin [1,43]. Based on compositional differences, the continental deposits exposed in the N-W part of the basin are grouped into the Densuş-Ciula Formation, while in the central and eastern parts they constitute the Sînpetru Formation (Figure 3 [3]). Densuş-Ciula Formation is further divided into three separate members: Lower, Middle, and Upper [3,44].
The Lower Member is exposed in N and N-W Haţeg and is represented by the Densuş and Răchitova volcaniclastic deposits, as proposed by previous authors [3,43]. However, recent studies suggest that the volcaniclastic deposits may represent a different lithostratigraphic unit [6].
The Middle Member crops out to the N and N-W of the basin and is predominantly made up of detrital sediments like sandstones, mudstones, and matrix-supported conglomerates, which also include some reworked volcanic material [45,46,47,48]. The Upper Member is mainly composed of coarse sediments, like sandstones, mudstones, red conglomerates, and siltites; unlike the other two members, it is entirely devoid of reworked volcanic material [3,43].
The Upper Cretaceous volcaniclastic deposits crop out in a discontinuous fashion in the north-western and northern parts of the Haţeg Basin, especially around Densuş and Răchitova localities (Figure 3 [1,2,3]), while outcrops of somewhat smaller extent occur near Peşteniţa and Ciula Mică localities (Figure 3 [1]). These deposits resulted from multiple mainly phreatomagmatic volcanic events, accompanied by some smaller-scale effusive activity [4,5].
Minerals 15 00111 g003
Figure 3. Geological map of the Haţeg Basin (modified after [49,50,51]); the location of the sampled volcaniclastic deposits is marked with a black star.
Figure 3. Geological map of the Haţeg Basin (modified after [49,50,51]); the location of the sampled volcaniclastic deposits is marked with a black star.
Minerals 15 00111 g003
The Densuş volcaniclastic deposits consist of two eruptive sequences of andesitic, dacitic, and rhyolitic composition, Densuş 1 (Figure 4a,a1,a2) and Densuş 2, located at the border between Ştei and Densuş villages. They are several tens of meters thick and represent small fragments of former edifices of composite volcanoes, each containing a succession of multiple primary, secondary, and epiclastic deposits [4,5]. In the vicinity of sequences Densuş 1 and Densuş 2, there are other similar volcaniclastic successions (Figure 3). These are, however, covered with soil and dense forest, making them difficult to study in the field.
The Răchitova volcaniclastic products (Figure 4b,b1,b2), positioned north of Densuş, resulted also from multiple phreatomagmatic eruptions of intermediate and rhyolitic magmas [5,7]. As in the case of the Densuş deposits, the Răchitova eruptive sequence contains multiple primary and secondary deposits and represents a piece of a former stratovolcano edifice [5].
The Peşteniţa volcanic deposits consist of relatively poorly exposed, isolated welded ignimbrite deposits of mostly intermediate to rhyolitic composition produced by highly explosive volcanism. Volcanic deposits of similar lithology and field exposure are also found near Ciula Mică, east of Răchitova [1].

3. Sample Selection and Methodology

The VDHB have relatively poor field exposures and, in general, exhibit signs of surface alteration and erosion due to weathering or diagenetic processes. For this study, we decided to use 20 volcaniclastic rock samples of different compositions collected from the Densuş, Răchitova, and Peşteniţa outcrops. Sample location and composition are given in Table 1. It should be mentioned that the Răchitova samples P6A and P6D, together with the two Peşteniţa samples P1 and P2 (Table 1), were collected from primary volcaniclastic deposits, specifically welded ignimbrites, while the rest of the samples represent volcanic rocks fragments formed directly by magma solidification as lavas. Due to the fact that they can sometimes contain foreign lithic clast and xenocrysts, pyroclastic rocks are not always suitable for geochemical analysis. Despite these drawbacks, the chemistry results we obtained for the ignimbrite rocks were successfully integrated in this study and, with few exceptions (see below), have proven reliable in determining magma evolution and origin.
The Densuş samples are partially altered, with a loss on ignition (LOI) ranging from 1.52 to 4.37, which indicates low to slightly moderate alteration. The LOI values of the Răchitova samples range from 0.24 to 9.54, showing low to high alteration. The two ignimbrite rock samples, P6A and P6D, display the highest LOI values of 8.47 and 9.54, respectively. The LOI values of the two Peşteniţa ignimbrite rock samples are 1.85 and 1.09, respectively, which translate to a relatively low degree of alteration. The LOI vs. SiO2 (wt%) plot is available as Supplementary Materials (Figure S1).
These data show that, at large, the selected rock samples are suitable for petrographic and geochemical analyses. However, the high LOI of samples P6A and P6D could reflect changes in the concentration of mobile elements such as Sr, Ba, or Pb. These changes were identified and are discussed in the following sections. The variation in K is most probably magmatic since it does not correlate with LOI (Supplementary File—Figure S2). Furthermore, K2O correlates well with large-ion lithophile elements (LILEs) such as U or Rb (Table 2).
Petrographic analysis of the rock samples was performed using an Olympus BX41 TF microscope (Oympus Corporation, Tokyo, Japan).
Whole-rock major and trace element compositions were determined at ALS Ireland (http://www.alsglobal.com/geochemistry (accessed on 28 July 2022 with certificate RM22173945). Powdered samples were submitted to lithium borate fusion followed by acid dissolution and then analyzed using inductively coupled plasma-atomic emission spectrometry (ICP-AES) for major elements (ALS code ME-ICP06) and inductively coupled plasma-mass spectrometry (ICP-MS) for trace elements including rare-earth elements (ALS codes ME-MS42, ME-MS61, ME-MS81). Weight loss on ignition (LOI) was determined gravimetrically after heating the rock powders to ~1000 °C using a thermogravimetric analyzer (TGA) (ALS code OA-GRA05).
Major and trace elements data are summarized in Table 2. The limit of detection (LOD) is given in the table. Major element values plotted on the diagrams were normalized to 100%.

4. Results

4.1. Petrographic Analysis

The selected rock samples are of intermediate to high-silica composition (Table 1 and Table 2), with porphyritic texture and aphanitic hypocrystalline groundmass (Figure 5A–F). The mineral assemblage of the Densuş and Răchitova andesite and dacite volcanic rock samples largely consist of euhedral and, more rarely, subhedral phenocrysts of plagioclase feldspar as the main mineral, followed by clinopyroxene, brown hornblende, and biotite, set in a groundmass composed of microcrystals, crystal fragments, and glass particles, sometimes with hyalophilitic texture (Figure 5A–C). Rhyolitic volcanic rock samples P4B and P7A are poorer in phenocrysts, and their mineralogy is dominated by potassium feldspar, plagioclase, and quartz set in a glass-rich groundmass (Figure 5D). Euhedral crystals of apatite, magnetite, ilmenite, zircon, and titanite commonly occur as accessory minerals.
Plagioclase crystals usually exhibit strong normal or patchy zonation, commonly with irregular cores. Hornblende is often partially or totally opacitized or, in some cases, is surrounded by microcrystals of opaque minerals (Figure 5C). Most of the time, hornblende crystals occur as glomerocrysts. Plagioclase is frequently altered to clay minerals, sericite, or epidote (Figure 5B), and in samples P4B, P5A, P6C, and P7A, crystals of this mineral are commonly fissured or fractured and impregnated with iron oxides, mostly along fissures. Mafic minerals, particularly pyroxene, are sometimes altered to chlorite or epidote. In some instances, pyroxene is replaced by hornblende. The transformation of hornblende into actinolite was also noticed is some of the samples. Other alteration minerals are celadonite and zeolites. In comparison with the rest of the samples, P4B, P5A, E5, and E6 have a more altered composition of both the phenocrysts and the groundmass. The mesostasis is sometimes altered to chlorite and carbonates.
The Răchitova and Peştenita ignimbrite samples, P1, P2, P6A, and P6D, are of trachytic and rhyolitic composition (Table 2) and largely share the same textural and mineralogical features. These samples exhibit apparent porphyritic texture, although not very pronounced, with a relatively low number of phenocrysts and xenocrysts, especially in the case of the Peşteniţa samples, set in a hypohyaline groundmass. The groundmass is mostly composed of flattened glass shards with angular, blocky, or cusp-like shapes, as well as crystal fragments, displaying a fluidal pattern and eutaxitic texture (Figure 5E,F). The main mineral phases are alkali feldspar, plagioclase feldspar, and quartz, with subordinate biotite and brown hornblende, which mostly occur as euhedral crystals (Figure 5E,F). The phenocrysts are often fractured or fragmented due to strong magma fragmentation. The xenocrysts are rare and are biotite and quartz. They can be distinguished from the phenocrysts by their heavily deformed state, which forms severe kink bands in biotite and undulatory extinction in quartz. The four samples also contain a small amount of volcanic and accessory metamorphic lithic clasts (Figure 5E,F).
Alkali feldspar crystals found in the Peşteniţa samples are often transformed to kaolinite, while those found in the two Răchitova samples, P6A and P6D, are frequently replaced by epidote or clay minerals. The plagioclase feldspar is, at times, sericitezed, and the hornblende is usually replaced by magnetite and/or pyroxene. Likewise, alteration of biotite crystals to chlorite can sometimes be observed. Small fissures or cracks that cut through the groundmass of the Peşteniţa samples are usually filled with calcite or quartz. Spherulitic aggregates can be observed in the Răchitova ignimbrite samples P6A and P6D.

4.2. Whole-Rock Geochemistry

Whole-rock major and trace element results obtained for the Densuş, Răchitova, and Peşteniţa rock samples are provided in Table 2. For comparison, on the major element diagrams, our rock samples were plotted together with the samples from [7] collected from Densuş and Răchitova (Supplementary File—Table S1). Whole-rock chemical composition of rhyolitic volcanic rocks and ignimbrite deposits from the Haţeg Basin has never been investigated before and is, therefore, provided for the first time in this study.
The Densuş rock samples are mostly intermediate in composition with respect to the SiO2 content, which varies between 54.5 and 69.2 wt% (Table 2). The Răchitova rock samples are intermediate and rhyolitic, with a SiO2 content in the range of 59.3–75.5 wt% (Table 2). The Peşteniţa rock samples are silica-rich, with a SiO2 content of 65.4 and 67.8 wt%, respectively (Table 2). The samples have low to moderate whole-rock Mg number (Mg#) in the range of 7.9–46.6 (Table 2).
According to the alkali vs. SiO2 (TAS) diagram (Figure 6 [52]), the Densuş samples fall into the categories of andesites, trachy-andesites, and dacites, except for the P4B sample, which is a rhyolite. The Răchitova rock samples are trachy-andesites, trachy-dacites, and rhyolites, while the two Peştenita rock samples, P1 and P2, fall into the category of trachytes.
On the K2O vs. SiO2 wt% diagram (Figure 7 [53]), the volcaniclastic rock samples plot into the calc-alkaline and high K-calc-alkaline fields are very slightly shifted toward less evolved compositions, compared to the TAS diagram (Figure 6). This is due to the fact that the ranges of SiO2 content employed for this diagram are different from the ones used for the TAS diagram. In this study, we use the TAS classification of [52]. On the SiO2 vs. Zr/TiO2 and Zr/TiO2 vs. Nb/Y diagrams [54], some of the samples plot in different compositional fields due to high igneous differentiation, which led to modifications of the concentration of some of the immobile trace elements such as Nb (Figure 8a,b).
On the chondrite-normalized rare earth elements plot (Figure 9a,b [55]), the volcaniclastic rock samples have slightly different concentrations of rare earth elements (REEs). All samples show enrichment in light rare earth elements (LREE), such as La, Ce, and Nd, relative to heavy rare earth elements (HREE), such as Dy, Er, and Yb. The rock samples with SiO2 > 64% P4B (Densuş), P6A, P6D, P7A (Răchitova), and P1 and P2 (Peşteniţa) are somewhat more enriched in REEs. They also exhibit negative Eu anomalies, although not very prominent (Figure 9b). Eu anomalies (Eu/Eu*) values are in the range 0.58–0.96 (Supplementary File—Table S2, Figure S3). In the case of these samples, two different trends can be seen emerging on the chondrite-normalized rare earth elements plot—samples P7A and P4B have similar trace element concentrations and show weak Eu anomaly, while ignimbrite samples P6A, P6D, P1, and P2 have higher REE abundances and display more pronounced negative Eu anomalies (Figure 9b). It can be noticed that the trend followed by samples P7A and P4B is identical to that of the samples of intermediate composition (Figure 9a). All samples display moderate to high LREE enrichment, with La/Yb ranging from 9.33 to 24.91 and an average value of 15.83 (Supplementary File—Table S3).
On the primitive mantle-normalized trace element plot (Figure 10a,b [55]), all samples are enriched in large-ion lithophile elements (LILEs), including Ba, K, Sr, and U and Th, and depleted in high-strength field elements (HFSEs). A high abundance of mobile incompatible elements such as Pb, Ba, and LREE is also a common feature of the rock samples.
On the SiO2 vs. major oxides diagrams, Al2O3, TiO2, Fe2O3, CaO, and MgO decrease with increasing SiO2 in all samples (Figure 11a–f). In contrast, Na2O shows a positive correlation with SiO2 in the intermediate samples and tends to decrease in the more high-silica samples, whereas K2O remains more or less constant but increases in the silica-rich samples (Figure 11g,h).
In comparison with the rest of the samples, the Răchitova and Peşteniţa ignimbrite samples P1, P2, P6A, and P6D have a higher content of some of the immobile incompatible elements, such as Zr, Y, and the LREEs, as shown on the SiO2 vs. selected trace element diagrams (Figure 12a–f). On the SiO2 vs. Nb diagram, relative to the other samples, the Răchitova rhyolite sample P7A displays a higher Nb concentration (Figure 12c).
The Ba content does not show a strong variation in andesitic and dacitic rock samples; however, it rapidly increases in more silica-rich samples (Table 2). Compared to the rest of the rock samples, samples P6A and P6D have much higher Sr concentrations, while the Peşteniţa samples P1 and P2 are richer in Rb (Figure 12g). The K/Rb of the samples varies between 0.01 and 0.05; the Ba/Rb is in the range of 4.25–27.9, while the Rb/Sr ranges from 0.02 to 1.58. For the majority of the volcaniclastic rocks samples, the Sr/Y ranges from 15.2 to 36.11 (Supplementary File—Table S3). Samples P1 and P2 show very low Sr/Y values, of 4.45 and 3.31, whereas samples P6A and P6D exhibit the highest ratios, of 63.52 and 74.19, respectively (Supplementary File—Table S3).

5. Discussion

5.1. Geobarometry

Applying the empirical Al-in-hornblende geobarometer proposed by [57], we tried to estimate the equilibrium pressure of magmas; that is, the depth of the magma chamber where the phenocrysts where last in equilibrium with the melt. For this purpose, we used the hornblende aluminum content (TAl) data for several andesitic rock samples collected from the Răchitova and Densuş areas, provided by [7] (Supplementary File—Table S4). Calculations yielded a wide range of pressures, between 1.1 and 4.3 kbar (±0.5 kb) corresponding to a depth range of roughly 4.2–16 km, assuming a lithostatic pressure gradient ~3.7 km/kbar (Supplementary File—Table S4) These results indicate the existence of two different magma crystallization levels, from shallow crustal regions corresponding to a depth range of 4–8 km to mid-crustal level reservoirs found at depths of about 9–16 km.

5.2. Open System Igneous Differentiation

5.2.1. Crystal Fractionation

Volcaniclastic rock samples of the Haţeg Basin display an evolved magma profile, with a fractionating trend ranging from 54 to 75 wt% SiO2 (Table 2) and a REE pattern with an average La/Yb value of 15.83 (Supplementary File—Table S3), suggesting that the primary composition was modified by various differentiation processes.
The samples that show the lowest degree of differentiation and are, therefore, closest to the parental magma composition are P4A and P5A, with the lowest SiO2 and highest MgO content and were collected from the Densuş area, whereas samples P6A, P6D, and P7A have amongst the most differentiated compositions and belong to the Răchitova volcaniclastic deposits (Table 2). In comparison with the chemical compositions of primitive, near-primary arc lava samples [58,59], the least evolved samples mentioned above have much lower MgO and FeO concentrations and higher SiO2, meaning that the parental magmas do not reflect the starting composition of primary magmas.
A different indicator of the degree of magma differentiation is the Mg#. Assuming that mantle peridotite-derived magmas have Mg# > 70 [59], the highly variable values obtained for the Haţeg rock samples, between 7 and 46 (Table 2), tell us that they could not have formed directly from parental magmas and suggest different degrees of differentiation.
The linear decrease in Al2O3 and CaO with SiO2 content (Figure 11a,c), together with the occurrence of negative Eu anomalies (Figure 10a,b) and low-to-moderate Sr/Y (<40) (Supplementary Data—Table S3), are indicators of plagioclase fractionation. However, judging from the rather mild Eu anomalies in some samples (Figure 10a,b; Supplementary File—Figure S3), plagioclase was probably not the most important fractionating phase in the majority of the rock samples. The negative correlation of Fe2O3, CaO, and MgO with SiO2 indicates hornblende, clinopyroxene, and magnetite fractionation (Figure 11b–d). Decreasing TiO2 with increasing SiO2 (Figure 11e) is usually attained by fractionation of titanium-bearing minerals, such as titanite, ilmenite, rutile, etc., which is also confirmed by the observed negative Ti and Nb anomalies (Figure 10a,b). The positive correlation of K2O and Rb with SiO2 (Figure 11g and Figure 12g) indicates that the crystallization of K-feldspar and biotite did not play an important role during the incipient stages of mineral fractionation. P2O5 increases with differentiation up to about 60 wt% SiO2, then starts to drop with a further increase in silica, forming a kink on the diagram (Figure 11f). This pattern is generated by the saturation of apatite, also indicated by the conspicuous negative P anomaly. Zr follows a similar trend, although zircon saturates at higher SiO2 concentrations of about 65 wt%. This behavior of P and Zr demonstrates that crystal fractionation was the principal mechanism driving magma differentiation [60].
On the Rb vs. Co diagram (Figure 13), a smooth decrease in Co, highly compatible in the mantle, with increasing Rb, which behaves incompatibly, is observed. This pattern further suggests that silica enrichment was mostly achieved by crystal fractionation.
Most of the analyzed samples show low to moderate Sr/Y values, in the range of 15–36.5 (Supplementary File—Table S3), suggesting predominant amphibole fractionation since Y is commonly retained in this mineral but also in garnet [61]. Equally, middle and heavy REE are also preferably incorporated into amphibole [62,63,64]. The observed decrease in (Dy/Yb)ₙ and increase in (La/Sm)n with differentiation also suggests amphibole fractionation (Supplementary File—Figures S4 and S5 [65]). On the contrary, garnet fractionation causes an increase in the Dy/Yb, as garnet assimilates only into the HREEs, producing flatter fractionation trends [63,65,66,67]. In light of these considerations, we can discard a garnet-rich mantle as a possible source of the parental melts. Amphibole fractionation is perhaps best observed on the Dy/Dy*vs. Dy/Yb plot (Figure 14), devised by [67]. Dy/Dy* is a proxy conceived by interpolation between LREE and HREE (La and Yb) and is used to measure the curvature of the REE pattern, which is finely controlled by mineral behavior. Amphibole and clinopyroxene fractionation decrease Dy/Dy* and Dy/Yb, while garnet increases Dy/Yb with a limited increase or decrease in Dy/Dy*.
Amphibole reaches thermal stability in the lower to mid-crustal levels, at higher pressures, between approximately 0.8 and 2 GPa (~25–60 km), and a high volatile content in the melt, but fractionation can also occur in the upper crust, at lower pressures [69,70]. This type of fractionating environment is usually associated with subduction or subduction-related settings, where highly hydrous magmas (>5% H2O) are usually produced and stored in the lower to mid-crustal levels [65,71,72]. Plagioclase crystallization is normally inhibited in volatile-rich melts [73,74,75,76], but this process can still take place in these conditions at lower pressures (<0.7 GPa) that are only achieved in mid-to-upper crustal levels [77].
Compared with the rest of the samples, a much more elevated Sr content was detected in ignimbrite samples P6A and P6D, of 2160 and 2070 ppm, respectively (Figure 12f). These two samples also show a relatively high Y content of 34 and 27.9 ppm. In contrast, the Peşteniţa samples P1 and P2 have the lowest Sr/Y. Since Sr has a high distribution coefficient in plagioclase, the low Sr/Y can be explained by plagioclase fractionation, which drives Sr depletion in magmas and suppresses amphibole crystallization [78]. Despite the elevated Sr content of samples P6A and P6D, the high Y concentration that correlates positively with SiO2 and the presence of negative Eu anomalies are indicators of plagioclase as a principal fractionation phase. Therefore, Sr enrichment must have happened after crystal fractionation (see below).
From the above analysis, we can recognize two individual differentiation trends—an amphibole—dominated trend of water-rich magmas in the mid-crustal levels that describe the Densuş and Răchitova andesite–dacite–rhyolite magma suite (samples D2-1, D2-2, D2-3, E2-E7, P4A, P4B, P5A, P5B, P6B, P6C), and a plagioclase-controlled trend that took place in the upper crust and determined the final composition of the Răchitova and Peşteniţa high-silica magmas that produced the ignimbrite deposits (samples P6A, P6D, P1, and P2). Trace element behavior, such as that of Dy/Yb and La/Sm, indicates that amphibole dominated the fractionating process, contributing to the calc-alkaline differentiation of all samples, regardless of their final composition.
To further test this result, we employed the Y vs. CaO plot, proposed by [79], who came up with two separate differentiation trends, depending on the enrichment or depletion of Y with respect to a standard calc-alkaline trend: a J-type arising from predominant hornblende fractionation and an L-type, defined by plagioclase and pyroxene crystallization. As shown in this diagram, the Densuş and Răchitova andesitic–dacitic–rhyolitic suite samples follow the J-type or hornblenditic trend, while the Răchitova and Peşteniţa ignimbrite samples conform to the L-type or plagioclase/pyroxenitic trend (Figure 15a).
Similarly, on the Yb vs. La/Yb plot, which is commonly used for discriminating between normal arc and adakite or adakite-like magmas, the Haţeg rock samples follow two distinct fractionating trends—an amphibole and a plagioclase fractionating trend (Figure 15b). In this diagram, sample P7A plots in the ‘adakite-like’ field due to its high La/Yb > 20, Sr/Y ~20, and low Y < 18 ppm and Yb < 1.9 ppm (Table 2 [80,81]). On the Y vs. Sr/Y diagram, the samples follow the same two differentiation paths, with few of the samples plotting in the ‘adakite’-like field. However, P6A and P6B diverge from the rest of the samples because of the alteration-induced high Sr concentration (Supplementary File—Figure S6). In agreement with the model proposed by [82] for the formation of ‘adakite’-like magmas identified in the East-Serbian segment of the ABTS belt, these compositions may form through a combination of abundant amphibole fractionation in the lower crust and plagioclase-fractionation and crustal assimilation in the upper crustal levels.

5.2.2. Crustal Assimilation

Crustal assimilation is now long recognized as one of the most important processes that play out during igneous evolution, particularly in continental arc or arc-related regions, and can significantly shift the chemical composition of magmas [83,84,85,86].
The high Sr content in the two Răchitova ignimbrite samples, P6A and P6D, could have come from crustal assimilation. However, for a rock sample (RA-12) collected from the same Răchitova ignimbrite deposit that samples P6A and P6D were collected from, [87] reported an 87Sr/86Sr value of 0.704749 ± 0.000009 and a δ 18O of 6.5 (biotite) and 11.9 (plagioclase). These isotopic ratios suggest that crustal assimilation cannot entirely account for the anomalously high Sr concentration in these samples, so it was most probably acquired from post-depositional alteration processes, as suggested by the high LOI values. Sr enrichment in these samples correlates well with the fluid-mobile elements such as Ba (Table 2).
Ref. [87] determined the 87Sr/86Sr and δ 18O for six additional volcaniclastic rock samples (RA-1, RA-2, RA-6, RA-7, RA-8, and RA-13) collected from the andesitic–dacitic–rhyolitic unit of the same Răchitova outcrop, which lies atop the rhyolitic ignimbrite that we mentioned earlier, and also for a rock sample (AVROM-204) collected from a tuff layer that crops out around Ciula Mică locality. The 87Sr/86Sr of these samples is in the range 0.705423–0.708061, and the δ 18O varies between 6.27 and 15.93 (Supplementary File—Table S5). Relative to the mantle 87Sr/86Sr, which can vary between 0.702 and 0.704 [88,89,90], the analyzed samples display higher values. The authors of [87] suggested that these results are indicative of crystal fractionation as the main magma differentiation process, accompanied by varying degrees of crustal assimilation, ranging from 0.5 to 1%. They identified two different trends in the samples: one trend that arises from the assimilation of a crustal material with low 87Sr/86Sr and high δ 18O values and a second trend resulting from the assimilation of rocks with a higher 87Sr/86Sr and moderate δ 18O.
The 87Sr/86Sr of the Răchitova volcaniclastic deposits published by [87] are, at large, consistent with the isotopic data for Banat rock samples, which range from 0.704421 to 0.707254 [14,90]. The 87Sr/86Sr of Apuseni rock samples are generally higher, between 0.704484 and 0.731214 [16,90]. This difference is not surprising given that magmatic bodies of the Banat and Apuseni regions were intruded into two different crustal mega-units: Dacia and Tisza.
The Rb/Th values of the rock samples can be used for a rough approximation of the degree of crustal assimilation since these elements are not substantially affected by fractional crystallization [91]. Rb is incorporated in biotite, but this mineral is relatively rare in the sampled volcaniclastic rocks, found only in some silica-rich dacites and rhyolites. For the Densuş and Răchitova andesitic–dacitic–rhyolitic rock suite, the Rb/Th is in the range of 1.2–13.8, pointing to variable degrees of crustal assimilation, from low to high, whereas for the Răchitova and Peşteniţa ignimbrite samples, it is between 3.6 and 9.7, recording a moderate degree of assimilated crustal material (Supplementary File—Table S3). The Rb/Th vs. SiO2 diagram is provided as Supplementary Materials (Figure S7).
The discovery of inherited zircon crystals from the crystalline basement in two of the Densuş sampled volcaniclastic rocks, D2-1 (143.68 Ma, 150.59 Ma, and 157.42 Ma) and D2-2 (between 2.54 and 2.8 Ga) [6], further suggests incorporation of the local crust. The author of [90] outlines the presence of inherited zircons in igneous rocks from the Banat and Apuseni segments of the ABTS belt, of Jurassic–Triassic, Carboniferous, Ordovician, Neoproterozoic, and even older crystallization ages (1–2 Ga).
In summary, the elements discussed so far lead to the conclusion that magma differentiation was mostly driven by combined crystal fractionation and crustal assimilation (AFC) in various proportions. However, in accordance with the AFC modeling results obtained by [16], we assume that more than 1% added crustal material is required to explain the composition of the VDHB, especially in the case of high-silica rhyolites.

5.3. Potential Magma Source and Geodynamic Implications

The magmas that produced the eruptive products of the Haţeg Basin have all the characteristics of subduction-related magmas belonging to the calc-alkaline magma suite (Figure 7). Calk-alkaline melts typically form in volcanic arcs above convergent plate boundaries by flux melting [92,93,94,95,96], but their occurrence is not restricted to active subduction zones. This type of magma can also be generated in post-collisional settings as a consequence of lithospheric extension by decompression melting of the mantle wedge or lithosphere enriched by a previous subduction event [97,98,99,100,101].
The calc-alkaline and high-K calc-alkaline character of the Haţeg volcaniclastic rock samples is testified by enrichment in LILEs relative to the HFSEs, enrichment in LREEs and depletion in HREEs, a high concentration of mobile elements (e.g., Pb, Ba), as well as the occurrence of strong negative Nb and Ta anomalies (Figure 9 and Figure 10 [102,103,104,105]).
As far as the VDHB are concerned, the precise location where volcanism manifested remains largely unknown, but according to paleomagnetism measurements conducted in the Haţeg Basin, it must have been situated closer to the Equator, in a sub-tropical environment, coinciding to a large extent with the initial location of the magmatic products of the Upper Cretaceous ABTS belt [106,107,108].
During the Late Cretaceous post-collisional stage of the ABTS belt, several sedimentary ‘Gosau-type’ basins were formed as a consequence of crustal-scale extension, among which was the Haţeg Basin [28,109]. ‘Gosau-type’ basins are mostly found in Austria, Germany, and Slovakia and expose a type of stratigraphic group containing largely marine deposits, whose age spans from Late Cretaceous to Eocene [110]. Considering the fact that the volcaniclastic deposits are at present emplaced within a ‘Gosau—type’ basin, postdating the formation of the basin, we can argue that the Late Cretaceous calc-alkaline volcanic activity likely developed in a post-collisional setting, under an extensional regime, rather than under active subduction conditions. This view is consistent with the orogenic collapse model proposed, for example, by [9] or [28] for the origin of ABTS magmatism. According to this geodynamic model, lithospheric thinning triggered the initiation of magmatism during the maximum extensional phase by partial melting of the upper lithospheric mantle. Collapse of the orogen along the ABTS belt and the installation of extensional or trans-tensional tectonics could have been initiated by slab rollback, even after slab subduction came to a complete halt [9] or, possibly, by slab tear or break-off [34].
Compositional characteristics suggest that magmas were produced mainly by partial melting of a sub-continental lithospheric mantle (see below) enriched during a previous subduction event, probably generating fractionated lherzolitic-type melts [111,112]. Recent studies have shown that in subduction-related environments, an important volume of magma can also be generated by pyroxenite melting [113]. Crystallization of hydrous minerals phases, such as amphibole, attests that the parental magmas were water-rich, oxidized melts derived from a metasomatized mantle source [72,114]. Formation of high-K calc-alkaline magmas, which are commonly generated in post-collisional settings [115], takes place within a mantle source region that is stable up to about 100 km [116], so melt generation likely took place at around this depth. The observed REE patterns are compatible with the presence of spinel in the mantle source and the absence of garnet [117].
Enrichment of the mantle source is also evidenced by the trace element ratios Ce/Pb and Nb/U [54,118]. For normal mid-ocean ridge basalt (N-MORB) and ocean island basalt (OIB), the Ce/Pb is ~25, and the Nb/U is ~47, whereas in island arc magmas and the continental crust, the values are usually <10 [119]. The Ce/Pb ratio is low even for the least evolved Haţeg samples (<5), whilst the Nb/U varies between 2.15 and 7.23 (Supplementary File—Table S3). These values are often observed in continental arcs or the crust [119], meaning that the parental magmas must have been produced by an enriched source akin to that found in subduction-related settings.
Further constraints on melt origin come from the Ta/Yb vs. Th/Yb and Nb/Y vs. Th/Y diagrams [120]. Compared to the MORB and the OIB, the Haţeg calc-alkaline magmas exhibit higher Th/Y values, which is suggestive of a metasomatized mantle component enriched in trace elements that abound in the crust (Figure 16a [120]). On the Nb/Y vs. Th/Y plot, all Haţeg rock samples adhere to one pattern, except for sample P7A, which plots in a different field due to its highly evolved composition (Figure 16a). On the Ta/Yb vs. Th/Yb diagram, the samples follow a similar trend (Supplementary File—Figure S8). The peculiar behavior of sample P7A is dictated by its high Nb, Ta, and Th concentrations. The ionic radius of Zr is similar to that of Nb and Ta, so these can easily replace Zr in zircon. This way, zircon crystals can concentrate significant amounts of Nb and Ta in their lattices [121,122]. Therefore, enrichment in these elements of sample P7A could have been produced by extensive assimilation of zircon-rich crustal rocks [123,124].
The authors of [126] suggested that certain trace element ratios, such as Nb/La and La/Yb, can be useful in tracing down different magma sources. High Nb/La ratios (~>1) indicate an OIB-like, asthenospheric mantle source, while lower ratios (<0.5) are common for lithospheric mantle regions. In the case of the Haţeg rock samples, the Nb/La is lower than 0.5, except for sample P7A, which has a slightly higher value of 0.73, while the La/Yb is in the range 9.33–24.9 (Supplementary File—Table S3). On the Nb/La vs. La/Yb diagram, the samples plot in the lithospheric mantle field (Figure 16b).
A possible explanation for the formation of high-silica rhyolite magmas that produced the Răchitova and Peşteniţa welded ignimbrite deposits and the other silica-rich rocks, is a polybaric evolution involving multiple stages of magma differentiation. At first, mantle-derived mafic melts percolated into and stalled at mid-crustal regions (>500 MPa), where they underwent mineral fractionation and crustal assimilation, forming crystalline mushes of evolved compositions [60,127]. Interstitial melts were then extracted and rose up to the upper crust (<8 km/250 MPs), where they further differentiated in an open system through a mixture of plagioclase-dominated crystal fractionation and crustal assimilation, finally producing wet rhyolitic magmas, in accordance with the model advanced by authors such as [128] or [129]. This model is compatible with the fractional crystallization paths discussed above and can explain the major influence of amphibole fractionation in the evolution of all rock samples. Additionally, it correlates well with the δ 18O values of [87], which range from 6.27 to 15.93.

5.4. How Many Magmatic Events?

Compositionally, the Densuş volcaniclastic successions are mostly intermediate with subordinate rhyolitic products (Figure 6). The Peşteniţa ignimbrite deposits are primarily of intermediate (trachytic) composition (Figure 6). The lowermost part of the Răchitova eruptive sequence is made up of a high-silica rhyolitic ignimbrite layer (samples P6A and P6D), which is unconformably overlain by a younger succession of andesitic, dacitic, and less often, rhyolitic volcaniclastic beds (Figure 6). A detailed stratigraphic description of the Răchitova and Densuş volcaniclastic successions is provided by [5].
The Densuş and Răchitova andesitic–dacitic–rhyolitic rock suite share the same geochemical profile, with similar major and trace element concentrations, following an identical pattern on the chondrite-normalized rare earth elements diagram (Figure 9a,b [10]). Moreover, the eruptive products from the two locations also have the same age of 80–82 Ma [6,37].
As shown in previous sections, the final composition of the magmas that fed the Upper Cretaceous calc-alkaline stratovolcanoes resulted from two different igneous differentiation trends. The Densuş and Răchitova andesitic–dacitic–rhyolitic magma suite was produced by a mid-crustal level amphibole-controlled differentiation trend, whereas the evolution of the silica-rich melts that generated the Răchitova and Peşteniţa ignimbrite deposits was governed by an upper crust plagioclase-controlled differentiation trend. The occurrence of two distinct REE trends on the chondrite-normalized REE diagram is concordant with this result (Figure 9a,b). Furthermore, on the SiO2 vs. incompatible immobile element variation diagrams, the concentration of Nb, Y, and Zr and the rare earth elements La, Nd, and Yb is notably higher in the four ignimbrite samples P1, P2, P6A, and P6D (Figure 12a–f). These samples stand out, following a trend that is markedly different from that of the rest of the samples, which is an important indicator of a distinct magma batch that evolved separately under different crustal conditions.
As per the data discussed above, we infer that the Upper Cretaceous VDHB were produced by multiple volcanic centers belonging to the same volcanic field that were formed during at least two independent magmatic episodes. During the older magmatic event, silica-rich rhyolitic magmas produced the Peşteniţa and Răchitova ignimbrite rocks, while a second, younger event generated mostly intermediate melts that formed the Densuş and Răchitova andesitic, dacitic and rhyolitic eruptive sequences. Certain differences in composition and texture, as shown by petrographic analyses (Figure 5E–F), as well as the different outcrop locations, indicate that the Peşteniţa and Răchitova welded ignimbrite deposits are the products of different volcanic centers. Since all volcaniclastic rock samples used in this study broadly display the same chemical profile, we interpret that they evolved from parental melts that largely shared the same compositional characteristics.

5.5. Correlation with the Late Cretaceous ABTS Subduction-Related Magmatism

Here, we compare the Haţeg volcaniclastic rock samples’ major and trace element results, including those of [7], with the geochemical data published by [14,15,16] for several intrusive/extrusive rock samples collected from the Banat and Apuseni regions, representing the Romanian sectors of the Upper Cretaceous ABTS belt.
All samples show comparable major element concentrations; however, some Apuseni and Banat samples have higher FeO and MgO concentrations, displaying a more basic composition [14,15,16]. The Haţeg Basin volcaniclastic rock specimens are somewhat richer in K2O and Na2O, whereas the Banat and Apuseni samples show a tendency towards higher MnO and lower TiO2 [14,15,16].
Extrusive rock samples from Banat and Apuseni cover a wide range of compositions, from basalts to rhyolites, but the majority have SiO2 contents greater than 60 wt%, corresponding to dacites and high-silica rhyolites [14,15,16].
On the K2O vs. SiO2 wt% diagram [53], most of the Apuseni and Banat samples fall into the field of calc-alkaline and high-K calc-alkaline magmas, with few samples plotting into the shoshonite field (Supplementary File—Figure S9). Likewise, the Hațeg rock samples are mostly of calc-alkaline and high-K calc-alkaline composition (Figure 7).
On the chondrite-normalized rare Earth elements diagram [55], the rock samples from all three regions follow a similar REE pattern, with enrichment in LREEs and depletion in HREEs (Supplementary File—Figure S10). Some of the ABTS rock samples display a flatter REE trend, with a slightly lower LREE content relative to the HREEs. In addition, larger Eu anomalies occur in some of the ABTS samples.
On the primitive mantle-normalized trace element diagram [55], the Apuseni/Banat and Haţeg samples overlap, displaying a general comparable pattern, with enrichment in LILEs relative to HFSEs, positive Pb and Rb anomalies and strong negative Ti, Nb, and P anomalies (Supplementary File—Figure S11). It is noticed, though, that, in comparison with the Haţeg samples, some of the Apuseni and Banat samples display a more pronounced positive Pb anomaly and are poorer in Ba and some of the LREEs, with more pronounced negative La, Ce, or Nd anomalies.
Altogether, our whole-rock major and trace element results are in good agreement with the geochemical data for ABTS rock samples collected from the Banat and Apuseni sectors, which are similar in major and trace element composition. As pointed out above, the isotopic data for the Răchitova eruptive products are concordat with those for Banat magmatic products [14,79,90]. Furthermore, we find a good correlation between the ABTS magma evolution and the differentiation of magmas that generated the VDHB [14,16,90]. This compatibility in terms of magma composition and evolution strongly supports the view that the Upper Cretaceous eruptive products of the Haţeg Basin are genetically related to the Late Cretaceous ABTS magmatism. This implies that the predominantly explosive volcanism developed from the same subduction-modified mantle source as the entire Neotethyan magmatism. Neo As stated previously, this hypothesis is further supported by similar ages and tectonic history, as well as geographic proximity.

6. Concluding Remarks

The Upper Cretaceous VDHB covers a wide range of compositions, from basaltic andesites to silica-rich rhyolites, and has calc-alkaline affinity, entailing that their parental magmas were generated in a subduction-related or subduction-modified tectonic setting.
Because of the way they are formed, pyroclastic rocks can sometimes contain foreign volcanic or non-volcanic material, so chemical analyses of this rock type can be influenced to a certain degree by these inclusions. However, as shown in this study, the composition of the analyzed ignimbrite samples is largely magmatic and has been successfully used to determine the properties of magmas and their evolution.
Major and trace element results indicate that the composition of the eruptive products of the Haţeg Basin was essentially produced by fractional crystallization of mantle-derived, hydrous melts, to which variable amounts of crustal material were added. Trace element patterns revealed the existence of two different igneous differentiation trends, suggesting distinct but closely genetically related magmatic events that supplied multiple volcanic centers during the Campanian age. Amphibole was, overall, a primary mineral phase controlling the fractionation process, which is characteristic of the evolution of subduction or subduction-related melts.
Our analysis has shown that the parental magmas were largely derived from partial melting of the lithospheric mantle. The mantle source composition was modified by a pre-subduction event, producing volatile-rich calc-alkaline and high-K calc-alkaline melts.
Major and trace element geochemical data overlap with previously reported geochemical analyses of rock samples from the Banat and Apuseni segments of the Upper Cretaceous ABTS belt, as well as with samples from the other segments of the belt, representing compelling evidence in favor of a common origin of magmas.
In view of a probable genetic relationship between the Late Cretaceous Banatitic magmatism and the volcaniclastic deposits of the Haţeg Basin, we propose that the magmatic activity likely developed in a post-collisional setting, set off by lithospheric mantle decompression melting.

Supplementary Materials

The following supporting information can be downloaded at: https://www.mdpi.com/article/10.3390/min15020111/s1, Figure S1: Loss on ignition (LOI) vs. SiO2 (wt%) of the Hateg rock samples (this study); there is no clear correlation between the LOI variation and SiO2 content, although a slight increase of LOI with decreasing SiO2 is observed; compared to the other rock samples, ignimbrite samples P6A and P6D display much higher LOI values. For legend see Figure 6. Figure S2: Loss on ignition (LOI) vs. K2O (wt%) of the Haţeg rock samples (this study); the LOI variation does not correlate with SiO2 content. For legend see Figure 6. Figure S3: Eu/Eu* vs. SiO2 (wt%) plot of the Haţeg volcaniclastic rock samples (this study). For legend see Figure 6. Figure S4: (Dy/Yb)ₙ vs. SiO2 (wt%) plot of the Haţeg volcaniclastic rock samples (this study); (Dy/Yb)ₙ was normalized to the chondrite values of [55]. For legend see Figure 6. Figure S5: (La/Sm)ₙ vs. SiO2 (wt%) plot of the Haţeg volcaniclastic rock samples (this study); (La/Sm)ₙ was normalized to the chondrite values of [55]. For legend see Figure 6. Figure S6: Y vs. Sr/Y (ppm) plot for the Haţeg volcaniclastic rock samples (this study). Fields for ‘adakite’-like and calc-alkaline magma compositions from are from [80] and mineral differentiation paths are from [81]. Abbreviations: amph—amphibole, plag—plagioclase. For legend see Figure 6. Figure S7: Rb/Th vs. Rb (ppm) plot of the Haţeg volcaniclastic rock samples (this study). For legend see Figure 6 Figure S8: Ta/Yb vs. Th/Yb (ppm) diagram of the Haţeg volcaniclastic rock samples (after [120]). The great majority of volcaniclastic rock samples are displaced away from the mantle array indicating a subduction-related enrichment. Sample P7A deviates from this general trend due to its high Nb and Ta concentrations. Mantle array compositions are from [55]. For legend see Figure 6. Figure S9: K2O vs. SiO2 wt% classification diagram [53] of the Haţeg volcaniclastic rock samples (this study)—orange circles, Haţeg rock samples from [7]—black circles; intrusive/extrusive rock samples from the Apuseni/Banat segments of the ABTS belt—gray dots. Figure S10: Rare Earth elements normalized to the chondrite values of [55] of Haţeg volcaniclastic rock samples (this study)—orange circles) and intrusive/extrusive rock samples from the Apuseni/Banat segments of the ABTS belt—grey dots. Figure S11: Primitive mantle normalized to the chondrite values of [55] of Haţeg volcaniclastic rock samples (this study)—orange circles and intrusive/extrusive rock samples from the Apuseni/Banat segments of the ABTS belt—grey dots. Table S1: Whole-rock major and trace elements for volcaniclastic rock samples of the Haţeg Basin [7]. Table S2: Eu/Eu* values of the Haţeg volcaniclastic rock samples (this study); normalized to the chondrite values of [55]. Table S3: Trace element ratios of the Haţeg volcaniclastic rock samples (this study).

Author Contributions

The authors carried out fieldwork, petrographic analyses and interpretation of the chemical and petrographic data. Conceptualization, V.M.V. and I.S.; validation, I.S. and V.M.V.; formal analysis, V.M.V.; investigation, V.M.V. and I.S.; resources, I.S.; writing—original draft preparation, V.M.V.; writing—review and editing, V.M.V. and I.S.; visualization, I.S.; supervision, I.S.; project administration, I.S.; funding acquisition, I.S. All authors have read and agreed to the published version of the manuscript.

Funding

This research was supported by the Romanian Executive Agency for Higher 814 Education, Research, Development, and Innovation; Funding projects PN-III-P4-ID-PCCF-2016-815 0014.

Data Availability Statement

The author confirms that the data supporting the findings of this study are available within the article and its Supplementary Materials. Further information are available from the corresponding author, upon request.

Conflicts of Interest

The authors declare no conflict of interest.

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Figure 1. The extension of the ABTS belt in relation to the main structural domains of the Alpine–Balkan–Carpathian–Dinaridic orogenic system. The ABTS belt is shown in red, and the yellow areas represent the Neogene magmatism. The Neotethys Ocean’s main suture zones are shown in black and dark gray [10].
Figure 1. The extension of the ABTS belt in relation to the main structural domains of the Alpine–Balkan–Carpathian–Dinaridic orogenic system. The ABTS belt is shown in red, and the yellow areas represent the Neogene magmatism. The Neotethys Ocean’s main suture zones are shown in black and dark gray [10].
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Figure 2. A simplified map of the Apuseni Mts. and Southern Carpathians showing the location of the Upper Cretaceous ABTS magmatic products of Banat and Apuseni segments and major tectonic units (see inset); the location of the Haţeg Basin is marked with a black rectangle (modified after [20,21] and the Geological map of Romania of the Geological Institute of Romania, scale 1:200.000).
Figure 2. A simplified map of the Apuseni Mts. and Southern Carpathians showing the location of the Upper Cretaceous ABTS magmatic products of Banat and Apuseni segments and major tectonic units (see inset); the location of the Haţeg Basin is marked with a black rectangle (modified after [20,21] and the Geological map of Romania of the Geological Institute of Romania, scale 1:200.000).
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Figure 4. General view of (a) volcaniclastic succession Densuş 1 (GPS coordinates: 45° 34′ 51″ N/22° 45′ 3″ E); (a1) base surge deposit of Densuș 1 volcaniclastic succession (a2) crystal-lithic tuff deposit of Densuș 1 volcaniclastic succession; (b) the Răchitova volcaniclastic succession (GPS coordinates: 45° 36′ 5″ N/22° 44′ 53″ E); (b1) pyroclastic surge deposit of the Răchitova volcaniclastic succession; (b2) lahar deposit of the Răchitova volcaniclastic succession.
Figure 4. General view of (a) volcaniclastic succession Densuş 1 (GPS coordinates: 45° 34′ 51″ N/22° 45′ 3″ E); (a1) base surge deposit of Densuș 1 volcaniclastic succession (a2) crystal-lithic tuff deposit of Densuș 1 volcaniclastic succession; (b) the Răchitova volcaniclastic succession (GPS coordinates: 45° 36′ 5″ N/22° 44′ 53″ E); (b1) pyroclastic surge deposit of the Răchitova volcaniclastic succession; (b2) lahar deposit of the Răchitova volcaniclastic succession.
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Figure 5. Microscope images of the selected volcaniclastic rock samples: (A) Sample P5A—basaltic trachy-andesite (Densuş, xpl image); (B) Sample E2—trachy-andesite (Densuş, ppl image); (C) Sample E5—dacite (Densuş, ppl image); (D) Sample P7A—rhyolite (Răchitova, xpl image); (E) Sample P6A—rhyolitic ignimbrite (Rachitova, ppl image); (F) Sample P2—trachytic ignimbrite (Peşteniţa, ppl image). Abbreviations: Plag—Plagioclase feldspar; Kfs—potassium feldspar; Qtz—quartz; Hbl—hornblende; Cpx—clinopyroxene; Bio—biotite; Ep—epidote; Chl—chlorite; ppl—plain-polarized light; xpl—cross-polarized light.
Figure 5. Microscope images of the selected volcaniclastic rock samples: (A) Sample P5A—basaltic trachy-andesite (Densuş, xpl image); (B) Sample E2—trachy-andesite (Densuş, ppl image); (C) Sample E5—dacite (Densuş, ppl image); (D) Sample P7A—rhyolite (Răchitova, xpl image); (E) Sample P6A—rhyolitic ignimbrite (Rachitova, ppl image); (F) Sample P2—trachytic ignimbrite (Peşteniţa, ppl image). Abbreviations: Plag—Plagioclase feldspar; Kfs—potassium feldspar; Qtz—quartz; Hbl—hornblende; Cpx—clinopyroxene; Bio—biotite; Ep—epidote; Chl—chlorite; ppl—plain-polarized light; xpl—cross-polarized light.
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Figure 6. Diagram showing total alkali vs. silica (wt%) (TAS) [52] of the Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol.
Figure 6. Diagram showing total alkali vs. silica (wt%) (TAS) [52] of the Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol.
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Figure 7. K2O vs. SiO2 wt% classification [53] of the Haţeg volcaniclastic rock samples, including the samples from [7], represented by the Y symbol. For legend, see Figure 6.
Figure 7. K2O vs. SiO2 wt% classification [53] of the Haţeg volcaniclastic rock samples, including the samples from [7], represented by the Y symbol. For legend, see Figure 6.
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Figure 8. (a) Classification of Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol, by SiO2 vs. Zr/TiO2 [54]; (b) classification of Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol, by Zr/TiO2 vs. Nb/Y [54]. For legend, see Figure 6.
Figure 8. (a) Classification of Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol, by SiO2 vs. Zr/TiO2 [54]; (b) classification of Haţeg volcaniclastic rock samples, including the samples from [7] represented by the Y symbol, by Zr/TiO2 vs. Nb/Y [54]. For legend, see Figure 6.
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Figure 9. Rare earth elements normalized to the chondrite values of [55]: (a) Haţeg volcaniclastic rock samples with SiO2 < 64%; (b) Haţeg volcaniclastic rock samples with SiO2 > 64%. For legend, see Figure 6.
Figure 9. Rare earth elements normalized to the chondrite values of [55]: (a) Haţeg volcaniclastic rock samples with SiO2 < 64%; (b) Haţeg volcaniclastic rock samples with SiO2 > 64%. For legend, see Figure 6.
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Figure 10. Primitive mantle normalized to the chondrite values of [55]: (a) Haţeg volcaniclastic rock samples with SiO2 < 64%; (b) Haţeg volcaniclastic rock samples with SiO2 > 64%. For comparison, we also plotted the trace element composition of the average upper crust (represented by black stars) and lower crust (represented by grey stars) of [56]. For legend, see Figure 6.
Figure 10. Primitive mantle normalized to the chondrite values of [55]: (a) Haţeg volcaniclastic rock samples with SiO2 < 64%; (b) Haţeg volcaniclastic rock samples with SiO2 > 64%. For comparison, we also plotted the trace element composition of the average upper crust (represented by black stars) and lower crust (represented by grey stars) of [56]. For legend, see Figure 6.
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Figure 11. (a–f) Variation in selected major oxides (wt%) with SiO2 content (wt%)—major oxide content decrease with increasing SiO2; (g) K2O vs. SiO2 (wt%) diagram—K2O is more or less constant in the intermediate rock samples and increases in the silica-rich samples; (h) Na2O vs. SiO2 content (wt%)—Na2O shows a positive correlation with SiO2 in the intermediate samples and tends to decrease in the more high-silica samples. For legend, see Figure 6.
Figure 11. (a–f) Variation in selected major oxides (wt%) with SiO2 content (wt%)—major oxide content decrease with increasing SiO2; (g) K2O vs. SiO2 (wt%) diagram—K2O is more or less constant in the intermediate rock samples and increases in the silica-rich samples; (h) Na2O vs. SiO2 content (wt%)—Na2O shows a positive correlation with SiO2 in the intermediate samples and tends to decrease in the more high-silica samples. For legend, see Figure 6.
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Figure 12. (a–f) Variation in selected immobile incompatible trace elements (ppm) with SiO2 content (wt%); (g) Rb (ppm) vs. SiO2 (wt%) diagram; (h) Sr (ppm) vs. SiO2 (wt%) diagram showing a slight decrease in Sr, whereas samples P6A and P6D exhibit a much more elevated concentration of this element. For legend, see Figure 6.
Figure 12. (a–f) Variation in selected immobile incompatible trace elements (ppm) with SiO2 content (wt%); (g) Rb (ppm) vs. SiO2 (wt%) diagram; (h) Sr (ppm) vs. SiO2 (wt%) diagram showing a slight decrease in Sr, whereas samples P6A and P6D exhibit a much more elevated concentration of this element. For legend, see Figure 6.
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Figure 13. Rb vs. Co (ppm) plot of the Haţeg volcaniclastic rock samples; a decrease in Co with increasing Rb is observed. For legend, see Figure 6.
Figure 13. Rb vs. Co (ppm) plot of the Haţeg volcaniclastic rock samples; a decrease in Co with increasing Rb is observed. For legend, see Figure 6.
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Figure 14. A plot of Dy/Dy* vs. Dy/Yb showing the decrease in Dy/Yb and Dy/Dy* with amphibole fractionation; garnet fractionation causes the increase in Dy/Yb; however, Dy/Dy* increases slightly or decreases (after [67]). The chondrite normalization constants are from [68]. Abbreviations: LREE—light rare-earth elements; amph—amphibole; cpx—clinopyroxene; grt—garnet. For legend, see Figure 6.
Figure 14. A plot of Dy/Dy* vs. Dy/Yb showing the decrease in Dy/Yb and Dy/Dy* with amphibole fractionation; garnet fractionation causes the increase in Dy/Yb; however, Dy/Dy* increases slightly or decreases (after [67]). The chondrite normalization constants are from [68]. Abbreviations: LREE—light rare-earth elements; amph—amphibole; cpx—clinopyroxene; grt—garnet. For legend, see Figure 6.
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Figure 15. (a) Y (ppm) vs. CaO (wt%) plot of the Haţeg volcaniclastic rock samples, compared with the calc-alkaline trend proposed by [79]; (b) Yb vs. La/Yb (ppm) plot for the Haţeg volcaniclastic rock samples. Fields for adakite-like and calc-alkaline magma compositions are from [80], and mineral differentiation paths are from [81]. Abbreviations: J-type—horblenditic trend; L-type—pyroxenitic trend; ol—olivine; opx—orthopyroxene; amph—amphibole; hbl—hornblende; cpx—clinopyroxene; plag—plagioclase. For legend, see Figure 6.
Figure 15. (a) Y (ppm) vs. CaO (wt%) plot of the Haţeg volcaniclastic rock samples, compared with the calc-alkaline trend proposed by [79]; (b) Yb vs. La/Yb (ppm) plot for the Haţeg volcaniclastic rock samples. Fields for adakite-like and calc-alkaline magma compositions are from [80], and mineral differentiation paths are from [81]. Abbreviations: J-type—horblenditic trend; L-type—pyroxenitic trend; ol—olivine; opx—orthopyroxene; amph—amphibole; hbl—hornblende; cpx—clinopyroxene; plag—plagioclase. For legend, see Figure 6.
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Figure 16. (a) Nb/Y vs. Th/Y (ppm) diagram of the Haţeg volcaniclastic rock samples (after [120]). The great majority of volcaniclastic rock samples are displaced away from the mantle array, indicating a subduction-related enrichment. Sample P7A deviates from this general trend due to its high Nb and Ta concentrations. Mantle array compositions are from [56]. (b) Nb/La vs. La/Yb (ppm) plot of the Haţeg volcaniclastic rock samples, showing the principal mantle source regions. The average lower crust is from [56], and the average OIB is from [55] (marked with black stars). The dividing line between the lithospheric and asthenospheric mantle fields is plotted using the data from [125]. For legend, see Figure 6.
Figure 16. (a) Nb/Y vs. Th/Y (ppm) diagram of the Haţeg volcaniclastic rock samples (after [120]). The great majority of volcaniclastic rock samples are displaced away from the mantle array, indicating a subduction-related enrichment. Sample P7A deviates from this general trend due to its high Nb and Ta concentrations. Mantle array compositions are from [56]. (b) Nb/La vs. La/Yb (ppm) plot of the Haţeg volcaniclastic rock samples, showing the principal mantle source regions. The average lower crust is from [56], and the average OIB is from [55] (marked with black stars). The dividing line between the lithospheric and asthenospheric mantle fields is plotted using the data from [125]. For legend, see Figure 6.
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Table 1. Sample name and location.
Table 1. Sample name and location.
SampleLocationLatLongPetrographic TypeGenetic Type
P5ADensuş45° 34′ 54″22° 46′ 53″basaltic trachy-andesitelava clast
E7Densuş45° 34′ 55″22° 46′ 17″trachy-andesitelava clast
P4ADensuş45° 35′ 04″22° 46′ 44″trachy-andesitelava clast
E4Densuş45° 34′ 54″22° 45′ 8″andesitelava clast
E3Densuş45° 34′ 52″22° 45′ 1″trachy-andesitelava clast
P6CRăchitova45° 36′ 2″22° 44′ 55″trachy-andesitelava clast
E6Densuş45° 34′ 55″22° 46′ 17″andesitelava clast
E5Densuş45° 34′ 55″22° 46′ 17″dacitelava clast
D2-2Densuş45° 34′ 54″22° 45′ 11″andesitelava clast
D2-3Densuş45° 34′ 54″22° 45′ 11″dacitelava clast
P5BDensuş45° 34′ 54″22° 46′ 53″trachy-andesitelava clast
E2Densuş45° 34′ 53″22° 45′ 2″trachy-andesitelava clast
P6BRăchitova45° 36′ 2″22° 44′ 54″trachy-dacitelava clast
D2-1Densuş45° 34′ 54″22° 45′ 11″dacitelava clast
P1Peşteniţa45° 33′ 45″22° 47′ 19″trachyteignimbrite deposit
P6ARăchitova45° 36′ 4″22° 44′ 55″rhyoliteignimbrite deposit
P4BDensuş45° 35′ 04″22° 46′ 44rhyolitelava clast
P2Peşteniţa45° 33′ 45″22° 47′ 19″trachyteignimbrite deposit
P6DRăchitova45° 36′ 3″22° 44′ 57″rhyoliteignimbrite deposit
P7ARăchitova45° 36′ 4″22° 44′ 53″rhyolitelava clast
Table 2. Whole-rock major, trace, and rare earth elements; D—Densuș; R—Răchitova; P—Peștenița.
Table 2. Whole-rock major, trace, and rare earth elements; D—Densuș; R—Răchitova; P—Peștenița.
SampleP5A
D		E7
D		P4A
D		E4
D		E3
D		P6C
R		E6
D		E5
D		D2-2
D		D2-3
D		P5B
D		E2
D		P6B
R		D2-1
D		P1
P		P6A
R		P4B
D		P2
P		P6D
R		P7A
R		LOD%
wt%
SiO253.355.655.856.558.459.359.460.260.361.061.161.96363.164.266.468.468.568.574.60.1
TiO20.780.730.770.610.630.880.770.580.630.560.620.520.740.560.820.370.330.770.280.120.1
Al2O218.8517.4517.4518.3519.217.817.4515.0517.6016.5016.4518.317.8516.3015.913.715.1515.2512.614.150.1
Fe2O36.347.676.986.495.776.235.826.095.905.135.645.975.896.383.561.862.844.141.462.520.1
MnO0.10.090.140.080.070.210.050.090.120.070.110.060.090.060.150.030.040.050.020.050.1
MgO4.393.114.732.871.921.762.342.481.252.062.191.561.221.470.691.140.680.321.150.20.1
CaO6.396.24.286.994.996.15.914.965.335.186.284.423.953.101.812.832.620.82.681.540.1
Na2O4.524.236.793.986.294.453.923.34.083.593.795.366.355.803.971.522.863.860.994.440.1
K2O2.631.921.161.71.342.242.042.572.092.453.081.931.140.626.692.965.466.882.533.430.1
P2O50.290.170.350.130.210.220.170.210.150.150.250.170.210.190.230.090.140.220.050.090.1
LOI3.434.372.772.551.871.851.963.31.522.351.71.521.431.661.858.472.821.099,540.240.1
Total101.18101.68101.33100.36100.8101.1499.9598.9599.0699.15101.36101.85101.9499.3099.9999.85101.54101.99100,19101.510.1
ppm
Ba6757853974194143735265974695375556322602919661970141083711708500.5
Rb47.629.527.962.738.887.663.863.789.476.999.128.839.310.4188.599.6120.5180.559.499.60.2
Sr523406385471460434424387414403542520326343147.5216023211420702650.1
Cs1.640.651.192.211.662.362.071.142.592.053.080.41.740.511.893.212.781.642.761.55
Ga24.922.320.921.219.623.623.119.518.516.621.118.919.612.620.718.418.813.417170.05
V155120168158144131160671301311378112367342516251755
Cr1242747593931592310100108451320<55<55<5<510
Co20.515.818.816.27.49.710.410.69.413.310.78.89.711.43.31.922.71.61.10.1
Ni22.21214.727.614.515.511.896.519.32411.410.213.22.42.81.22.72.52.10.2
U2.671.353.31.582.632.111.922.522.322.483.311.722.542.276.134.0735.644.174.30.05
Th10.257.0711.55.167.716.966.388.386.468.569.986.83108.6020.116.611.4518.616.2516.10.05
Zr1561241831171272611321441301501461291721573062401722862471102
Hf4.263.514.872.953.76.513.614.093.64.04.233.594.844.08.274.787.797.3230.1
Nb9.187.7510.557.418.856.928.147.057.06.77.146.458.677.416.513.48.5315.612.731.10.1
Ta0.50.40.70.50.70.40.50.50.50.50.40.40.50.61.10.90.610.92.10.1
Zn111479869697445418266695266557546427037382
Pb33.325.614.216.317.224.122.323.731.325.22427.717.217.427.527.22230.327.821.20.5
Y23.420.321.11617.819.115.810.619.318.316.614.421.716.433.13413.134.427.913.70.1
La3326.137.21626.325.623.423.720.927.132.632.131.528.350.354.730.448.943.242.10.1
Ce62.941.973.130.344.446404439.648.159.550.854.649.698100.554.795.179.167.20.1
Pr7.564.799.153.374.975.524.634.814.726.016.875.376.655.7811.611.55611.159.16.520.02
Nd291936.513.82022.518.118.117.622.926.819.62620.5474321.944.434.621.70.1
Sm6.83.557.093.164.074.83.433.363.584.375.243.434.823.759.387.723.548.96.743.380.03
Eu1.660.981.780.891.041.121.030.811.051.181.360.991.30.951.71.430.871.61.20.760.02
Gd5.473.055.513.043.784.093.112.523.614.204.333.184.283.257.416.532.887.025.832.350.05
Tb0.790.470.720.470.530.620.470.390.590.590.580.460.710.481.050.970.411.010.950.380.01
Dy43.063.812.963.133.782.682.163.543.493.332.83.922.826.025.912.495.875.222.230.05
Ho0.840.710.70.610.650.70.580.450.750.700.670.570.80.541.241.220.51.191.10.430.01
Er2.241.972.091.81.741.951.621.222.241.961.741.842.161.623.43.341.383.263.011.220.03
Tm0.30.280.290.280.260.30.260.180.330.290.250.280.330.240.530.470.220.490.480.220.01
Yb1.921.891.871.621.7521.81.172.241.831.681.912.191.443.63.631.63.363.331.690.03
Mg#43.4831.0642.9532.9526.9923.8930.8831.1519.0530.8530.1422.5018.6120.3817.7240.5121.017.9146.678.10
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Vornicu, V.M.; Seghedi, I. Petrography and Geochemistry of the Upper Cretaceous Volcaniclastic Deposits of the Haţeg Basin (Southern Carpathians): Inferences on Petrogenesis and Magma Origin. Minerals 2025, 15, 111. https://doi.org/10.3390/min15020111

AMA Style

Vornicu VM, Seghedi I. Petrography and Geochemistry of the Upper Cretaceous Volcaniclastic Deposits of the Haţeg Basin (Southern Carpathians): Inferences on Petrogenesis and Magma Origin. Minerals. 2025; 15(2):111. https://doi.org/10.3390/min15020111

Chicago/Turabian Style

Vornicu, Violeta M., and Ioan Seghedi. 2025. "Petrography and Geochemistry of the Upper Cretaceous Volcaniclastic Deposits of the Haţeg Basin (Southern Carpathians): Inferences on Petrogenesis and Magma Origin" Minerals 15, no. 2: 111. https://doi.org/10.3390/min15020111

APA Style

Vornicu, V. M., & Seghedi, I. (2025). Petrography and Geochemistry of the Upper Cretaceous Volcaniclastic Deposits of the Haţeg Basin (Southern Carpathians): Inferences on Petrogenesis and Magma Origin. Minerals, 15(2), 111. https://doi.org/10.3390/min15020111

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