Whole-rock geochemistry and U-Pb ages of Devonian bimodal-type rhyolites from the Rudny Altai, Russia: Petrogenesis and tectonic settings Share Link: https://authors.elsevier.com/a/1aP0h,UYEnRqHe

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Whole-rock geochemistry and U-Pb ages of Devonian bimodal-type
rhyolites from the Rudny Altai, Russia: Petrogenesis and tectonic settings
M.L. Kuibida
a,b,
⁎, O.V. Murzin
c
,N.N.Kruk
a,b
, I.Y. Safonova
a,b
,M.Sun
d
, T. Komiya
e
, J. Wong
d
, S. Aoki
e
,
N.M. Murzina
c
,I.Nikolaeva
a
, D.V. Semenova
a
, M. Khlestov
a
, R.A. Shelepaev
a,b
,P.D.Kotler
a,b
,
V.A. Yakovlev
a,b
, A.V. Naryzhnova
a
a
Sobolev Institute of Geology and Mineralogy SB RAS, Koptyuga ave. 3, Novosibirsk, Russia
b
Novosibirsk State University, Pirogova st. 2, Novosibirsk, Russia
c
Siberian Research Institute of Geology, Geophysics and Mineral Resources, Krasny prospect, 67, Novosibirsk, Russia
d
Department of Earth Sciences, the University of Hong Kong, Pokfulam Road, Hong Kong, China
e
University of Tokyo, 3-8-1 Komaba, Meguro-ku, Tokyo 153-8902, Japan
abstractarticle info
Article history:
Received 25 August 2019
Received in revised form 25 November 2019
Accepted 9 December 2019
Available online 27 December 2019
Handling Editor: S. Kwon
Keywords:
Rifted-arc
Back-arc basin
Bimodal basalt-rhyolite association
Melnichno-Sosnovsky volcanic complex
Central Asian Orogenic Belt
The paper presents new original data on the Devonian felsic volcanism of the NW Rudny Altai (Russia) in the
west of Central Asian Orogenic Belt (CAOB) –the front part of the Altai convergent margin of the Siberian conti-
nent. Two geochemical types of subvolcanic rhyolites were emplaced synchronously with the bimodal rhyolite-
basalt association, which began to form in the end-Emsian, and clearly manifested on the border of the Givetian
and the Frasnian. The rhyolites yield zircon U-Pb ages of ca. 390 Ma (R1-type) and 380 Ma (R2- and R3-types),
reflecting two peaks of the volcanic activity. Most of these rocks have extreme petrochemical characteristics of
high SiO
2
contents and have contrast Na/K ratios. Their compositions are transition between calc-alkaline and
tholeiite series: (La/Yb)
n
~2–7, Zr/Y ~ 4 (Zr b350 ppm) and La/Sm ~ 0.55–1. Rhyolites bear the distinctive geo-
chemical signature of A-type felsic magma, such as enrichments in Zr, Nb, Y and Ce (N350 ppm), Zr (N250 ppm),
and high Ga/Al (N2.6) values. The island-arc-like R1-rhyolite formed immediately after the beginning of rifting
due to widespread crustal melting under reduced conditions. The generation of rift-like R2- and R3-rhyolites
took place under non-equilibrium conditions, synchronously with the rise in the upper crust of Givetian-
Frasnian basic magmas, as a result of the active lithospheric extension and high thermal input from the underly-
ing hot mantle. We propose an extension regime in the transition area between the island-arc and back-arc basin
for the origin of rhyolites. The study of the Devonian volcanism of the Rudny Altai gives important information
about the processes that occurred at the initial stage of the formation of the Altai convergent margin.
© 2020 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
1. Introduction
In the context of the analysis of magmatic formations, the early geo-
synclinal evolution of the marginal parts of the platforms is indicated by
the formation of the so-called spilite-keratophyric association, which is
a result of specific submarine explosive volcanism of essentially sialic-
sodium composition (Daly, 1914;Rosenbursh, 1923;Battey, 1955;
Kuznetsov, 1964). In case of rift-related extension of continental mar-
gins, mature island-arc, and marginal back-arc basin association is char-
acterized by the predominance of felsic eruption and genetically related
subvolcanic intrusions. Bimodal-felsic and bimodal-siliciclastic types
are common (Barrie and Hannington, 1999), as manifested on the an-
cient Pacific margin of North America (New-Brunswick and Isok lake
deposits in Canada), the modern Pacific margin of Asia (e.g. Hokuroku
district of Japan), and the Paleo-Asian margin of Siberian continent
(Rudny Altai; Lentz, 1998;Hart et al., 2004;Piercey et al., 2006;Galley
et al., 2007;Gaboury and Pearson, 2008;Gaskov, 2015). In this paper,
felsic volcanic rocks from the bimodal association of such geodynamic
settings, are called “bimodal-type extensional rhyolites”(Bachmann
and Bergantz, 2008;Tamura et al., 2009;Deering et al., 2008, 2010).
The Rudny Altai is located in the western part of the Central Asian
Orogenic Belt (CAOB) or Altaid Tectonic Collage (Altaids) –the world's
largest accretionary orogen, a Pacific-type system that comprises multi-
ple subduction-accretionary orogens (e.g., Zonenshain et al., 1990;
Mossakovsky et al., 1993;Sengör et al., 1993;Dobretsov et al., 1995;
Jahn et al., 2000;Buslov et al., 2001;Yakubchuk, 2004;Buslov et al.,
2004;Windley et al., 2007;Xiao et al., 2010;Safonova et al., 2011;
Kröner et al., 2014;Xiao and Santosh, 2014;Chen et al., 2017;
Safonova et al., 2018). The western Altaids segment extends for
N2500 km from the Russian Ural in the west, across Eastern
Gondwana Research 81 (2020) 312–338
⁎Corresponding author at: Sob olev Institute of Geology and Minera logy SB RAS,
Koptyuga ave. 3, Novosibirsk, Russia.
E-mail address: maxkub@igm.nsc.ru (M.L. Kuibida).
https://doi.org/10.1016/j.gr.2019.12.002
1342-937X/© 2020 International Association for Gondwana Research. Published by Elsevier B.V. All rights reserved.
Contents lists available at ScienceDirect
Gondwana Research
journal homepage: www.elsevier.com/locate/gr

Kazakhstan and Northwest China to Southern Mongolia. It developed in
the Late Neoproterozoic-Late Paleozoic by closure and suturing of the
Paleo-Asian Ocean (PAO), which once separated the Siberian Craton
from the Tarim, North China and Kazakhstan continents. The PAO his-
tory continued for N800 Ma and included multiple events of subduction
and accretion of oceanic island/plateau, island-arc, and terranes, of both
western Pacific- and Andean-styles, which had eventually amalgamated
into the Altaids in the Late Paleozoic.
The Altai orogenic belt (Bespaev et al., 1997;Scherba et al., 1998)or
the Altai accretion-collision system (Vladimirov et al., 2003)isa
Hercynian fold belt, extending for 2000 km from Russia and
Kazakhstan through northern China to southwestern Mongolia
(Sengör et al., 1993;Jahn et al., 2000;Buslov et al., 2001;Windley
et al., 2007;Xiao et al., 2010;Pirajno, 2010;Cai et al., 2011). It developed
along the Caledonian marginal terranes of the Siberian continent, due to
the subduction and closure of Paleo-Asian Ocean, collision and post-
collision tectonic processes. The formation of the Altai convergent mar-
gin is associated with the Devonian oblique subduction of the Chara
oceanic plate under the terrain-orogenic margin of the Siberian conti-
nent. Tectonically, it consists of the frontal part (Rudny Altai), the
axial (Central and Western Gorny Altai and Salair) and the rear part
(Uimeno-Lebed in the Eastern Gorny Altai and Western Sayan); Fig. 1
(Yolkin et al., 1994;Shokalsky et al., 2000).
In the last decades a significant progress has been made on the study
of the Chinese and Mongolian Altai (Yakubchuk, 2004;Xiao et al., 2010;
Wan et al., 2010;Cai et al., 2011;Yarmolyuk et al., 2013;Wu et al., 2015;
Chen et al., 2017;Yang et al., 2018;Şengör et al., 2018), but detailed and
systematic geochemical and isotope geochronological study is limited
for the Rudny Altai one of the largest Devonian polymetallic provinces
(Scherba et al., 1998). Some of the sulfide deposits are considered to
be an analogue of the “Kuroko”type, i.e. in connection with long-lived
hot metal-rich hydrothermal systems of black smokers; whereas others
are considered as syngenetic with high temperature subvolcanic felsic
intrusions, which were a conductor of heat and fluids and control con-
vection of metal-precipitating hydrothermal flows (Gaskov, 2015).
Here we report new geochemical data, zircon U-Pb and Nd isotopic
compositions of the subvolcanic rhyolites from the bimodal association
from the NW Rudny Altai (Russia) and discuss their petrogenesis,
magma sources and geodynamic setting.
2. Tectonic setting and geological framework
The Rudny Altai tectonic block is located on the territory of Russia,
Kazakhstan and Chinese Altai, and extends in the North-Western direc-
tion for N500 km with a width of about 100 km, in the strip between the
North-Eastern regional faults (NEF) and Irtysh Shear Zone (ISZ); Fig. 2a.
On the basis of tectonic and volcanic zoning (Kulkov, 1980), its territory
is subdivided into five zones: (i) Aley in Russia; (ii–iv) Sinukha-Holzun,
Nizhnebukhtarma and Yuzhnoaltai in Eastern Kazakhstan (Bespaev
et al., 1997;Scherba et al., 1998); and (v) Ashele in Xinjiang of China
(e.g. Wan et al., 2010;Wu et al., 2015;Yang et al., 2018).
Until the start of Devonian volcanism this territory existed as part of
the shelf on the margin of the Siberian continent with terrigenous-
carbonate sedimentation, similar to the modern passive continental
margins (Yolkin et al., 1994). Suggested that the passive margin was
“dragged”in the process of Paleozoic drift and rotation of the Siberian
continent (Scherba et al., 1998). The Devonian volcano-plutonic belts
were formed as a result of the powerful destruction of the Caledonian
passive margin, probably on the border of Pragian and Emsian
(Rotarash et al., 1982;Scherba et al., 1998). Volcanism developed in
the environment of a shallow-sea basin with depths about 0.5 km and
not exceeding 1.5 km (Rotarash et al., 1982;Yolkin et al., 1994;
Bespaev et al., 1997;Saraev et al., 2012). NW Rudny Altai is represented
by the so-called Alei block –the outcrop of the Caledonian deformed
basement, with a total size of 200 × 50–60 km. In Early Devonian the
Alei Block was subsided and formed a rhomboidal-shaped structure,
like pull-apart basin, now looks like half-graben (Fig. 2a). This structure
is bounded by the general systems of north-western and sub-latitude
faults. The first of them correspond to the ISZ and the NEF zone formed
at the stage of formation of the Altai convergent margin, since they con-
trol the emplacement of two volcanic belts (Fig. 2c). According to geo-
logical mapping at a scale of 1:200,000 (Murzin et al., 2001), these
volcanic belts in the sides of the Alei structure were formed synchro-
nously; although it has not yet been determined whether the composi-
tions of volcanic rocks from these belts are identical. The formation of
feathering sub-latitudinal faults occurred at the boundary of Givetian
and Frasnian (Scherba et al., 1998). In general, based on the structural
pattern of the Alei block, it can be argued that the Chara subduction oce-
anic plate under the Rudny-Altai block was obliquely, WE direction, in
modern coordinates. Thus, this differs from the proposed northward
subduction of the Junggar plate under the Devonian margin of the Sibe-
rian continent (e.g. Xu et al., 2003;Zhang et al., 2018;Ma et al., 2018).
3. Evolution of the volcanism and characteristics of subvolcanic
intrusion
3.1. Evolution of the volcanism
The Emsian-Frasnian volcanism was developed in two major stages
involving five volcanic rhythms (Fig. 3;Murzin et al., 2001). In this pe-
riod powerful outbreaks of contrasting predominantly shallow-sea vol-
canism, accompanied by the injection of subvolcanic and hypabyssal
intrusions associated with linear zones of extension; and the formation
of numerous polymetallic deposits alternating with partial attenuation
stages of volcanic activity (Scherba et al., 1998;Shokalsky et al., 2000;
Murzin et al., 2001). The first stage (Late Emsian –Early Givetian) in-
cludes two rhythms of felsic volcanism, formed the Melnichno-
Sosnovsky (MS) volcanic complex. The first rhythm (Late Emsian –
Eifelian) was represented by rare submarine eruptions of felsic lavas
and their tuffs, which had minor distribution in comparison with the
prevailing fine-grained sedimentary deposits. In contrast, the second
rhythm (Early Givetian) was characterized by active eruptions from
large volcanoes of the Central-type,and powerful pyroclastic felsic com-
position deposition, and intrusion of syngenetic subvolcanic rhyolites.
The first stage total rock thickness is ~400 m. The second pronounced
stage or so-called “Givetian –Frasnian rhyolite-basalt flash”
corresponded to the clear mode of extension, and the formation of the
so-called Shipunikha rift-like structure in the rear part of the NW
Rudny Altai, accompanied by fractured effusions of basalts (Fig. 2a).
This stage includes three volcanic rhythms, generally called the third,
fourth and fifth. Accordingly, the third rhythm covers the period of for-
mation of the varied stratified strata, consisting of lava, lava-breccia,
tuffs and felsic ignimbrites composition (Fig. 3). The fourth rhythm is
reflected by the eruptions of basalt flows during the subsidence of the
crustal blocks. The fifth rhythm appeared as felsic eruption from numer-
ous volcanoes of the Central-type, inherited on the site of Givetian-
Frasnian volcanic structures.With the depletion of magmatic chambers
in some areas of the Rudny-Altai block, the subsidence of pre-existing
volcanic apparatus and the formation of volcanic-tectonic depressions
in their place took place, which, at least until the middle of theFrasnian,
maintained a tendency to dive. The total rock thickness is ~450 m. In a
first approximation, U-Pb SHRIMP-II age of syngenetic granitoids reveal
of the first and second stage: ~395–384 Ma (Kuibida et al., 2015) and
~378–372 Ma (Murzin et al., 2001), respectively. In general, a character-
istic feature of volcanism in the NW Rudny Altai wasthe formation of a
contrasting rhyolite-basalt association with the antidromic order of the
rhyolites-basalt volcanism in the Emsian-Eifelian; and the pendulum
nature of the rhyolite-basalt-rhyolite volcanism in Givetian-Frasnian
period. The total volume of felsic volcanism prevailed over the basic vol-
canism in a ratio of 75:25%; and volcanic products of the intermediate
composition were almost absent (Rotarash et al., 1982;Murzin et al.,
2001).
313M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Fig. 1. Structural-geological map of the western part of the Altai accretion ary-collision system (Russian Altai) modified after 1:200,000 Geological Maps of the Russ ian Federation; illustrating the location of the two Devonian volcanic systems (data
source from http://www.vsegei.ru/en/geology-of-russia). Thetectonic framework is from (Buslov, 2011). Tectonic blocks: RA –Rudny Altai, GA –Gorny Altai (integrated), AM –Altai-Mongolian. Volcanic belts: RA –Rudny Altai, KA –Korgon-Aksai,
KhS –Kholzun-Sarymsakty. The numbers in the rectangles indicate the ages (Ma). Normal font (1:200,000 Ge ological Maps of the Russian Federation (http://www.vsegei.ru/en/geology-of-russia), and (Kuibida et al., 2015;Kuibida, 2019); italic font
(Glorie et al., 2011). Ages data for Melnichno-Sosnovsky (MS) volcanic complex is from this study.
314 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

3.2. Characteristics of subvolcanic intrusion
Subvolcanic MS-rhyolites were collected mainly from the NE volca-
nic belt (Fig. 2) including several volcanic centers located in a chain on
a direction NW-SE: Voronezh, Karaulny, Butochny and Sadovushka
(Fig. 4a–h). Butochny and Sadovushka were originally parts of a single
volcanic center, but now they are separated by the local pull-apart
structure inside the Sipunikha rift (Fig. 2). The volcanic center Kryuchki,
the only exposure of MS-rhyolite in the SW volcanic belt, was also col-
lected for this study (Fig. 2). Subvolcanic intrusions occur as stock-,
sill- and dike-like bodies localized inside the Late Emsian –Early
Givetian stratified volcanogenic-sedimentary deposits of the first stage
(Fig. 3). A large number of rhyolites have experienced significant brittle
deformations, and are quite often subject to secondary alterations, in-
cluding limonitization within micro-fracture.
On the basis of field observations, study of thin sections and geo-
chemical compositions we have identified three main types of MS-
rhyolites described in this study as R1-, R2- and R3-rhyolites. Rhyolites
of the first type compose the main volume of all subvolcanic intrusions
in NW Rudny Altai. Rhyolites of the second type are related to the late
phases, which form a small-sized subvolcanic body in the frame of the
main volcanic edifices. The third type of rhyolite was found only within
the Butochniy volcanic center, where they are located, both inside the
volcanic structure and in its frame to the North and South.
Fig. 2. (a) Shaded relief image (illuminated from the NW), representing the digital elevation model of the NW Rudny Altai, illustrating the location of the two Devonian volcanic belts
formed along the NW direction of deformation in the boards of the Alei basement rhomboidal-shaped structure. (b) The principal tectonic scheme of the Rudny Altai, illustrating the
division on structure-formation units. (c) Detailed three-dimensional topographic mapping scheme on panel a, illustrating the position Alei structure and Shipunikha rift in structure
of NW Rudny Altai. The vertical scale is increased five times.
315M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

4. Analytical methods
4.1. Zircon U-Pb dating
Zircon separation of an 8 kg rhyolite samples were conducted in the
Institute of Geology and Mineralogy by a standard procedure of
crushing, panning, heavy liquid and magnetic separation techniques.
Zircon grains were hand-picked under a binocular and then mounted
in a 20 mm and 6 mm epoxy resin disc. All grains were half-polished
to observe the internal structure of zircons at cathodoluminescence
(CL) images. The CL imaging was performed at the Novosibirsk Institute
Geology and Mineralogy using a Jeol JSM-6510LV scanning electron mi-
croscope; and at the University of Tokyo using a JSM-6060VL scanning
electron microscope (JEOL, Ltd., Tokyo, Japan) equipped with SSM-
7CLS cathodoluminescence (CL) system (Sanyu Electron Co., Ltd.,
Tokyo, Japan), which was operated at an accelerating voltage of 15 kV.
The locations for the spot analysis on zircon grains were selected from
CL images and photomicrographs, in transmitted and reflected light,
to avoid mineral inclusions and cracks. The CL-images of zircons and
the results of
206
Pb/
238
U dating are shown in Tables 1, 2 and Figs.6,7.
In situ zircon U-Pb datings of MS-10/1 and MS-13/1 samples were
conducted by using an Agilent 8800 single-collector triple quadrupole
ICP-MS (Agilent Tech., Santa Clara, USA) coupled to a NWR-213 Nd:
YAG LA system (ESI, Portland, USA) at the Gakushuin University. The
sample discs were set inthe two-volume sample chamber of the LA sys-
tem. The locations for the spot analysis on zircon grains were chosen by
referring the CL images and observations with the LA camera to avoid
mineral inclusions and cracks. Before the analysis, the locations were
ablated using a pulse of laser to remove potential contaminants on the
zircon surfaces. At the analysis, they were ablated for 30 s by the laser
with fluence of 2.0 J/cm
2
, repetition rate of 5 Hz, and laser spot size di-
ameter of 30 μm after laser shooting with laser shutter closed for 30 s
(laser warming up). The ablated materials were carried by He gas,
which was introduced into the two-volume sample chamber at a flow
rate of 0.6 l/min. They were mixed with Ar make-up gas introduced at
aflow rate of 1.2 l/min before their introduction into the ICP-MS. On
the ICP-MS, 6 nuclides (
202
Hg,
204
Pb,
206
Pb,
207
Pb,
232
Th and
238
U)
were analyzed. In order to reduce isobaric interference of
204
Hg to
204
Pb, small amount of mixed ammonia/helium was flowed into colli-
sion/reaction cell between the tandem quadropole mass spectrometers
(e.g. Kasapoğlu et al., 2016). We succeeded to reduce the blank count on
204 amu (
204
Hg +
204
Pb) b10 cps on average. The background and ab-
lation data for each analysis were collected for 15 s of the laser
warming-up time and 20 s of the ablation time, respectively. Those
data were acquired for multiple groups of 15 unknown grains bracketed
by trio of analyses of the 91500 zircon standard (Wiedenbeck et al.,
2004) and NIST SRM610 glass standard. The 91500 was analyzed for
correction of elemental fractionation bias of
206
Pb/
238
U and Th/U ratios.
For the correction, apparent
206
Pb/
238
U ratio without common Pb cor-
rection by Sakata et al. (2017) and Th/U ratio (Wiedenbeck et al.,
2004) were used as normalization values, respectively. The NIST SRM
610 was analyzed for correction of mass fractionation bias of
207
Pb/
206
Pb ratios. The isotopic ratio compiled by Jochum and Nohl
(2008) was used as a normalization value for the correction. The back-
ground intensities collected at the laser warming-up time were
subtracted from following signals at the ablations.
235
U was calculated
from
238
U using a
238
U/
235
U ratio of 137.88 (Jaffey et al., 1971). The in-
tensity of
202
Hg of all analyses was used to correct the isobaric interfer-
ence of
204
Hg on
204
Pb. Corrected
204
Pb intensities were too low to
correct U-Pb ages for common Pb contamination with sufficient preci-
sion based on
204
Pb (Stern, 1997). Thus, in this study, no common Pb
correction was made. All uncertainties of the data are quoted at a 2
sigma level to which repeatability of each six measurements of 91500
zircon and NIST SRM 612 data bracketing unknown sample groups is
propagated. Elemental fractionation of U/Pb and Th/U ratios and mass
fractionation of
207
Pb/
206
Pb ratio were linearly interpolated by the mea-
sured data of each six analysis of 91500 zircon and NIST SRM610, re-
spectively. Resulting age interpretations, weighted mean U-Pb ages
and Concordia plots were constructed using Isoplot software
(Ludwing, 2003). Through all the analyses, Plešovice zircons (Slama
et al., 2008) were measured multiple times as secondary standards for
quality control. Mean concordia age obtained for Plesovice standard
yields
206
Pb/
238
U age of 337.2 ± 1.8 Ma (MSWD = 1.0) and
207
Pb/
235
U age of 340.2 ± 5.8 Ma (MSWD = 1.2) (n = 25), which is
in accordance with the reported ID-TIMS age of 337.1 ± 0.4 Ma
(Slama et al., 2008).
For rhyolite samples MS-8, MS-26 and MS-32 the U-Pb geochrono-
logical studies were performed by the LA-SF-ICP-MS method on a
Thermo Fisher Scientific ElementXR high-resolution mass spectrometer
Fig. 3. Stratigraphic columns of end Early Devonian –early Late Devonian volcanic and
sedimentary rocks in th e NW Rudny Altai (modified fro m 1:200,000 geological maps
(Murzin et al., 2001)), illu strating the two stage evolution of bimodal volcanism, and
syngenetic subvolcanic intrusions.
316 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Fig. 4. Representative photos showing survey panoramas of the mainvolcanic centersand intrusive bodiesof MS-rhyolite: (a,b) Karaukny; (c) Voronezh; (c) Butochny (d, e);Kruchki (g,
h) of the main petrographic types of volcanic rocks: R1-rhyolite (i–m); R2-rhyolite (n, o); and R3-rhyolite (p, q).
317M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Table 1
LA-ICP-MS zircon U-Pb isotopic data of the Melnichno-Sosnovsky subvolcanic rhyolites (NW Rudny Altai).
Point no Isotope rations Ages (Ma) Concentration (ppm)
207
Pb/
235
U2σ
206
Pb/
238
U2σRho
207
Pb/
235
U2σ
206
Pb/
238
U2σ% conc. U Th Th/U
R1: sample MS-10/1. LA-ICP-MS (Gakushuin University, Japan)
1 0,45,529 0,00744 0,06103 0,00058 0,58 375,5 30,0 381 5,2 99 397 166 0,42
2 0,45,373 0,00912 0,06134 0,00062 0,5 356,2 39,2 379,9 6,4 94 231 82 0,35
3 0,45,308 0,00866 0,06117 0,00061 0,52 359,4 36,9 379,4 6,1 95 262 84 0,32
4 0,45,311 0,00905 0,06059 0,00061 0,51 381 38,7 379,5 6,3 100 235 78 0,33
5 0,45,132 0,00902 0,06049 0,00061 0,51 375,7 38,8 378,2 6,3 99 236 79 0,34
6 0,45,881 0,00894 0,06075 0,00061 0,51 403 37,4 383,4 6,2 105 248 87 0,35
7 0,45,750 0,00812 0,06092 0,00059 0,55 390,2 33,4 382,5 5,7 102 316 126 0,4
8 0,45,404 0,01064 0,06162 0,00094 0,65 347,7 40,1 380,1 7,4 91 230 82 0,35
9 0,44,859 0,01082 0,06003 0,00109 0,75 379,3 35,6 376,3 7,6 101 301 118 0,39
10 0,46,179 0,01098 0,06003 0,00109 0,76 444,2 34,2 385,5 7,6 115 312 127 0,41
11 0,46,257 0,01340 0,06051 0,00114 0,65 430,3 48,9 386 9,3 111 150 59 0,39
12 0,45,673 0,01062 0,06050 0,00109 0,78 402,1 32,7 382 7,4 105 349 145 0,41
13 0,45,231 0,01122 0,06002 0,00110 0,74 398,4 37,6 378,9 7,8 105 265 91 0,34
14 0,45,656 0,01119 0,06039 0,00110 0,74 405,2 36,6 381,9 7,8 106 276 101 0,37
15 0,45,841 0,01105 0,06094 0,00111 0,75 394 35,5 383,2 7,7 103 294 110 0,37
16 0,45,827 0,01068 0,06188 0,00112 0,78 359,1 33,2 383,1 7,4 94 341 147 0,43
17 0,46,914 0,01126 0,06187 0,00113 0,76 412,2 35,0 390,6 7,8 106 296 122 0,41
18 0,46,250 0,00895 0,06207 0,00062 0,52 373 37,3 386 6,2 97 249 103 0,41
19 0,46,964 0,00940 0,06216 0,00063 0,51 403,8 38,6 390,9 6,5 103 226 84 0,37
20 0,46,789 0,01018 0,06222 0,00094 0,69 393,5 35,2 389,7 7 101 289 108 0,37
21 0,46,751 0,01034 0,06206 0,00094 0,68 397,5 36,1 389,5 7,2 102 273 101 0,37
22 0,46,255 0,00969 0,06237 0,00093 0,71 362,1 33,1 386 6,7 94 334 146 0,44
23 0,46,182 0,01216 0,06198 0,00098 0,6 372,8 47,4 385,5 8,4 97 159 60 0,38
24 0,46,121 0,01085 0,06229 0,00096 0,65 358,6 40,3 385,1 7,5 93 224 77 0,34
25 0,47,375 0,01006 0,06229 0,00094 0,71 418,6 33,5 393,8 6,9 106 312 137 0,44
26 0,46,519 0,01097 0,06124 0,00111 0,77 415,9 33,7 387,9 7,6 107 320 133 0,42
27 0,47,351 0,01009 0,06238 0,00094 0,71 414,3 33,7 393,6 7 105 310 135 0,44
28 0,46,942 0,01088 0,06299 0,00114 0,78 373,1 32,6 390,8 7,5 95 342 142 0,41
R1: sample MS-14/1. LA-ICP-MS (Hong-Kong University)
1 0,46,905 0,01061 0,06430 0,00120 0,82 390,5 7,3 401,7 7,3 97 413 150 0,36
2 0,46,940 0,00921 0,06265 0,00105 0,85 390,8 6,4 391,7 6,4 99 379 156 0,41
3 0,46,984 0,00938 0,06077 0,00081 0,66 391,1 6,5 380,3 4,9 97 362 198 0,55
4 0,47,001 0,01219 0,06296 0,00073 0,45 391,2 8,4 393,6 4,5 99 404 136 0,34
5 0,46,992 0,00880 0,06018 0,00083 0,73 391,1 6,1 376,7 5,0 96 360 184 0,51
6 0,46,956 0,01283 0,05723 0,00107 0,69 390,9 8,9 358,7 6,5 91 373 183 0,49
7 0,46,928 0,00920 0,06246 0,00091 0,74 390,7 6,4 390,6 5,5 99 360 175 0,49
8 0,47,123 0,01088 0,06140 0,00119 0,84 392,0 7,5 384,1 7,2 97 347 189 0,54
9 0,46,955 0,01412 0,05891 0,00158 0,89 390,9 9,8 369 9,6 94 345 194 0,56
10 0,47,157 0,00893 0,05809 0,00087 0,79 392,3 6,2 364 5,3 92 400 164 0,41
11 0,46,961 0,00882 0,06222 0,00080 0,69 390,9 6,1 389,1 4,9 99 366 180 0,49
12 0,46,813 0,00866 0,06211 0,00087 0,76 389,9 6,0 388,5 5,3 99 350 195 0,56
R1: sample MS-12. LA-ICP-MS (Hong-Kong University)
1 0,46,726 0,00402 0,06146 0,00047 0,90 389,3 2,8 384,5 2,9 98 345 218 0,63
2 0,46,730 0,00263 0,06246 0,00030 0,85 389,3 1,8 390,6 1,8 99 331 226 0,68
3 0,46,741 0,00366 0,06319 0,00025 0,51 389,4 2,5 395 1,5 98 314 168 0,54
4 0,46,730 0,00276 0,06218 0,00027 0,73 389,3 1,9 388,9 1,6 99 301 249 0,83
5 0,46,792 0,00396 0,06186 0,00022 0,41 389,7 2,7 386,9 1,3 99 342 188 0,55
6 0,46,699 0,00485 0,06135 0,00044 0,69 389,1 3,4 383,8 2,7 98 307 172 0,56
7 0,46,768 0,00267 0,06259 0,00033 0,93 389,6 1,8 391,3 2,0 99 395 194 0,49
8 0,46,743 0,00351 0,06304 0,00037 0,78 389,4 2,4 394,1 2,2 98 400 175 0,44
9 0,46,785 0,00411 0,06232 0,00039 0,71 389,7 2,8 389,7 2,3 99 351 213 0,61
10 0,46,704 0,00225 0,06196 0,00025 0,83 389,1 1,6 387,5 1,5 99 356 220 0,62
11 0,46,757 0,00247 0,06229 0,00029 0,87 389,5 1,7 389,5 1,7 99 349 198 0,57
12 0,46,726 0,00314 0,06106 0,00036 0,89 389,3 2,2 382,1 2,2 98 342 219 0,64
13 0,46,738 0,00291 0,06190 0,00036 0,95 389,4 2 387,2 2,2 99 347 223 0,64
14 0,46,751 0,00414 0,06137 0,00052 0,95 389,5 2,9 384 3,1 98 349 220 0,63
15 0,46,750 0,00383 0,06227 0,00046 0,91 389,5 2,6 389,4 2,8 99 409 186 0,45
R1: sample MS-26. LA-ICP-MS (IGM SB RUS, Novosibirsk, Russia)
1 0,46,573 0,01189 0,06158 0,00102 0,65 388,2 8,2 385,2 6,2 100 141 310 0,45
2 0,46,575 0,01104 0,06190 0,00102 0,70 388,2 7,7 387,2 6,2 99 190 351 0,54
3 0,46,305 0,01423 0,06173 0,00105 0,55 386,4 9,9 386,2 6,4 99 136 167 0,81
4 0,48,381 0,01101 0,06195 0,00102 0,72 400,7 7,5 387,5 6,2 102 535 635 0,84
5 0,47,126 0,01209 0,06180 0,00103 0,65 392,1 8,3 386,6 6,2 100 209 388 0,54
6 0,46,576 0,00873 0,06236 0,00101 0,86 388,3 6,1 390 6,1 101 804 1183 0,68
7 0,46,450 0,01172 0,06195 0,00103 0,66 387,4 8,1 387,5 6,3 99 213 301 0,71
8 0,46,715 0,01114 0,06173 0,00102 0,69 389,2 7,7 386,2 6,2 100 580 546 1,06
9 0,46,638 0,01228 0,06195 0,00104 0,64 388,7 8,5 387,4 6,3 99 139 263 0,53
10 0,48,012 0,01191 0,06195 0,00103 0,67 398,2 8,2 387,5 6,3 102 298 462 0,65
11 0,46,822 0,01068 0,06188 0,00102 0,72 390 7,4 387,1 6,2 100 505 472 1,07
12 0,48,308 0,01199 0,06230 0,00104 0,67 400,2 8,2 389,6 6,3 102 282 352 0,80
13 0,47,824 0,01145 0,06205 0,00103 0,69 396,9 7,9 388,1 6,3 101 356 426 0,84
318 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Table 1 (continued)
Point no Isotope rations Ages (Ma) Concentration (ppm)
207
Pb/
235
U2σ
206
Pb/
238
U2σRho
207
Pb/
235
U2σ
206
Pb/
238
U2σ% conc. U Th Th/U
14 0,46,902 0,01359 0,06227 0,00106 0,59 390,5 9,4 389,4 6,4 99 147 249 0,59
15 0,46,591 0,01697 0,06212 0,00109 0,48 388,4 11,8 388,5 6,6 99 64 108 0,59
16 0,46,814 0,01071 0,06243 0,00103 0,72 389,9 7,4 390,4 6,3 101 412 530 0,78
17 0,46,621 0,01080 0,06193 0,00102 0,71 388,6 7,5 387,3 6,2 99 278 388 0,72
18 0,46,946 0,01219 0,06193 0,00104 0,65 390,8 8,4 387,3 6,3 100 360 412 0,87
19 0,46,917 0,01027 0,06229 0,00103 0,76 390,6 7,1 389,5 6,2 99 631 731 0,86
20 0,48,887 0,01164 0,06229 0,00103 0,69 404,1 7,9 389,5 6,3 103 399 456 0,87
21 0,46,917 0,01244 0,06191 0,00104 0,63 390,6 8,6 387,2 6,3 100 248 326 0,76
22 0,46,810 0,01115 0,06227 0,00103 0,69 389,9 7,7 389,4 6,3 99 353 463 0,76
23 0,50,836 0,02233 0,06194 0,00113 0,42 417,3 15 387,4 6,9 107 68 114 0,6
24 0,46,832 0,01290 0,06172 0,00104 0,61 390 8,9 386,1 6,3 100 359 404 0,89
25 0,47,093 0,01302 0,06269 0,00106 0,61 391,8 9 392 6,4 101 193 300 0,64
26 0,46,615 0,01374 0,06167 0,00105 0,58 388,5 9,5 385,8 6,4 100 122 220 0,56
27 0,46,536 0,01563 0,06183 0,00107 0,52 388 10,8 386,8 6,5 99 241 193 1,25
28 0,46,769 0,01558 0,06199 0,00107 0,52 389,6 10,8 387,7 6,5 99 125 172 0,73
29 0,49,424 0,01643 0,06189 0,00107 0,52 407,8 11,2 387,1 6,5 104 161 215 0,75
30 0,46,573 0,01174 0,06220 0,00104 0,66 388,2 8,1 389 6,3 101 204 399 0,51
31 0,48,378 0,01556 0,06198 0,00107 0,54 400,7 10,7 387,6 6,5 102 115 254 0,45
32 0,46,449 0,01137 0,06197 0,00104 0,69 387,4 7,9 387,6 6,3 101 192 319 0,6
33 0,46,919 0,01373 0,06254 0,00107 0,58 390,6 9,5 391,1 6,5 101 305 418 0,73
34 0,48,080 0,01550 0,06229 0,00108 0,54 398,6 10,6 389,5 6,6 101 152 199 0,76
35 0,46,409 0,01281 0,06199 0,00105 0,61 387,1 8,9 387,7 6,4 101 98 273 0,36
36 0,46,771 0,01133 0,06241 0,00105 0,69 389,6 7,8 390,2 6,4 101 226 403 0,56
37 0,48,443 0,01621 0,06222 0,00109 0,52 401,1 11,1 389,1 6,6 102 209 330 0,63
R1: sample MS-8. LA-ICP-MS (IGM SB RUS, Novosibirsk, Russia)
1 0,46,279 0,01488 0,06178 0,00118 0,59 386,2 10,3 386,4 7,2 100 358 122 0,34
2 0,47,005 0,01461 0,06242 0,00119 0,61 391,2 10,1 390,3 7,2 100 317 135 0,43
3 0,46,202 0,01212 0,06154 0,00115 0,71 385,7 8,4 385 7,0 100 350 149 0,42
4 0,47,220 0,01654 0,06207 0,00119 0,55 392,7 11,4 388,2 7,2 101 253 85 0,34
5 0,46,907 0,01394 0,06209 0,00117 0,63 390,5 9,6 388,3 7,1 101 312 112 0,36
6 0,46,449 0,01861 0,06192 0,00122 0,49 387,4 12,9 387,3 7,4 100 254 78 0,31
7 0,46,955 0,01258 0,06237 0,00116 0,69 390,9 8,7 390,1 7,1 100 428 167 0,39
8 0,46,773 0,01666 0,06224 0,0012 0,54 389,6 11,5 389,2 7,3 100 535 175 0,33
9 0,46,433 0,01671 0,06196 0,00119 0,53 387,3 11,6 387,5 7,2 100 394 130 0,33
10 0,46,949 0,01384 0,06201 0,00116 0,63 390,8 9,6 387,8 7,1 101 362 174 0,48
11 0,46,945 0,01762 0,06196 0,00119 0,51 390,8 12,2 387,6 7,2 101 159 56 0,36
12 0,47,301 0,01362 0,06258 0,00117 0,65 393,3 9,4 391,3 7,1 101 386 203 0,73
13 0,47,170 0,01249 0,06274 0,00116 0,7 392,4 8,6 392,3 7,1 100 429 224 0,64
14 0,47,388 0,01866 0,06252 0,00122 0,5 393,9 12,9 391 7,4 100,7 263 139 0,52
15 0,46,299 0,01145 0,06178 0,00114 0,75 386,3 8 386,4 6,9 100 352 282 0,73
16 0,47,212 0,01301 0,06231 0,00116 0,68 392,7 9 389,6 7,1 100,8 249 147 0,62
17 0,46,870 0,01307 0,06239 0,00116 0,67 390,3 9 390,1 7,1 100,1 530 412 0,54
R2: sample MS-13. LA-ICP-MS (Gakushuin University, Japan)
1 0,47,126 0,01564 0,06116 0,00125 0,61 392,1 10,8 382,7 7,6 102 128 40 0,31
2 0,45,028 0,01411 0,05945 0,00119 0,64 377,5 9,9 372,3 7,2 101 154 58 0,37
3 0,44,671 0,01211 0,05939 0,00115 0,71 375 8,5 371,9 7,0 101 261 126 0,48
4 0,48,640 0,01486 0,06160 0,00123 0,66 402,5 10,1 385,4 7,5 104 153 56 0,37
5 0,47,731 0,01638 0,06299 0,00186 0,86 396,2 11,3 393,8 11,3 101 247 78 0,32
6 0,45,552 0,01577 0,06087 0,00180 0,85 381,1 11,0 380,9 10,9 100 241 98 0,41
7 0,48,376 0,01635 0,06182 0,00182 0,87 400,7 11,2 386,7 11,1 104 272 133 0,49
8 0,46,729 0,01498 0,06268 0,00183 0,91 389,3 10,4 391,9 11,1 99 466 177 0,38
9 0,52,926 0,01695 0,06342 0,00186 0,91 431,3 11,3 396,4 11,2 109 422 264 0,62
10 0,46,994 0,01565 0,06269 0,00184 0,88 391,1 10,8 392 11,2 100 316 132 0,42
11 0,48,584 0,01618 0,06294 0,00185 0,88 402,1 11,1 393,5 11,2 102 306 152 0,5
12 0,46,730 0,01523 0,06264 0,00183 0,9 389,3 10,5 391,7 11,1 99 386 203 0,52
13 0,46,369 0,01498 0,06202 0,00181 0,91 386,8 10,4 387,9 11,0 100 429 224 0,52
14 0,52,450 0,01644 0,06466 0,00189 0,93 428,2 10,9 403,9 11,4 569 569 344 0,6
15 0,50,739 0,01616 0,06777 0,00198 0,92 416,7 10,9 422,7 12,0 463 463 380 0,82
16 0,52,978 0,01784 0,06797 0,00200 0,87 431,7 11,8 423,9 12,1 257 257 86 0,34
R3: sample MS-32. LA-ICP-MS (IGM SB RUS, Novosibirsk, Russia)
1 0,46,943 0,00985 0,06275 0,00106 0,81 390,8 6,8 392,3 6,4 106 537 154 0,29
2 0,51,548 0,01000 0,06294 0,00105 0,86 422,1 6,7 393,5 6,4 106 2567 933 0,36
3 0,45,145 0,01158 0,06074 0,00103 0,66 378,3 8,1 380,2 6,3 129 447 146 0,33
4 0,49,630 0,01041 0,06315 0,00105 0,79 409,2 7,1 394,8 6,4 111 2068 1172 0,57
5 0,45,193 0,00911 0,05930 0,00097 0,81 378,6 6,4 371,4 5,9 108 3858 1310 0,34
6 0,49,839 0,01000 0,06171 0,00101 0,82 410,6 6,8 386 6,2 110 3578 1911 0,53
7 0,48,221 0,00893 0,06405 0,00107 0,9 399,6 6,1 400,2 6,5 94 2980 1298 0,44
8 0,48,426 0,01312 0,06407 0,00111 0,64 401 8,98 400,3 6,71 100,2 303 181 0,6
9 0,48,221 0,00893 0,06405 0,00107 0,9 399,6 6,12 400,2 6,5 99,9 2980 1298 0,44
10 0,48,995 0,01216 0,06485 0,00111 0,69 404,9 8,28 405,1 6,7 100 390 185 0,48
11 0,49,264 0,00932 0,06487 0,00108 0,88 406,7 6,34 405,2 6,56 100,4 2140 867 0,41
12 0,49,777 0,00956 0,06512 0,00109 0,87 410,2 6,48 406,7 6,58 100,9 2840 839 0,3
(continued on next page)
319M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

with a New Wave Research UP-213 laser ablation system, at the Institute
of Geology and Mineralogy (Russia, Novosibirsk). The analysis was ap-
plied to 20–50 single zircon grains, similar in morphology and size in
each sample, which were embedded in epoxy resin together with
Temora (Black et al., 2004)andPlešovice (Slama et al., 2008) standard
zircons. The zircon grains were cut off to about a half of their thickness
and polished. The dating points on the grain surfaces were selected
using BSE and CL images. Signal was tuned to maximum sensitivity for
Pb and U while keeping oxide production, monitored as
254
UO/
238
U,
well below 2% by ablation of NIST SRM 612 glass. The laser was operated
at 5 Hz using a spot size of 30 μmandalaserfluence of ~3 J/cm
2
.Data
were acquired for masses 202–204–206–207–208–232–235–238 in
low-resolution E-scan mode during a 25 s background measurement
followed by a 30 s sample ablation. The evaporated particles were carried
from the laser system to the mass spectrometer in a stream of pure He.
Data reduction was carried out using the “Glitter”software package
(Griffin et al., 2008). The laser induced elemental fractionation and
mass bias was corrected by normalizing the data to the Temora reference
zircon (Black et al., 2004). Resulting age interpretations, weighted mean
U-Pb ages and Concordia plots were constructed using Isoplot software
(Ludwing, 2003). All uncertainties reported are standard deviations at
the 1-sigma confidence-level. The Plešovice (Slama et al., 2008)zircon
standard was measured multiple times throughout each sequence as
an unknown for accuracy check. Mean concordia age obtained for
Plesovice standard yields 337.5 ± 3 Ma (n = 12), which is in accordance
with the reported ID-TIMS age of 337.1 ± 0.4 Ma (Slama et al., 2008).
For rhyolite samples MS-12 and MS-14 the U-Pb isotopic composi-
tions of zircon grains were analyzed on a VG PQ Excell ICP-MS equipped
with a NewWave Research UV213 laser ablation system in the Depart-
ment of Earth Sciences, the University of Hong Kong. The laser system
delivers a beam of 213 nm UV light from a frequency quintupled Nd:
YAG laser. Most analyses were carried out with a beam diameter of
30 μm, at a 6 Hz repetition rate. This gave a
238
U signal of 3 × 104 to
200 × 104 counts per second, depending on U contents. Typical ablation
time was 30–60 s, resulting in pits 20- to 40-μm-deep. Before measure-
ment, samples were ablated for 10 s to eliminate common lead contam-
ination on sample surfaces. In addition,
202
Hg was monitored to control
the isobaric interference of
204
Hg on
204
Pb. Data acquisition started
with a 15 s measurement of a gas blank during the laser warm-up
time. The
204
Pb signal was so small that the common lead correction is
therefore regarded as unnecessary (Xia et al., 2004). The standard zircon
91500 was used to evaluate the magnitude of mass bias and inter-
elemental fractionation. The instrumental settings and detailed analyti-
cal procedures are described in Xia et al. (2004). The U-Pb ages were cal-
culated using the U decay constants of 238 U = 1.55125 × 10
−10
year
−1
,
235
U = 9.8454 × 10
−10
year
−1
and the Isoplot 3 software (Ludwing,
2003). Individual analyses are presented with 1σerrors, and uncer-
tainties in pooled age results are quoted at the 95% confidence level (2σ).
4.2. Whole-rock compositions
The concentrations of major elements were measured by a CMP-25
Xray fluorescence (XRF) device according to the state standard of the
USSR Ministry of Geology (GOST 41-08-212-82) at the Institute of Geol-
ogy and Mineralogy (IGM, Novosibirsk). Trace elements were deter-
mined by the inductively coupled plasma mass spectrometry (ICP-MS)
on a FinniganElement ICP-MS analyzer, at IGM, following the protocols
of Jenner et al. (1990). Powdered samples were digested in a HF-HNO3
(2:1) mixture in a screw-top Teflon beaker for 2 days at ~100 °C, then
evaporated to dryness, refluxed in 6 N HCl and dried twice, and then
redissolved in 1 N HCl. The procedure was repeated till complete disso-
lution of the powder. The final solution was evaporated to dryness,
refluxed in 6 N HNO3, dried three times, and dissolved in 2% HNO3.
The wet chemistry analytical work was conducted under clean lab con-
ditions.The precision and accuracy of theanalyses were checked against
the BHVO-1 (Jenner et al., 1990), BCR-1 (Jochum and Nohl, 2008), and
JB-3 (Orihashi and Hirata, 2003) international standards and estimated
to be 2–7% for rare earth and high-field strength elements. The element
abundances and ratios were normalized to chondritic (e.g., La/Sm
n
)and
primitive mantle (PM) values (e.g., Th/Nb
pm
), with reference to Sun and
McDonough (1989) and McDonough et al. (1992), respectively. Repre-
sentative analyses of major and trace element contents (in wt% and
ppm, respectively) are given in Table 3.
4.3. Isotopes
Radiogenic Sm-Nd isotope studies were carried out for bulk samples
at the Geological Institute of the Kola Science Center (Apatity) on a
Finnigan MAT 262 (RPQ) seven-channel solid-state mass spectrometer.
Isotope ratios were normalized to
146
Nd/
144
Nd = 0.7219 and then
recalculated for the
143
Nd/
144
Nd = 0.511860 ratio assumed for the La
Jolla Nd standard, with a
143
Nd/
144
Nd weight average ratio of
0.511837 ± 12 (2σ), from 36 measurements. The εNd(t) values and
model ages T
DM
were calculated with reference to
143
Nd/
144
Nd =
0.512638 and
147
Sm/
144
Nd = 0.1967 (Jacobsen and Wasserburg,
1984) and the respective depleted mantle (DM) ratios of
143
Nd/
144
Nd = 0.513151 and
147
Sm/
144
Nd = 0.2136 (Goldstein and
Jacobsen, 1988). Two-stage age models T
DM-2
were obtained assuming
a middle crust ratio of
147
Sm/
144
Nd = 0.12 (Liew and Hofmann,
1998). Sm-Nd isotopic data are presented in Table 3; all errors quoted
in this paper are 2σ. All the εNd(t) values of the R1-rhyolites, and R2
and R3-rhyolites have been calculated using the average U-Pb age of
390 and 380 Ma, respectively.
5. Results
5.1. Petrography and mineralogy
The R1-type is mainly a porphyry rhyolite, rarely aphyric, commonly
yellowish, sometime mauve or reddish. They have a massive, some-
times striped texture, where discontinuous strips and lenses are com-
posed of spherulites (Fig. 5a–e). Within the same subvolcanic
intrusions of porphyritic varieties the central parts can be oligophyric
and aphyric. Phenocrysts of quartz and alkaline feldspar (from 0.2 to
2 mm) make up 1–2to10–40% of the rock volume. Quartz forms iso-
metric grains with melted edges; plagioclases are represented by
Table 1 (continued)
Point no Isotope rations Ages (Ma) Concentration (ppm)
207
Pb/
235
U2σ
206
Pb/
238
U2σRho
207
Pb/
235
U2σ
206
Pb/
238
U2σ% conc. U Th Th/U
13 0,49,483 0,00935 0,06512 0,00108 0,88 408,2 6,35 406,7 6,56 100,4 2686 1337 0,5
14 0,52,103 0,01114 0,06498 0,00109 0,78 425,8 7,44 405,8 6,57 104,9 2155 747 0,35
15 0,49,852 0,01566 0,06476 0,00112 0,55 410,7 10,61 404,5 6,8 101,5 255 98 0,38
16 0,49,509 0,00941 0,06484 0,00107 0,87 408,4 6,39 405 6,49 100,8 3129 875 0,28
17 0,50,803 0,01038 0,06485 0,00107 0,81 417,1 6,99 405 6,47 103 2459 671 0,27
18 0,49,917 0,00979 0,06473 0,00106 0,83 411,1 6,63 404,3 6,44 101,7 3436 1186 0,35
19 0,49,018 0,00969 0,065 0,00107 0,83 405 6,6 405,9 6,45 99,8 3578 1911 0,32
20 0,49,736 0,01013 0,06468 0,00106 0,8 409,9 6,87 404 6,43 101,5 3217 1068 0,33
320 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Table 2
Major oxides (wt%) and trace elements (ppm) in the MS-rhyolites.
Volcanic center Voronezh Karaulny Butochny
Phase 1 1 1 1 11111 1 111
№ms-25/2 ms-25/4 ms-26 ms-30/1 ms-4/1 ms-4/2 ms-4/4 ms-4/5 ms-10/1 ms-10/3 ms-6 ms-8 ms-15
SiO
2
76.73 78.87 86.3 77.26 78.99 76.93 79.45 81.7 76.42 77.48 79.32 79.12 77.09
TiO
2
0.22 0.13 0.15 0.17 0.17 0.18 0.14 0.02 0.16 0.1 0.02 0.04 0.13
Al
2
O
3
12.47 12 8.59 12.57 11.67 12.38 11.68 10.95 12.22 12.11 10.42 11.47 12.12
Fe
2
O
3
* 3.1 1.6 1.61 2.61 1.69 3.23 1.56 1.05 3.02 3.03 3.15 2.3 3.65
MnO 0.06 0.04 0.01 0.07 2.26 1.58 1.51 0.01 0.02 0.03 0.02 0.01 0.05
MgO 0.86 0.42 0.49 0.34 0.05 0.06 0.02 0.86 0.29 0.19 0.09 0.06 0.29
CaO 0.27 0.2 0.1 0.29 0.21 0.15 0.14 0.1 0.08 0.25 0.05 0.05 0.16
Na
2
O 3.83 3.83 0.18 4.8 4.82 3.64 2.6 4.57 3.79 3.78 3.41 4.27 2.75
K
2
O 2.41 2.9 2.53 1.86 0.12 1.83 2.88 0.73 3.96 3.01 3.5 2.67 3.74
P
2
O
5
0.04 0.02 0.04 0.03 0.01 0.02 0.01 0.01 0.04 0.03 0.02 0.02 0.02
LOI 1.35 0.9 1.87 1.15 1.34 1.98 1.59 0.97 1 0.92 0.38 0.59 1.34
Th 6.85 8.7 6.09 5.62 6.9 8.85 9.07 8.63 6.43 7.42 6.41 7.87 6.92
U 2.56 2.64 2.16 2.19 3.96 2.9 2.71 2.41 2.8 2.8 2.52 2.77 2.66
Ga 11 12 12 15 8 19 15 10 13 12 9 12 15
Rb 32 42 63 39 29 53 61 29 35 45 64 45 97
Ba 834 1286 388 338 220 594 572 205 283 406 694 520 443
Cs 0.11 0.19 0.65 0.37 0.11 0.47 0.43 0.22 0.11 0.11 0.26 0.22 0.67
Sr 107 111 14 88 56 58 52 74 53 223 65 74 181
La 32.1 37.8 25.4 21.4 25 20 22.8 30.1 47.8 25.4 28.7 19.7 30
Ce 62 66 55 51 49 40 49 62 110 49 56 33 56
Pr 7.9 9.1 6.8 5.5 6.8 5.6 6.9 7.9 11.5 7 6.9 5.4 6.9
Nd 31.2 35.6 28 22 26.1 20.9 25.6 27.9 45.5 25.1 26.9 19.1 26.9
Sm 7.7 8.1 5.9 5.1 5.7 4.9 5.7 5.4 10.5 5.4 6.1 4.3 6.3
Eu 1.3 0.99 1.08 1.25 1.22 1.08 0.86 0.81 1.43 1.3 1.07 1.08 1.19
Gd 7.7 8.56 5.5 4.92 6.76 6 6.14 5.53 10.22 5.94 6.01 4.53 6.78
Tb 1.29 1.48 0.9 0.84 1.26 1.14 1.17 0.96 1.53 1.05 1.05 0.87 1.24
Dy 7.54 9.23 5.45 5.17 8.22 7.63 7.58 5.84 8.94 6.42 6.44 5.78 7.06
Ho 1.7 1.93 1.14 1.14 1.6 1.65 1.65 1.22 1.79 1.33 1.3 1.22 1.56
Er 4.74 5.68 3.2 3.6 4.82 4.97 5.09 3.66 4.68 3.78 3.78 3.84 4.49
Tm 0.81 0.94 0.47 0.56 0.74 0.82 0.83 0.6 0.76 0.57 0.59 0.65 0.68
Yb 5.06 5.84 3.2 3.9 4.73 5.43 5.2 3.96 4.82 3.76 4.05 4.24 4.57
Lu 0.8 0.88 0.47 0.6 0.72 0.84 0.83 0.6 0.74 0.6 0.6 0.66 0.68
Zr 247 180 148 183 213 273 207 109 182 172 207 199 242
Hf 6.69 5.53 4.04 4.9 5.22 6.6 5.52 3.52 5.39 4.45 5.58 5.18 6.19
Ta 0.63 0.81 0.74 0.59 0.69 0.82 0.86 0.82 0.75 0.63 0.63 0.63 0.63
Nb 8.3 8.9 10 8.4 10.8 13 12.9 11.8 9.9 9.2 8 9.9 8.9
Y 48 56 313257564934 50 42 373846
V 1.2 0.4 10.4 1.2 15.9 69 0.4 17.1 1 0.4 0.9 28.8 0.7
Cr 13 17 25 20 12 17 –9 10 28 112 8 70
Co 1.4 0.9 1.2 0.8 –––– 0.9 1.2 1.7 1.1
Ni –– 10 9 53 ––59 29 37 36 45 39
Pb 7 9 31117122 4 12 6 4 4 7
Cu 29 45 16 16 20 16 17 19 17 20 25 17 18
Zn 170 173 26 80 187 108 66 29 93 93 77 77 84
Eu/Eu* 0.5 0.4 0.6 0.8 0.6 0.6 0.4 0.4 0.4 0.7 0.5 0.7 0.6
La/Yb
n
4 4 544235 7 5 534
Sr/Y 2.2 2 0.5 2.7 1 1 1.1 2.2 1.1 5.4 1.8 1.9 3.9
ΣРЗЭ 172 192 143 127 143 121 139 157 260 137 150 104 154
Volcanic center Sadovushka Kruchki Sadovushka Voronezh Butochny
Phase 1 1 1 1 1 1 2 2 2 3 3 3 3
№ms-1/2 ms-1/3 ms-12 ms-14/1 ms-14/2 ms-14/3 ms-13/1 ms-30/2 ms-16 ms-17 ms-19 ms-22 ms-32
SiO
2
76.2 75.53 77.17 79.1 76.21 76.44 78.1 84.95 77.41 77.8 78.22 78.53 82.68
TiO
2
0.15 0.21 0.19 0.11 0.17 0.13 0.02 0.06 0.11 0.04 0.03 0.06 0.04
Al
2
O
3
12.38 12.34 12.51 10.44 12.06 12.35 11.93 8.28 11.42 12.2 11.69 11.8 9.52
Fe
2
O
3
* 1.54 2.19 2.46 1.48 2.38 2.29 1.88 1.94 2.19 1.89 1.81 0.84 1.34
MnO 0.52 0.22 0.54 0.02 0.03 0.03 0.02 0.02 0.01 0.01 0.02 0.02 0.02
MgO 0.01 0.01 0.02 0.26 0.41 0.49 2.02 0.13 0.08 0.18 0.05 0.14 0.35
CaO 0.19 0.15 0.27 0.93 1.17 0.4 0.79 0.2 0.05 0.05 1.35 0.07 0.1
Na
2
O 1.28 1.63 4.56 1.33 3.23 4.14 3.34 4.17 0.14 1.97 2.65 1.3 0.12
K
2
O 7.72 7.71 2.26 6.31 4.33 3.7 1.87 0.22 8.58 5.84 4.15 7.24 5.81
P
2
O
5
0.01 0.01 0.01 0.02 0.02 0.02 0.03 0.02 0.01 0.02 0.02 0.02 0.02
LOI 0.2 0.82 0.72 0.38 0.52 0.77 2.03 0.46 0.69 0.88 0.54 0.43 1.07
SUM 100 100 100 100 100 100 100 100 100 100 100 100 100
Th 11.8 7.81 9.48 7.78 8.4 7.66 9.41 4.58 6.64 9.62 8.64 9.03 7.48
U 3.39 2.51 3.04 3.13 3.47 3.1 2.63 1.66 2.77 3.08 3.42 5 2.65
Ga 8232110192012 4 91316139
Rb 98 27 57 64 42 45 190 5 90 136 80 141 99
Ba 450 238 333 583 410 534 864 125 732 656 816 417 639
Cs 0.4 0.25 0.19 0.24 0.2 0.49 1.57 0.11 0.52 1.6 0.75 0.78 0.45
(continued on next page)
321M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

prismatic crystals more often with simple twins, less often with
polysynthetic twins; sometimes form glomerular coalescence. The alka-
line feldspar phenocrysts are slightly sericitized and sometimes re-
placed by hydromuscovite on micro-cracks. The groundmass of
quartz-feldspar sometimes recrystallized. Rare small flakes of
chloritized biotite with rutile inclusions along cleavages are observed.
The ore minerals are magnetite and titanomagnetite; and in subordi-
nate –by leucoxene. Xenocrysts are absent.
The R2-rhyolite easily diagnosed due to the light colors and aphyric
or aplite-like appearance (Fig. 5f–g). Relict porphyritic texture is ob-
served. Phenocrysts (1–3%) are represented by semi-decomposed pris-
matic crystals of alkaline feldspar (0.4–0.7 mm), with inclusions of
micro-lamellae rutile and leucoxene, and alteration of clay minerals
such as kaolinite. The groundmass is made of secondary-spherulitic
quartz-feldspar aggregates, black dusty particles of leucoxene (up to
1–3%), and yellowish-brown kaolinite.
The R3-rhyolite is more alterated than the other types. They have a
light yellowish colors, and glomeroporphyritic textures (Fig. 5h, i). The
rock is massive and heterogeneous, taxitic, flow-banded, festoon-
fluidized. Phenocrysts (from 0.2 to 2.5 mm), in an amount of 10–20%
of the rock volume are represented by quartz and feldspar. Quartz
forms isometric less often comminuted grains with melted edges. The
feldspar phenocrysts are intensively exposed to secondary processes.
Etching shows spotted K-feldspathization at the periphery and along
the micro-cracks, often in the form of solid microcline lattice or in a
staggered manner on the central part. There is a germination of small
differently oriented prisms of albite. Felsitic texture in combination
with micro-poikiloblastic, relict- and secondary-spherulitic textures
are dominated. These rocks are characterized by wide development of
chloritized biotite flakes with rutile secretions along cleavage. Second-
ary alterations are expressed in K-feldspatization, silicification,
spotted-banded sericitization, weak chloritization. In some cases,
there are magmatic breccias of rhyolites, resembling automagmatic
breccias. Captured material is represented by rhyolites xenoliths and
xenocrysts of quartz and feldspars (0.25–0.75 mm).
5.2. Zircon U-Pb geochronology
U-Pb dating of zircons from samples of the R1-rhyolites (MS-26, MS-
10/1, MS-8, MS-12, MS-14/1), R2-rhyolites (MS-13/1), and R3-rhyolites
(MS-32) provided time constraints on the emplacement of the
subvolcanic intrusions of the Melnichno-Sosnovsky volcanic complexes
in the NW Rudny Altai. The
206
Pb/
238
U ages of the R1-rhyolite were ob-
tained for six samples of main phase of the Voronezh, Karaulny,
Butochny, Sadovushka and Kruchki volcanic centers. The zircon grains
of sample №MS-26 (Voronezh) are colorless or pale yellow euhedral
prismatic crystals (80–350 μm, KL =2–3), with oscillatory zoning in
CL-images, indicating magmatic origin (Fig. 6a). A weighted mean
206
Pb/
238
U age of 388 ± 2 Ma (MSWD = 0.12) was obtained for 37 an-
alytical spots in different partsof zoned magmatic crystals (Nos. 1–37 in
Table 1) and interpreted as the age of emplacement (Fig. 6f). The zircons
of №MS-10/1 (Karaulny) are transparent, translucent, colorless, pale
yellow and pale brown euhedral prismatic crystals (150–270 μm,
KL =2–2.5) with thin oscillatory zoning (Fig. 6b). A weighted mean
206
Pb/
238
U age of 389 ± 3 Ma (MSWD = 0.06) was obtained for 12 an-
alytical spots in different parts of zoned magmatic crystals (Nos. 17–28
in Table 1) and interpreted as the age of emplacement (Fig. 6g). How-
ever, 16 analyses for the same sample (Nos. 1–16; Table 1)gave
weighted mean
206
Pb/
238
U age of 381 ± 2 Ma (MSWD = 0.44). Since
these rhyolites are the same as R1-samples; and the zircons from this
sample were separated in one batch with the zircons from R2-
rhyolite, we assume that clogging of this sample could occur. If we cal-
culate the total weighted average age obtained for all zircon grains
from this sample, it yields a value of ~384 Ma, which contradicts the
stratigraphic data, since at this point there was a volcanic pause in the
NW Rudny Altai (Fig. 3;Murzin et al., 2001). On the other hand, the
Table 2 (continued)
Volcanic center Sadovushka Kruchki Sadovushka Voronezh Butochny
Phase 1 1 1 1 1 1 2 2 2 3 3 3 3
№ms-1/2 ms-1/3 ms-12 ms-14/1 ms-14/2 ms-14/3 ms-13/1 ms-30/2 ms-16 ms-17 ms-19 ms-22 ms-32
Sr 33 36 35 129 158 112 41 96 50 79 144 69 21
La 27.5 24 42.5 26.1 40.8 32.2 18.3 15.4 19.2 19.9 22.7 23.6 13.5
Ce 56 57 81 65 85 71 34 30 37 41 41 40 32
Pr 7.9 8.9 10.5 8.3 10.4 9 4 3.4 4.3 5.2 5.5 5.7 3.7
Nd 28.7 36.1 40.4 32.3 42 38.3 15 12.7 16 20 20.9 21.6 15.1
Sm 6.8 8.7 9.4 7.4 9.3 9.7 3.1 2.7 3.4 5 5.3 5.6 3.7
Eu 0.9 1.63 1.36 0.83 1.51 2.07 0.44 0.27 0.28 0.31 0.47 0.37 0.26
Gd 6.93 9.89 10.4 6.92 9.12 9.93 3.54 2.75 3.57 4.65 5.21 6.48 3.79
Tb 1.32 1.89 1.72 1.1 1.57 1.69 0.67 0.5 0.68 0.72 0.93 1.2 0.7
Dy 8.36 12.17 10.9 6.93 9.67 9.78 4.73 3.41 4.07 3.86 5.35 7.45 4.56
Ho 1.8 2.58 2.36 1.44 2.17 1.97 1.08 0.75 0.91 0.84 1.14 1.62 0.97
Er 5.44 7.95 6.73 4.08 6.94 5.48 3.17 2.2 2.53 2.4 3.2 4.47 2.82
Tm 0.91 1.28 1.1 0.68 1.21 0.87 0.48 0.34 0.4 0.38 0.51 0.72 0.45
Yb 5.86 8.26 7.24 4.6 8.15 6.4 3.2 2.24 2.77 2.31 3.36 4.57 3.1
Lu 0.9 1.26 1.11 0.69 1.18 0.94 0.51 0.33 0.4 0.36 0.51 0.71 0.47
Zr 217 378 245 179 265 249 171 56 59 78 68 72 59
Hf 6.25 9.46 7.93 6.08 8.25 7.08 5.04 1.78 2.3 3.22 2.98 3.07 2.4
Ta 0.73 0.82 0.81 0.77 0.77 0.77 0.66 0.55 0.78 1.22 1.14 1.2 0.77
Nb 10.9 15.3 9 11.7 12.8 11 7.3 6.7 8.2 13.4 11.6 12.1 10.4
Y 52777331535030 26 2521314827
V 32.3 23 1 3.8 3.1 1.9 3 3.7 1.1 2.1 0.8 0.7 1.5
Cr 15 16 25 54 33 19 22 60 59 582 29 31 44
Co –– 0.4 2 1.5 0.8 0.5 1.6 1.5 4.9 1 1.1 1.5
Ni 11 54 39 30 16 10 23 39 28 64 8 21 22
Pb 109 76953 20 41214165
Cu 16 19 18 18 18 18 29 24 19 32 14 19 17
Zn 30 115 82 9 42 80 56 15 30 2048 44 29 28
Eu/Eu* 0.4 0.5 0.4 0.3 0.5 0.6 0.4 0.3 0.2 0.2 0.3 0.2 0.2
La/Yb
n
32 44334 5 56533
Sr/Y 0.6 0.5 0.5 4.1 3 2.2 1.4 3.7 2 3.9 4.6 1.4 0.8
ΣРЗЭ 160 182 227 166 229 199 92 78 96 107 116 124 85
322 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

obtained variations of
206
Pb/
238
U isotopic ages are still within the accu-
racy of the LA-ICP-MS method used. Zircon grains of sample №MS-8
(Butochny) are euhedral prismatic crystals (70–120 μm, KL =
1.5–2.0), translucent, colorless or pale yellow. CL-images reveal thin os-
cillatory magmatic zoning (Fig. 6c). The weighted mean
206
Pb/
238
Uage
for 17 analyses (Nos. 1–13; Table 1;Fig. 6h) is 389 ± 3 Ma (MSWD =
0.4). The zircon grains of sample №MS-12 (Sadovushka) are translucent
colorless or with a yellowish tint euhedral prismatic crystals
(70–110 μm, KL =1.5–2), with oscillatory zoning in CL-images
(Fig. 6d). Fifteen analyses yield a weighted mean
206
Pb/
238
U age of
389 Ma (MSWD = 0.01) for different parts of zoned magmatic crystals
(Nos. 1–5; Table 1) and interpreted as the age of emplacement (Fig. 6i).
The zircon grains of sample №MS-14/1 (Kruchki) are colorless or pale
yellow euhedral prismatic crystals (70–200 μm, KL =1.5–4), with oscil-
latory zoning in CL-images (Fig. 6e). A twelve sub-concordant analyses
yield a weighted mean
206
Pb/
238
U age of 391 Ma (MSWD = 0.03) for
different parts of zoned magmatic crystals (Nos. 1–12; Table 1)and
interpreted as the age of emplacement (Fig. 6j). The zircon grains of
sample №MS-13/1 (Sadovushka) are transparent-translucent, colorless
or slightly yellowish euhedral prismatic and isometric crystals
(90–200 μm, KL =2–3), with oscillatory zoning in CL-images (Fig. 7a).
A weighted mean
206
Pb/
238
U age of 381 Ma (MSWD = 0.86) was ob-
tained for 13 analytical spots in different parts of zoned magmatic crys-
tals (Nos. 1–13; Table 1) and interpreted as the age of emplacement
(Fig. 7c). Additional 3 analyses gave a weighted mean
206
Pb/
238
U age
of 422 Ma (MSWD = 0.49) interpreted to record the age of inherited
zircons (Nos. 14–16; Table 1).
The zircon grains of sample №MS-32 (Butochny) are translucent
colorless or slightly yellowish euhedral prismatic and isometric
crystals (80–170 μm, KL =2–4), with oscillatory zoning in
cathodoluminescence (CL) images (Fig. 7b). A weighted mean
206
Pb/
238
U age of 381 Ma (MSWD = 2.2) was obtained for 6 analytical
spots in different parts of zoned magmatic crystals (Nos. 1–6; Table 1)
and interpreted as the ageof emplacement (Fig. 7d). Thirteen additional
analyses gave a weighted mean
206
Pb/
238
U age of 404 Ma (MSWD =
0.09), which may record the age of xenocrytic or inherited zircons
(Nos. 7–19; Table 1).
5.3. Mobile element geochemical systematic
Twenty-eight samples of rhyolites have been analyzed for major-
and trace-element abundances and results are listed in Table 2. Loss
on ignition mostly ranged from 0.2 to 2% and the major elements com-
positions have been recalculated on an anhydrous basis to 100 wt%. Al-
though rhyolite samples were taken away from the ore bodies to avoid
the possible effect of secondary mineralization, the rhyolites still show
chemical vagaries due to feldspathization (enrichment in soda or pot-
ash) and silicification (enrichment in silica).
Most of the MS-rhyolites have very high contents of SiO
2
(75.5–79.4 wt%), low Al
2
O
3
(8.3–12.5 wt%) and very low CaO
(0.05–1.35 wt%) and MgO (0.01–0.9 wt%). These rocks have a wide var-
iation in the amount of Na
2
O+K
2
O(2.7–9.3 wt%), and formally have
affinity to sub-alkaline series, in the TAS diagram (Fig. 8a; Le Maitre
et al., 1989); and to low-K tholeiite and medium-K calc-alkali series, in
the SiO
2
-K
2
O diagram (Fig. 8b; Peccerillo and Taylor, 1976). This deter-
mines their division into two groups: most R1 (n = 15) and R2-
rhyolites (n = 2) have a weak predominance ofsodium over potassium
(Na
2
O = 2.6–4.8 wt%; K
2
O = 0.1–4.3 wt%); and in contrast, all R3-
rhyolites (n = 6) and some samples from the first group have a pre-
dominance of potassium over sodium (Na
2
O = 0.1–2.6 wt%; K
2
O=
4.1–8.6 wt%). In addition, these rock have a predominantly high Fe*
Index (Fe* =FeO
T
/(FeO
T
+ MgO); Frost et al., 2001;Fig. 8c) from 0.88
to 1.00 (n = 15); and extremely strongly peraluminous values of the
Aluminum Saturation Index from 1.2 to 5.6 (ASI =Al/(Ca–
1.67P + Na + K) molar ratio; Chappell and White, 1992;Fig. 8d).
In fact, MS-rhyolites consist of relatively immobile Si and Al, which
do not vary greatly in amount, and mobile K and Na components,
which have wide variations (Table. 3). Some strongly altareted rhyolites
have extremely felsic composition(SiO
2
= 84.9–86.3 wt%) or extremely
strongly contrast Na
2
O/K
2
O (0.02 to 39.6 wt%) ratios. From the entire
rocks, only three R1-samples (ms-10/1; 14/2 and 14/3) have the lowest
SiO
2
~ 76 wt% and alkalis from 6.2 to 7.8 wt%, which are probably the
closest reflection of the primary composition of these rocks. In term of
major elements, this causes difficulties and can give confusing conclu-
sions on the classificationof such high-siliceous rocks, with sodium/pot-
ash dichotomy; given also the absence of lessfelsic varieties with which
they could form a single series suitable for analysis and identification.
5.4. Alteration effect
The mobile-element variations in the volcanic rocks of submarine
environment are a result of spilite-keratophyric alteration –mass ex-
change components due to active fluid-magma/rock interactions, due
to magma reaction with down-welling cold sea water and/or out-
flowing hot hydrothermal fluids. Essentially this is the result of
chloritization (Mg-Fe alterations), and albitization (Na-K exchange re-
actions) –dramatic increase in normal albite balance in hydrolysis of
feldspar, accompanies by relatively addition or removal of silica and alu-
mina (Alt and Teagle, 2003).In previous works, it was noted that (i) the
felsic effusive of the Rudny Altai are not a particular type of ultra-felsic
and ultra-sodium rocks, but were subject to intense secondary alter-
ation pronounced in silicification and albitization; and (ii) although
the volcanic rocks of the Rudny Altai were formed in submarine condi-
tions, albitization was subject not only to effusions, but also to Devonian
granites, which excludes the major role of sea water in this process
(Chernov, 1974). Thus, subvolcanic intrusions with MS-rhyolites proba-
bly could not interact directly with seawater.
Based on our data, only R1-rhyolite could be effected to some so-
dium enrichment by fluid-related alterations, as evidenced by the posi-
tive correlation between Na and LOI, and ASI Index, and the high values
of this index in samples with high Na contents (Fig. 9a). In contrast, for
high-K R3-rhyolite, there is not relationship between ASI Index and Na,
and between LOI and Na, suggesting that they are free of sodium fluid-
related alterations. On the one hand, the “Igneous Spectrum”diagram
(Hughes, 1972), reflecting K/Na variations in spilite-keratophyric
rocks, the R1-rhyolite plot in the field of weakly altered rocks; while
all R3- and some of R2-rhyolites are in the field of compositions modi-
fied by potassic metasomatism (Fig. 9b). Also it corresponds to discrim-
ination of compositions on the “Alteration Box Plot”diagram (Fig. 9c;
Large et al., 2001), reflecting the geochemical trends of diagenetic and
hydrothermal alterations in the rocks, which shows most R1- and R2-
rhyolites are in the field of least altered rhyolites; and, in contrast, the
R3-rhyolites are the boundary between diagenetic and hydrothermal
alterations fields. On the other hand, for R3-rhyolites there is no direct
relationship between K
2
O and LOI; and LOI values are not the highest
(Fig. 9a; Table. 2). This suggests a lack of relationships between the
high K
2
O contents in these rocks and any potash enrichments, associ-
ated with exposure to hydrothermal fluid on the solid rock. Our analysis
shows that the strongly altered rocks from are only six samples of R1-
and R2-rhyolites, with the highest SiO
2
contents and extreme Na/K ra-
tios (Table. 2). On the “Alteration-Box-Plot”diagram, they lie outside
the last change of least altered rhyolitic field and show a clear dichot-
omy of compositions into soda- and potash-rich varieties (Fig. 9c).
Clearly, the wide variations in the petrochemical compositions of
strongly altered rocks are not a primary igneous feature, but instead
are a result of intense secondary alteration that occurred most likely
at the post-magmatic stage.
The effects of alkaline-fluid alterations are also shown by variations
of some other major and trace elements. Of course, the fluid-mobile el-
ements (e.g. Ba, Rb, Cs and Sr) concentrations could be strongly modi-
fied in the increasing degree of feldspathization process because they
323M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

replace Na, K, and Ca, due to their mobility during hydrothermal alter-
nation and metamorphism (Ayers et al., 2012;Wilke et al., 2012). For
example, Rb/Sr ratios for most MS-rhyolites are very low from 0.05 to
0.9, reaching ~4.5 in only three samples. However, as indicated by vari-
ation of the Na and K, for weakly altered samples the values of elements
such as Al, Ti, Ga, Ga/Al, and Zn are probably modified only marginally.
For this study, it is relevant that the aqueous solubility of Zr may in-
crease with the addition of Na-Al-Si-bearing hydrothermal fluids, and
it can play an important role for mobilization of HFSEs and change of
the isotopic system during fluid-rock interaction (Ayers et al., 2012;
Wilke et al., 2012). Although most of the R1- and R3-rhyolites have
wide variations of Na/Al ratios reflecting relatively modification of the
compositions by alkaline fluid, nevertheless they are observed in a nar-
row range of Zr contents, which do not have a clear correlation with Na/
Al ratios (Fig. 9d). The exception is the four strongly altered potash-rich
R1-samples on the left side of this diagram, which have low Na/Al ratios
correlating with Zr values.
The rare earth elements (REE), and even more so Eu, could also be
mobile during intense hydrothermal or low-grade metamorphic alter-
ation; and light REE (LREE) are considered to be more susceptible to
mobility than heavy REE (HREE) under most conditions (Polat and
Hofmann, 2003). However, as shown by Fig. 9e, for weakly altered R1-
Table 3
Sm-Nd isotopic compositions for representative samples of the MS-rhyolite.
Sample Volcanic center Type Sm Nd
147
Sm/
144
Nd
143
Nd/
144
Nd Err (2σ)E
(Nd)
0E
(Nd)
-t T
DM
(Ma) T
DM2
(Ma)
ms-26 Voronezh I 5,27 26,75 0,119,157 0,512,504 9 −2,61 1,25 1044 1050
ms-10/1 Karaulny I 11,63 53,1 0,130,190 0,512,622 7 −0,31 3,01 966 903
ms-8 Butochny I 6,52 29,2 0,134,932 0,512,648 14 0,20 3,28 974 880
ms-12 Sadovushka I 4,61 20,59 0,135,373 0,512,701 20 1,23 4,29 877 795
ms-14/2 Kruchki I 7,52 31,27 0,145,409 0,512,700 16 1,21 3,70 1008 838
ms-32 Butochny III 3,68 14,86 0,149,607 0,512,449 9 −3,69 −1,41 1669 1263
Fig. 5. Micro-photos of representative rhyolite samples. R1-rhyolite (a–e); R2-rhyolite (f–h); and R3-rhyolite (i–l). Pl = plagioclase, Kfsp = K-feldspar, Qtz = quartz, Bt = biotite. For
details of data selection see Section 5.1.
324 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

325M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

rhyolites, there is no correlation between La/Sm and Gd/Yb ratios, and
the values of “Alteration Index”(AI;Lentz, 1998). In addition, the REEs
(Fig. 9f) and HFSEs (Fig. 10) for all R1-rhyolite show linear trends
with Zr contents, and as shown below, there is a coherent character of
chondrite-normalized REE-patterns, attesting to their relative immobil-
ity. As for R3-rhyolites, our analysis clearly shows that wide variations of
La/Sm ratios in these rocks were clearly not related to either
feldspathization or chloritization, as their compositions form a
subvertical trend on the La/Sm vs AI diagram (Fig. 9e). There is no
clear correlation between Zr and REEs; but their compositions are not
scattered randomly on the diagram, as if they were secondary modified.
The observed distribution of compositions most closely resembles the
magmatic trend of differentiation. In general, we suggest that for most
MS-rhyolites their REEs and HFSEs were relative immobile during hy-
drothermal alterations, and we thus use these elements to identify the
original geochemical type of these rocks.
5.5. Immobile element geochemical systematics
Firstly, on the Nb/Y-Zr/TiO
2
diagram (Winchester and Floyd, 1977),
rocks of this study demonstrate affinity to a “normal-type”rhyolite, al-
though several samples are on the border of the field comendite/
pantellerite (Fig. 10a). When incompatible elements, such as REEs, Zr,
and Y are used, the MS-rhyolites are characterized by transitional compo-
sitions between calc-alkaline and tholeiitic series: (La/Yb)
n
~3–6
(Fig. 10b; Lentz, 1998; C1 chondrite values of Sun and McDonough,
1989); and Zr/Y ~ 4 при Zr b350 ppm (Fig. 10c); and La/Sm ~ 0.55–1
(Fig. 10d; Lentz, 1998;Barrett and MacLean, 1999). On the discrimination
diagram of La/10-Y/15-Nb/8 (Cabanis and Lecolle, 1989), the R1-rhyolite
is marked on the boundary between the field of the arc-related calk-
alkaline series and the within-plate series; and R2- and R3-rhyolites
have a geochemical affinity with the within-plate series (Fig. 11).
The relationships between Nb, Y, Zr and Ce, and ratios of Ga/Al
(1.26–3.52), and Zn (9–187 ppm) concentrations, show that the R1-
rhyolite has continuous transitional characteristics between the OGT
(unfractionated M-, I-andS-granite) and A-type felsic rocks (Fig. 12a-
f; Whalen et al., 1987). In contrast, the R2- and R3-rhyolites are marked
only in the OGT-field, since the total Nb, Y, Zr and Ce (119–242 ppm),
and Ga/Al (0.51–1.39) values; and Zn (52–171 ppm) contents are rela-
tively low. In the Nb vs. Y and Rb vs. Y + Nb tectonic discrimination di-
agrams (Pearce et al., 1984), all of the rhyolites are transition between
volcanic-arc (I-type) and within-plate (A-type) fields (Fig. 12g-h;
Whalen et al., 1987). Although the variations in Rb are coherent with
K
2
O, we neglected it in this case because the reason for that the rhyolites
compositions plot on the border between the fields of normal arc-
related and within-plate felsic magmas is due to Y + Nb (32–92 ppm)
and Ta + Nb (7–15 ppm) abundances. All rhyolites possess low Rb
(b200 ppm) contents, which is incompatible with a syn-collision setting
(Fig. 12h). Consequently, some of the R1-rhyolite bear the distinctive
geochemical signature of A-type felsic magmas, such as: (i) total enrich-
ment in Zr, Nb, Y and Ce (N350 ppm), Zr (N250 ppm), and high Ga/Al
values (N2.6; Whalen et al., 1987). In general, if we consider all MS-
rhyolites, they havetransitional geochemical characteristics that resem-
ble those of felsic rocks of the subduction-related extension geodynamic
settings. However, this magmatism was unlikely to be associated with
the rebound of the oceanic plate following slab break-off and exten-
sional detachment of subduction orogen, like the Cordillera of North
America (Whalen and Hildebrand, 2019); Fig. 12j.
Secondary, analysis of geochemical patterns shows that all rhyolites
can be divided into two groups, which corresponds to the main and late
phases of magmatic activities. The difference in REEs concentration and
the degree of Eu anomaly are distinctive feature of these rhyolites
(Fig. 13). One group (R1-rhyolite) has higher REEs (104–260 ppm)
concentration, generally show coherent spectra with more fractionated
REEs ((La/Yb)
n
=1.57–6.69), nearly flat to positive HREEs patterns
((Gd/Yb)
n
=0.81–1.71), and moderate negative Eu anomalies (Eu/
Eu* = 0.35–0.76; Eu* = 2Eu
n
/(Sm
n
+Gd
n
)), in the chondrite-
normalized REEs diagrams (Fig. 13a; C1 chondrite values of Sun and
McDonough, 1989). In contrast, another group (R2- and R3-rhyolites)
has relative low REEs concentration (78–124), less fractionated REEs
((La/Yb)
n
= 2.95–5.82), and nearly flat to negative HREEs patterns
((Gd/Yb)
n
=0.38–1.52), along with a pronounced negative Eu anomaly
(Eu/Eu* = 0.18–0.4); Fig. 13c. For all of the rhyolites, the N-MORB-
normalized multi-element patterns (Sun and McDonough, 1989) are
characterized by sub-parallel spectra and homogeneous with promi-
nent and uniform enrichment in Rb, Ba, Th, U and negative anomalies
in Nb, Ta and Sr, and with slight to moderate variations in element
abundances. In general, the rhyolites of the first group are more
enriched in Zr, Hf, and Zn, Ni and V, and the rhyolites of the second
group are enriched in Rb, Ba and Pb, Cr, although these regularities are
not systematic. As noted in manyworks, some of the geochemical signa-
tures, such as depletion in HFSEs: Nb, Ta and Ti relative to LILEs and
LREEs, it is a common feature of felsic magmas formed in convergent
margin settings, and derived from a crystal source, which itself was de-
rived from arc crust (e.g. Whalen et al., 1996).
Thirdly, although wehave summarized thegeochemical characteris-
tics of MS rhyolites, their initial geochemical type is still difficult to iden-
tify. As suggested to us Prof. Bruno Scaillet “All the types discussed here
are metaluminous”. (i) Obviously, the parental magmas of the MS-
rhyolites were hardly peralkaline in nature, as was observed from
their Zr/Ti and Nb/Y ratios (Fig. 10аand Fig. 12j). (ii) It is also unlikely
that these magmas were peraluminous type, because they have A-
type evolutionary trend (Fig. 12f), and are clearly different from highly
fractionated S-andI-type granites, that would increase their Ga/Al ra-
tios при crystal differentiation (Wu et al., 2017). (iii) As discussed
below, the R2- and R3-rhyolites are somewhat resemble high-silica re-
sidual and non-equilibrium melts associated with A-type magmas
(Barboniand Bussy, 2013); Fig. 12f. In general, this indicates that the pa-
rental magmas of the MS-rhyolites should have been metaluminous.
Fourth, in comparison with the bimodal-type extensional rhyolites
of different subduction-related environments (Fig. 13a), the R1-
rhyolites REE-patterns have a close resemblance to those of felsic volca-
nic rocks of intra-oceanic back-arc basin, as shown by the example of
the Izu-Bonin convergent system (Tamura et al., 2009). They show
even greater geochemical affinity to the extensional rhyolites of the bi-
modal association of the Coastal Range in the Central Chile, related to
Late Triassic intra-continental rifting in the Pacific Gondvana margin
(Fig. 13b; Morata et al., 2000). There is also a geochemical similarity to
some of the bimodal-type rhyolites of Taupo Volcanic Zone, New
Zealand (Deering et al., 2011), although the latter are mostly character-
ized by more depleted concentrations of MREEs and HREEs (Fig. 13b).
As for the R2- and R3-rhyolites, it have the so-called seagull-like form
of REE-patterns, which, in the simplifying assumptions, resemble
those of some fractionated granites (FG)andA
2
-type granites attributed
to post-orogenic extension tectonic setting (Chappell and White, 1992;
King et al., 1997;Chappell, 1999). However, all of the MS-rhyolites are
clearly distinguished from post-orogenic granites, by a general deple-
tion in REEs, lower HFSEs (e.g. Nb ~ 7–13 ppm; Th ~ 5–10 ppm) con-
tents, and less pronounced Eu anomaly (Fig. 12c, d), for example, if we
compare them with A
2
-granite from Southern Altai Range, China
(Shen et al., 2011). By the same geochemical characteristics, they differ
from bimodal-type rhyolites associated with initial rifting, as in Miocen
ridge-subduction-related California-type environments (Johnson and
O'Neil, 1984;Cole and Basu, 1992). Compared to R2- and R3-rhyolites,
some bimodal-type rhyolites from Eastern Tianshan Belt in Northwest
China (Chen et al., 2011) have a similar form of REE-patterns, reported
Fig. 6. (a–e)Representativecathodoluminescence images(CL) of zircons of R1-rhyolites. (f–j)U-Pb dating resultsand concordia diagrams for the selectedages. For detailsof data selection
see Section 5.2.
326 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

for the initial post-orogenic rifting settings (Fig. 13d); and Okinawa
Through rhyolites are associated with rifted-arc Ruykuy (Shinjo and
Kato, 2000); and Kuroko Rift rhyolites are associated with volcanic ac-
tivity in the NE Honshu back-arc system (Yamada et al., 2012). Our anal-
ysis shows that there is some regularity in the affinity of REEs
characteristics of MS-rhyolites with those of the bimodal-type rhyolites
related with the modes of extension of the continental type of the lith-
osphere, as in mature/marginal island-arc and continental margin, and
in some post-collisional settings.
5.6. Isotopes
Most of the MS-rhyolites have elevated values of
143
Sm/
144
Nd ~ 0.13–0.14, compared to values for the average continental
crust (~0.12; Liew and Hofmann, 1998); and εNd(t) from +3.0 to +4.3,
and model ages of T
Nd
(DM-2) ~ 0.79–0.9 Ga (Fig. 14;Table. 3). Coherent
normalized REE patterns and lack of correlation withthe variations of Na/
Al (inset to Fig. 14), indicate that secondary alteration have not signifi-
cantly changed their Nd isotopic composition. The exception is the
strongly altered rhyolite №MS-26 (Tables 2, 3), which has values εNd
(t) = +1.25 and model age of T
Nd
(DM-2) = 1.05 Ga. Since this sample
has very low Na/Al values, the modification of its isotopic composition
was probably due to post-magmatic potassium metasomatic alterations.
6. Discussion
6.1. Rocks source
Firstly,it is suggested thatthe main sources producing felsic magmas
of bimodal association in the subduction-related extensional
environments may be: (i) amphibolitic crust of ancient basement,
such as the Pacific convergent margin of the North American plate
(Piercey et al., 2008); (ii) newly formed basalt-andesitic crust, as for ex-
ample, from Izu-Bonin intra-oceanic subduction system (Tamura et al.,
2009); or (iii) the greywackes or previously emplaced igneous rocks,
like in Taupo Volcanic Zone, New Zealand (e.g. Deering et al., 2011). In
term of the experimental constraints on the formation of silicic magmas
(see review in Scaillet et al. (2016)), metaluminous melts are typically
originated from igneous sources, either via fractional crystallization of
mafic magmas or by partial melting of older, crustal metaluminous in-
trusive during which biotite and amphibole breakdown takes place.
The absence of adakites signature (e.g. Sr/Y b40 and Yb N1.8 ppm;
Xiao and Clemens, 2007) in the MS-rhyolites suggests that their parent
magma was not in equilibrium with a garnet-bearing residue. During
partial melting, garnet is not stable in residue under pressures below
8 kbar for tonalitic (Rutter and Wyllie, 1988;Watkins et al., 2007)and
6 kbar for volcanoclastic (Castro et al., 2000)sources.
Secondary, it may seem that the moderate isotope signatures (εNd
(t) from 3.0 to 4.3; Table 3;Fig. 14) and relatively young Nd model
ages (mostly 0.8–0.9 Ga) of the MS-rhyolite indicate possible presence
of mixed mantle and crustal components, for example, mixing/mingling
of mafic and felsic magmas caused by basaltic underplating (Wilson
et al., 2005). However, the stratigraphic position of the mafic
volcanogenic-sedimentary rocks filling the Shipunikha rift clearly indi-
cates their generation after the formation of subvolcanic intrusions of
MS-rhyolites (Murzin et al., 2001). Thus, the partial melting of under-
plating new mafic crust, in the form of known Devonian basic rocks,
could not play an important role. In addition, trend of the zircon-
saturated crystallization for the MS-rhyolites differs from the trend of
high-TI-type felsic magmas produced by the partial melting of mafic
Fig. 7. (a, b) Representative CL-imagesof zircons of R2- and R3-rhyolites, respectively.(c, d) U-Pb dating results and concordia diagrams for the selectedages. For details of dataselection
see Section 5.2.
327M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

metaigneous source (Fig. 16a; King et al., 1997); and MS-rhyolites are
clearly enriched in LREEs with respect to metabasalt-derived felsic
magmas when compared to intra-oceanic arc-related bimodal-type
rhyolites (Fig. 13a; Tamura et al., 2009). Note also that in addition to
hybrid isotope values in the MS-rhyolite, and later tuffs of mixed com-
position, there is no clear geological evidence of mantle-crust interac-
tion in the Rudny Altai. We suggest that the isotope compositions for
the rhyolites could be related to the involvement of immature sedi-
ments in their petrogenesis, which have juvenile isotopic compositions.
Volcanoclastic and continental detritus sediments were widespread
along the continental margin, which consists of biotite, amphibole, pla-
gioclase, and quartz and may have been brought to the magma sources
of granitoids (Vielzeuf et al., 1990). Fig. 14 shows that the isotopic com-
position of MS-rhyolites are closest to the field of evolution of Nd isoto-
pic composition for the Early Devonian (our data) and Caledonian
volcanic rocks (Kruk et al., 2010;Kruk, 2015), which, in the form of clas-
tic material, could be transported from the Altai Mountains (Hu et al.,
2000), and sedimented on the continental shelf prior to Devonian
volcanism.
6.2. Zircon saturation
The R1-rhyolite is characterized by enrichment in Zr
(109–378 ppm), and there are no inherited zircon crystals (see
Section 6). The high Zr content is more likely to suggest partial melting
of the upper crustal materials (Zr ~ 160–240 ppm; Rudnick and Gao,
2003), than formation due to mantle-related source. As discussed
above, the MS-rhyolites belong to the “normal”type felsic volcanic
rocks. This means that the enrichment in Zr does not come from the
peralkaline nature of their magmas (Scaillet et al., 2016). We suggested,
that the absence of ancient zircons is indirectly confirmed the high zir-
con saturation temperatures (Watson and Harrison, 1983;Clemens
et al., 1986;Eby, 1990;Boehnke et al., 2013). At dissolution of accessory
phases during the melting processes, melts gain heat producing and in-
compatible elements is facilitated in A-type granitic magmas because
they are high in temperature and largely molten (King et al., 1997).
For example, Zr vs. SiO
2
variation diagrams (Fig. 16a) shows that Zr in
the R1-rhyolites decrease toward more felsic compositions, similar to
the trends in A-types felsic magmas derived from a quartz-feldspathic
source withZr-saturation (King et al., 1997). In addition, it is empirically
established that A-type felsic rocks have 2 to 3 times higher Zn contents,
than compositionally similar S-, I-, and M-types granitoids and related
volcanic rocks (b60 ppm), because of the high solubility of Zn at high
temperature in such magmatic systems (Fig. 16b; Lentz, 1998). The
R2- and R3-rhyolite have low Zr concentrations (53 to 78 ppm, except
for one sample); and contain several inherited zircon crystals with U-
Pb isotopic ages of ~405 and 422 Ma. Thus, (i) these MS-rhyolites are
probably related to hot-dry-type felsic rocks without or with poor in-
heritance (Miller et al., 2003); and (ii) this confirms once again that
the source for MS-rhyolites could have been the Early Devonian
supra-crustal (metagreywacke) material from the upper level of the
Rudny-Altai basement.
Fig. 8. Major element classification of the MS-rhyolites. (a) Total alkalis vs. Silica diagram (Irvine and Baragar, 1971;Le Maitre et al., 1989), division according to (Middlemost, 1994).
(b) K
2
O vs. SiO
2
alkalinity diagram (Rickwood, 1989). (c) FeO
T
/(FeO
T
+ MgO) vs. SiO
2
(wt%) diagram (Frost et al., 2001). (d) Aluminum Saturation Index (Al/(Ca-1.67P + Na + K vs.
Al/(Na + K) molar ratio)) diagram (Shand, 1943). The diagrams show the extreme petrochemical characteristics of the MS-rhyolites.
328 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

6.3. Trace elements constraints
Firstly,as noted by our reviewer “There is little one can do, geochem-
ically speaking, when dealing with weathered old volcanic rocks, except
to use immobile trace elements”. It has been experimentally shown
(Conrad et al., 1988;Patiño Douce and Beard, 1996), that higher-silica
melts can be produced by vapor-absent melting of biotite-
hornblende-bearing assemblages only at low aH
2
O and high tempera-
tures (N800 °C), due to anhydrous-minerals as the dominant residual
phases during melting. Alternatively, felsic orthogneisses of tonalitic to
granodioritic composition, consisting of quartz-feldspathic mineralogy,
as well as amphibole and biotite, and titanite, apatite, and zircon, have
been recognized as one of the potential source for A-type granites and
related volcanic rocks (e.g. Collins et al., 1982;Anderson, 1983;White
and Chappell, 1983;Rutter and Wyllie, 1988;Creaser et al., 1991;
Skjerlie and Johnston, 1993;Patiño Douce, 1997). In term of incompat-
ible elements, geochemical characteristics of the MS-rhyolites ((La/Yb)
n
~ 4; Zr/Y ~ 4; La/Sm ~ 3–5), transitional between those of upper-
crust material and mantle-derived materials (Fig. 10b–d; Sun and
McDonough, 1989;Rudnick and Gao, 2003), support biotite-
hornblende-bearing source. Based on Zr and Y partition coefficient be-
tween the mineral and melt (D
Zr
Amph/Melt
b1; D
Y
Amph/Melt
N1; e.g. Pearce
and Norry, 1979;Arth, 1976),in the absence of amphibole residue, Y be-
comes incompatible, as is Zr, resulting in low overall Zr/Y with increas-
ing Y. Amphibole will have the similar effect on the La/Yb ratio because
LREEs are more incompatible than HREEs in this mineral; and, thus, the
values of their relationsshould similarly approach the valueclose to that
of the tholeiite series (Fig. 16c; Lentz, 1998). Speaking of LREEs, they
have a minor dependence on fractionation of phenocryst phases such
as quartz and feldspars in granite systems, but are strictly controlled
by the melting of biotite (D
LREE
Bt/Melt
N1; D
La
Bt/Melt
≫D
Sm
Bt/Melt
), and have a de-
pendence on the stability of some accessory phases, such as alanite and
monazite (Mahood and Hildreth, 1983). In general, the Yb vs. La/Yb di-
agram (Fig. 16c) indicates that partial melting of amphibole-bearing
source (Beard and Lofgren, 1991) rather than hypothetical fractional
differentiation of minerals in viscous high-silica magma played a domi-
nant role in the evolution of the R1-rhyolites.
Secondary, if we consider the mineral vectors in the system of La vs.
La/Yb, reflecting the relationship between geochemical indices and frac-
tionation, the involvement of minor amounts of accessory phases in the
petrogenesis of these rhyolites is also an unlikely scenario (Fig. 16d).
The diagram La vs. La/Sm (Fig. 16e) shows that R1-rhyolites have a
cogenetic distribution with vector reflecting the partial melting of a
biotite-containing source (Koester et al., 2002). Although it is clear
that the compositions of R3-rhyolites again behave in some other way.
In general, we believe that some of the processes associated with
high-temperature melting of metagreywacke materials or felsic meta-
igneous rocks in the shallow-crustal level played a major role in the or-
igin of MS-rhyolites.
6.4. Mechanism of the MS-rhyolite formation
Firstly, the geotectonic environments in favour of the widespread
high-silica bimodal-type rhyolitic magmatism are usually associated
with episodes of extensional geodynamic activity of rifted mature island
arcs, rifted continental margin and marginal back-arc basin, or within-
plate extensional settings, in contrast to the compositional continuum
in normal subduction environments (Gribble et al., 1996;Lentz, 1998;
Fig. 9. Major andtrace elements alteration plots for the MS-rhyolites. (a): LOI vs. Na
2
O and K
2
O (wt%) diagrams, indicates, that sodium alterations could be related to hydrothermal fluid,
and, in contrast, potash alterations were not associated with it. (b) Igneous Spectrum diagram (Hughes, 1972), indicates, that R1-rhyolites lie inside the field weakly altered magmatic
rocks; and R2- and R3-rhyolite marked inside the field K-metasomatic rocks. (c) Alteration-Box-Plot (Large et al., 2001), showing MS-rhyolites alteration trends (CCPI = Chlorite-
Carbonate-Pyrite Index = 100 ∗(MgO + FeO
T
)/(Mg + FeO
T
+K
2
O+Na
2
O); and Al = Hashimoto alteration index = 100 ∗(K
2
O + MgO)/(K
2
O + MgO + Na
2
O + CaO)). The R1-
rhyolites lie inside least altered box, reflecting their weakly alteration; and the R2- and R3-rhyolite lie on arrays from the least altered box toward the Ser-Kfsp side, associated with
hydrothermal alteration. (d) Zr vs. Na
2
O/Al
2
O
3
diagram, indicating a lack of clear correlation between Zr content and sodium hydrothermal fluid for weakly altered MS-rhyolites.
(e) La/Sm and Gd/Yb vs. Alteration Index diagram (Lentz, 1998), reflecting the minor modification of LREEs during chloritization (Fe + Mg) or feldspathization (Na + K) processes.
(f) REEs vs. Zr (ppm) diagram indicate that REEs were relatively immobile during hydrothermal alteration, since ithas the strong correlation with Zr concentrations.
329M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Barrie and Hannington, 1999;Orozco-Esquivel et al., 2002;Bachmann
and Bergantz, 2004, 2008;Piercey et al., 2006;Galley et al., 2007;
Gaboury and Pearson, 2008;Tamura et al., 2009;Deering et al., 2010).
A huge contribution to understanding the processes of petrogenesis
and evolution of felsic magmatism was made by Bachmann and
Bergantz (2008),andDeering et al. (2008, 2010). They showed that it
is mainly generated by extraction of viscous melts from solid residues
either in (i) long-lived crystallizing mush zones fed by mafic to interme-
diate magmas (dominantly down-temperature evolution with residual
silicic melt), or in (ii) partial melting zones within the crust (dominantly
up-temperature evolution with newly formed silicic melt). In the first
case, dominated by flux melting of the supra-subduction mantle lead
to generation of the cold-oxidized felsic magmas. As noted (Orozco-
Esquivel et al., 2002), high-silica rhyolite or crystal-poor rhyolite is ex-
tracted by squeezingof the partly crystallized mush,percolates through
pores and fractures, concentrates in small pockets and ascends to its
final emplacement level using the existing plumbing system. However,
in the Rudny Altai there are no low-silica granitoids, syngenetic to R1-
rhyolite, which could be considered as enriched in crystals cumulated
material originating from crystalline mushes (Kuibida et al., 2015). In
second case, depending on the decompression melting of the mantle
(higher Tand/or lower fO
2
/aH
2
O) lead to origin of the hot-reduced felsic
magmatism. Extensional regimes where major crustal fracturing may
occur, controls injection and emplacement of the basic magmas into
the different levels of continental crust, producing volumetric melting
of the crustal material. In these conditions the high temperatures and
low to moderate water contents felsic magmas are relatively fluid,
Fig. 10. Trace element classificationof the MS-rhyolites. (a)Zr/TiO
2
-Nb/Y (Winchester andFloyd, 1977); (b) La
n
vs. Yb
n
; (c) Y vs. Zr (ppm); and(d) La
ucn
vs. Sm
ucn
diagrams(Lentz, 1998;
Piercey et al.,2008), illustrating the transitional geochemical characteristicsof MS-rhyolites between mantle-derived and upper crust material. UCN = Upper crust normalized (Rudnick
and Gao, 2003); values for Depleted mantle and Chondrite are from Sun and McDonough (1989).
Fig. 11. Tectonic discriminant diagram La-Y-Nb (Cabanis and Lecolle, 1989), illustrating
the transitional geochemical characteristics of MS-rhyolites between normal arc-related
and within-plate series. 1A –calc-alcalaine series; 1B –calc-alkaline and tholeiite series;
1С–arc-related tholeiitic series; 2A –within-plate series.
330 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

highly mobile melts, able to reach very shallow depths in the crust
(Clemens et al., 1986). Accordingly, the rapid extraction of magma
from its source and ascent to the surface and probably inhibited its stag-
nation in magma chambers, contributes to (i) non-equilibrium melt
generation conditions; and (ii) limited degree of assimilation and frac-
tional crystallization, which are important factors affecting the compo-
sitional evolution of magma. Although there may be transitional
varieties between those generated by mantle upwelling and those
formed in arc-related environments (Bachmann and Bergantz, 2008;
Deering et al., 2008).
Secondary, the trace-element concentrations and their elemental ra-
tios are a reflection of melting reactions, related to intensive parameters
(P,T,fO
2
,aH
2
O) and controlled by the equilibrium between melt and re-
sidual assemblage in different source region and geotectonic environ-
ments. The distinctly different compositions (REEs, HFSEs) between
the two geochemical types of the MS-rhyolites may reflect the differ-
ence in their petrogenetic histories. At first sight, the REE-patterns of
the R1-rhyolites has some similarities to those of wet-cold-oxidized
felsic magmas (Fig. 15c), suggesting that the melting of the source re-
gion occurred under conditions resemble to those of the subduction
volatile-rich system (Bachmann and Bergantz, 2008;Deering et al.,
2010). However, R1-rhyolites were produced under relatively more re-
ducing conditions, resulting in: (i) an increase in MREEs and Y, Zr, and
Zr/Y ratios; and (ii) low crystallinity and dominance of anhydrous
phases. For example, Deering et al. (2008) showed distinctive geochem-
ical features between normal arc-related wet-cold-oxidized felsic
magmas and reduced arc-related extensional felsic magmas, by the ex-
ample of Taupo Volcanic Zone, New Zealand rhyolites (Fig. 15c). The
seagull-like REE-patterns of the R2- and R3-rhyolites resemble those
of rhyolites formed in dry reduced crustal conditions associated with in-
crease in temperature induced by thermal input from the underlying
hot mantle (Wilson et al., 2005); Fig. 15d. Felsic magmas with such geo-
chemical characteristics are commonly attributed to rift-related areas of
mantle upwelling, such as hotspots, continental rifts, and mid-ocean
ridges (Bachmann and Bergantz, 2008;Deering et al., 2010), although
noted also for the back-arc basin, as summarized in (Barboni and
Fig. 12. Tectonic discriminant diagrams. (a–f) Y, Nb, Zr, Zn, Ce and Zr + Nb + Ce + Y vs. 10
4
∗Ga/Al diagrams (Whalen et al., 1987;Sylvester, 1989;Wu et al., 2017), showing A-type
evolutionary trend of the MS-rhyolites. (g–j) Nb vs. Y and Rb vs. Y + Nb diagrams (Pearceet al., 1984;Whalen and Hildebrand, 2019), illustrating the transitional geochemical character-
istics of MS-rhyolites between normal arc-related and within-plate series, and the fact that they do not belong to compositional field of the slab-failure-related series. SYNCOLG = syn-
collisional granites, WPG = within-plate granites, ORG = oceanridge granites, VAG = volcanic arc granites.
331M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Bussy, 2013). Interestingly, although R2- and R3-rhyolites have contrast
specificity of alkalis, but they have close similarityof REE-patterns, com-
mon low La/Sm ratios and Zr concentrations, which suggests their ge-
netic relationship with each other. In the simplifying assumptions, the
following two-stage petrological model is proposed: initial non-
equilibrium melting at deeper hot levels, and subsequent melt segrega-
tion and degassing at subsurface levels. Previous studies indicate that
Zr- and LREEs-undersaturated dry crustal high-K rhyolites have similar
REE-patterns, and can be produced as a result of high temperature
water-undersaturated disequilibrium melting, in case of melt extraction
from its residuum occurred faster than extensive dissolution of acces-
sory phases (Fig. 15d; Watt and Harley, 1993;Orozco-Esquivel et al.,
2002). In the shallow-crustal levels, further segregation of parent
magma into two contrasting potash-rich and sodium-rich types
magmas could occur, as a result of the abrupt “dehumidification”of
volatiles-rich magmatic chamber during its decompression and
degassing (Abramov, 2004;Samuel et al., 2007). It is suggested that
the active unloading of the magmatic system in many cases leads to
the loss of volatile components even before the final crystallization of
magma, and results in its segregation into two melts, the first of
which, more potassium, accumulates in the lower part of the magmatic
chamber, and the second, more sodium, saturated with volatile compo-
nents and mobile, is injected intothe host rocks (Inshin, 1972;Avdonin,
1975;Samuel et al., 2007). For example, some extreme high-Na rhyo-
lites were interpreted as residual melt extracted from partially crystal-
lized granite-like mush (ca. 50% of K-feldspar and quartz, and minor
biotite), at a late stage of crystallization (Fig. 15d; Barboni and Bussy,
2013). This scenario is supported by the aplitic petrographic texture of
the Na-rich R2-rhyiolites, which could represent a melt removed from
the magmatic chamber during its depressurization; and, in contrast,
the porphyritic type potash-rich R3-rhyolite are in association with
magmatic breccia that could represent early magmatic accumulation.
Thirdly, we believe that in the course of the longevolution of the vol-
canic system felsic magmas with island-arc-like geochemical features
may change torift-like one in response to the changes in the crustal ar-
chitecture and magnitude of extension in the lithosphere (e.g. Gribble
et al., 1996;Deering et al., 2010). In other words, we believe that arc-
like extensional R1-rhyolite formed immediately after the beginning
of rifting, and widespread upper crustal melting occurred under rela-
tively more reduced conditionsthan the formation ofwet-cold-oxidized
magma in normal volatile-rich supra-subduction systems (Fig. 17a).
The generation of rift-like R2- and R3-rhyolites occurred under dry
and/or non-equilibrium conditions, synchronously with the rise in the
upper crust of the Givetian-Fransian basic magmas, as a result of the
prolonged strike-slip deformation, the Alei pull-apart basin opening, ac-
tive lithospheric extension and high thermal input from the underlying
hot mantle (Fig. 17b). We can draw an important preliminary conclu-
sion that the initial rifting in the Rudny Altai could have started in the
Late Emsian –Early Eifelian, since R1-rhyolites bear some arc-like geo-
chemical signatures, but formed under more reduced conditions. In ad-
dition, although in the evolution of the entire volcanic system there
have been changes in the mechanisms of magma generation, it has
Fig. 13. Chondrite-normalized REEs diagrams (Sun andMcDonough, 1989)for the MS-rhyolites, in comparison with rhyolites from (a)rhyolite-dominate intra-oceanic arc(Tamura et al.,
2009), and marginal island-arc (Shinjo and Kato, 2000); (b) intra-oceanic back-arc basin (Tamura et al., 2009); rifted mature island-arc (Deering et al., 2011), and continental-margin
extensional setting (Morata et al., 2000); (c) initial ridge-subduction setting (Johnson and O'Neil, 1984;Cole and Basu, 1992;Shen et al., 2011); (d) initial post-collisional setting
(Chen et al., 2011), and marginal back-arc basin (Shinjo and Kato, 2000;Yamada et al., 2012). These diagrams illustrate that the REE-patterns of the MS-rhyolites are most related to
those of the bimodal associations of the continental-margin extensional setting, rifted mature island-arc and marginal back-arc basin settings; and from some post-collision settings.
332 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

not affected the spatial distribution of the emplaced magmas, which in-
dicates the long-living nature of the permeable tectonic zones of the
Alei pull-apart structure and emphasizes its rift-setting.
6.5. Tectonic implication
Firstly, the Altai orogen was under a convergent environment in the
Devonian, but debate exists on whether it was an active continental
margin of Andean-type, that evolved in the sialic island arc of Japan-
type (Rotarash et al., 1982;Shokalsky et al., 2000;Glorie et al., 2011;
Chen et al., 2019), or a continental margin rift setting (Kozlov, 1995;
Bespaev et al., 1997;Promyslova, 2004). Some geologists consider that
the frontal part of the Altai continental margin resembles Okinawa
Through behind the Ryukyu SW of Japan island arc or Taupo Rift of
the North Island of New Zealand in Tonga-Kermadek arc system
(Karaulov et al., 1992). A ridge subduction was proposed for the contin-
uation of the Rudny-Altai volcanic belt in the SE (Ashele basin, Xinjiang,
China; e.g. Wan et al. , 2010;Wu et al., 2015;Yang et al., 2018), in such as
scenario hot oceanic plate was subducted and subsequent slab window
was formed (e.g. Ma et al., 2018; and reference above). It should be
noted that the question of the collapse of the overthickened orogen
was never considered in any of the geodynamic models, which would
inevitably have to happen if the Andean-type subduction orogen
evolved. In the light of the newest world geodynamic models, it is
widely known that the collapse of the Andean orogens on both sides
of the Pacific was accompanied by lithospheric delamination and as-
thenospheric upwelling (e.g. Wang et al., 2007;Xu et al., 2006;Zhu
et al., 2010), and generation of adakite-like magmas as a result of partial
melting of overthickened continental crust (Foster et al., 2001;Kapp
et al., 2002;DeCelles, 2004). However, Devonian adakites (Scherba
et al., 1998;Murzin et al., 2001), have not yet been found in either the
Russian or the Kazakhstan parts of the Altai orogen, as the evidence of
the its collapse; or subduction spreading zone and slab melting in sub-
duction windows (e.g. Defant and Drummond, 1993;Drummond
et al., 1996;Sajona et al., 2000). It is possible that they do exist, but
have not yet been discovered due to the lack of a detailed regional ana-
lytical database for these territories. It is also possible that at the post-
subduction stage the frontal part of Rudny Altai was fragmented and
moved by large-amplitude sinistral deformation (Buslov, 2011), as evi-
denced, for example, by an anomalously distance of ~20 km between
the Rudny Altai volcanic belt and Irtysh Shear Zone.
Secondary, the formation of the volcanic belt of the NW Rudny Altai
could be related to a continental-margin region extension, induced by a
distant spreading of oceanic ridge, consistent withits sub-parallelorien-
tation to the continental margin (e.g. Ma et al., 2018). For example, Ter-
tiary rift-like felsic magmatism in the Mesa Central, Mexico, occurred in
the rear part of Baja California (Orozco-Esquivel et al., 2002). On the
other hand, the ridge-subduction-related processes should have oc-
curred later than the formation of a normal supra-subduction volcanic
belt; otherwise it is difficult to imagine the subduction of the spreading
zone without subduction the oceanic plate. However, in the Rudny Altai
there is no normal arc-related volcanism occurred before the formation
of the bimodal-type MS-rhyolites (~390–380 Ma), and there is a genetic
link of volcanism and volcanogenic massive sulphide deposits with
rifting. Thus, these do not contribute a geodynamic setting of normal is-
land arc.
Thirdly, it could be a reflection of an oblique subduction and slab
roll-back processes of “extension accretionary orogen”(Collins, 2002),
that induced regional extension of the continental margin.For example,
the Early Devonian initialization of the Altai active margin occurred
Fig. 14. εNd(t) vs. T
DM
(Ma) age plot for the MS-rhyolites. Data for the Gorny Altai arc-
related Cambrian volcanic rocks and Ordovician metaturbidites are according to (Kruk
et al., 2010;Kruk, 2015); those fo r the Early Devonian volcanic rocks from the Alta i-
Mongolian block (from basalt to rhyolite) are ours. For details of data selec tion see
Table 3. Inset with the diag ram εNd(t) vs. Na
2
O
3
/Al
2
O
3
illustrates the composition of
rocks altered and unaltered by potassium metasomatism.
Fig. 15. (a, b) Chondrite-normalized REEs diagrams (Sun and McDonough, 1989), for the
MS-rhyolites, in relative to cold-wet-oxidized an d hot-dry-reduced rhyolites. These
diagrams show that the REE-patterns of the R1-rhyolites have affinity with some of the
transitional hot-dry-reduced rhyolites of subduction-related extensional settings. The
REE-patterns of the R2- and R3-rhyolites resemble those of some Na-rich residual
rhyolites and K-rich disequilibrium rhyolites, also marked for subduction-related
extensional settings. References to data sources are shown in the figure.
333M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

within the terrain-orogenic belt (Gorny Altai) of the Siberian continent,
~20 Ma before the formation of the volcanic belt on the front of the ac-
tive margin (Rudny Altai), as it follows from numerous geochronologi-
cal data (Shokalsky et al., 2000;Babin et al., 2004;Glorie et al., 2011;Cai
et al., 2014;Krupchatnikov et al., 2018;Chen et al., 2019); Fig. 1. This
geological record resembles the Japan magmatic arc formed at a terres-
trial continental margin via a stage of spreading in a back-arc basin
followed by multiple stages of submarine rifting (e.g. Yamada et al.,
2012). It is assumed the Rudny-Altai block is a back-arc basin fragment
formed after the destruction of the Siberian passive margin in the Early
Devonian. Based on all the lines of evidence presented in this study, we
propose extension regime in the transition area between the island-arc
and back-arc basin for the origin of the NW Rudny Altai volcanic belt.
The retreat and change in the angle of the subducting plate (from flat
to steep) may be one of the causes for the migration of volcanism
from the continent to the ocean at the Altai convergent margin.
7. Conclusion
Our new field and geochemical data for the Devonian magmatism in
the NW Rudny Altai volcanic belt (Russia) suggest:
(1) Two different geochemical types of subvolcanic silica-rich intru-
sions (MS-rhyolites) were emplaced at ~390 and 380 Ma, respec-
tively. They were triggered by intensive melting of the upper
crust materials beneath the Rudny Altai. The consistent forma-
tion of the arc-like rhyolites (R1) and then rift-like rhyolites
(R2 and R3) suggests progressive rifting.
(2) The island-arc-like R1-rhyolite formed immediately after the be-
ginning of rifting due to widespread crustal melting under re-
duced conditions. The rift-like R2- and R3-rhyolites were
generated under non-equilibrium conditions, synchronously
with the ascent of basaltic melts to the upper crust as a result
of lithospheric extension and high thermal input from theunder-
lying hot mantle.
(3) The RA-rhyolites formation was related to riftingin the transition
area between the island-arc and back-arc basin, ~20 Ma after the
initial continental-margin region extension.
Declaration of competing interest
The authors declare that they have no known competing financial
interests or personal relationships that could have appeared to influ-
ence the work reported in this paper.
Fig. 16. (a)Zr (ppm) vs. SiO
2
(wt%) content for A-andI-type granites at Zr-saturatedcrystallization(King et al., 1997). The chemical trends showcontinuously decreasing Zr with increas-
ing SiO
2
and indicating derivation from a quartz-feldspathic source and primary zircon saturation. (b)Zn vs. SiO
2
diagram forA-andM-, I-andS-type granites, illustrating the relationship
betweenhigh Zn contents and high-TA-type felsic magmas(Lent z, 1998).(c) La/Yb vs. Yb plot showingrelationship between the changeof REEs of MS-rhyolites and dehydration melting
amphibole-bearingsource, with a non-garnetresidual phase (Beardand Lofgren, 1991). (c)La/Yb vs. La plot showing, that variations ofthese elements are notcontrolled by fractionation
of accessory phases (Li et al., 2013); (d) La/Sm vs. Sm plotshowing the relationship between the change of REEs of MS-rhyolites and the dehydration melting of a biotite-bearingsource,
with a non-garnet residual phase (Koester et al., 2002).
334 M.L. Kuibida et al. / Gondwana Research 81 (2020) 312–338

Acknowledgments
Writing this paper would be impossible without years of the hard
field and laboratory work of skillful geologists from Russia, China and
Kazakhstan. We are sincerely grateful to the reviewers who have en-
abled us to make this paper better. The study was carried out as part
of a basic research program at the Ministry of Science and Education
of the Russian Federation (Projects Nos. 14.Y26.31.0018), the Russian
Foundation for Basic Research (project no. 16-05-01021), Hong Kong
RGC grants 17302317 and 17303415; and according to the plan of the
State Assignment of the Institute of Geology and Mineralogy, Siberian
Branch, Russian Academy of Sciences.
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