The Dovyren Intrusive Complex (Southern Siberia, Russia): Insights into dynamics of an open magma chamber with implications for parental magma origin, composition, and Cu-Ni-PGE fertility

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DOI: 10.1016/j.lithos.2018.01.001
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The Dovyren Intrusive Complex (DIC, Northern Baikal region, 728 Ma) includes the layered dunite-troctolite-gabbronorite Yoko-Dovyren massif (YDM), associated mafic-ultramafic sills, and dykes of olivine-rich to olivine-free gabbronorite. Major rock types of the DIC are presented, including a diversity of olivine orthocumulates to olivine-plagioclase and gabbroic adcumulates, carbonate-contaminated ultramafics and Cu-Ni-PGE mineralisation. Detailed comparisons of complete cross-sections of the YDM in its centre and at the NE and SW margins demonstrate differences in the cumulate succession, mineral chemistry, and geochemical structure that likely reflect variations in parental magma compositions. Combining petrochemical reconstructions for most primitive rocks and calculations using the COMAGMAT-5 model, it is shown that the central and peripheral parts of the intrusion formed by olivine-laden parental magmas ranged in their temperatures by 100 °C, approximately from 1290 °C (~11 wt% MgO, olivine Fo88) to 1190 °C (~8 wt% MgO, olivine Fo86). Thermodynamic modelling suggests that the most primitive high-Mg magma was S-undersaturated, whereas its derivatives became S-saturated at T < 1240–1200 °C. These estimates are consistent with geological observations that mostly sulphide-poor mineralisation occurs in the centre of the intrusion, whereas Cu-Ni sulphide ores (locally net-textured) occur in its NE and SW parts, as well as in the underlying peridotite sills. The primitive S-undersaturated olivine cumulates became sulphide-saturated at a post-cumulus stage. As a result, Ni-rich immiscible sulphides formed within and migrated through the early olivine-rich cumulate piles to generate poorly-mineralised plagiodunite. In the troctolite and gabbroic parts of the Dovyren chamber, sulphide immiscibility likely occurred at lower temperatures, producing Cu-rich sulphide precursors, which gave rise to the ‘platinum group mineral’ (PGM-containing) troctolite and low-mineralised PGE-rich anorthosite in the Main Reef. The geochemical structure of the YDM demonstrates C-shaped distributions of TiO2, K2O, P2O5, and incompatible trace elements, which are 3–5 fold depleted in the cumulate rocks from the inner horizons of the intrusion with respect to the relatively thin lower and upper contact zones. In addition, a marked misbalance between estimates of the average composition of the YDM and that of the proposed olivine-laden parental magmas is established. This misbalance reflects a significant deficit of the YDM in incompatible elements, which argues that 60–70% of basaltic melts had to have been expelled from the Dovyren magma chamber during its consolidation. A possible scenario of the evolution of the open magma chamber is proposed.
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Review article
The Dovyren Intrusive Complex (Southern Siberia, Russia): Insights into
dynamics of an open magma chamber with implications for parental
magma origin, composition, and Cu-Ni-PGE fertility
Alexey Ariskin
, Evgeny Kislov
Andrew McNeill
, Yuri Kostitsyn
, Sandrin T. Feig
, Alexey Malyshev
Faculty of Geology, Moscow State University, Leninskie Gory 1, 119234 Moscow, Russia
Vernadsky Institute, Kosygin Str. 19, 119991 Moscow, Russia
CODES CoE and Earth Sciences, University of Tasmania, Private Bag 79, Hobart, TAS 7001, Australia
Geological Institute, Sakhyanovoy Str., 6a, 670047 Ulan-Ude, Russia
Centre for Exploration Targeting, School of Earth and Environment, ARC Centre of Excellence for Core to Crust Fluid Systems, The University of Western Australia, 35 Stirling Highway, 6009,
Crawley, Perth, Western Australia, Australia
Mineral Resources Tasmania, PO Box 56, Rosny Park, Hobart, TAS 7018, Australia
Central Science Laboratory, University of Tasmania, Private Bag 74, Hobart, TAS 7001, Australia
abstractarticle info
Article history:
Received 21 April 2017
Accepted 7 January 2018
Available online 13 January 2018
The Dovyren Intrusive Complex (DIC, Northern Baikal region, 728 Ma) includes the layered dunite-troctolite-
gabbronorite Yoko-Dovyren massif (YDM), associated mac-ultramac sills, and dykes of olivine-rich to
olivine-free gabbronorite. Major rock types of the DIC are presented, including a diversity of olivine
orthocumulates to olivine-plagioclase and gabbroic adcumulates, carbonate-contaminated ultramacs and
Cu-Ni-PGE mineralisation. Detailed comparisons of complete cross-sections of the YDM in its centre and at the
NE and SW margins demonstrate differences in the cumulate succession, mineral chemistry, and geochemical
structure that likely reect variations in parental magma compositions. Combining petrochemical reconstruc-
tions for most primitive rocks and calculations using the COMAGMAT-5 model, it is shown that the central
and peripheral parts of the intrusion formed by olivine-laden parental magmas ranged in their temperatures
by 100 °C, approximately from 1290 °C (~11 wt% MgO, olivine Fo88) to 1190 °C (~8 wt% MgO, olivine Fo86).
Thermodynamic modelling suggests that the most primitive high-Mg magma was S-undersaturated, whereas
its derivatives became S-saturated at T b12401200 °C. These estimates are consistent with geological observa-
tions that mostly sulphide-poor mineralisation occurs in the centre of the intrusion, whereas Cu-Ni sulphide
ores (locally net-textured) occur in itsNE and SW parts, as well as in the underlying peridotite sills. Theprimitive
S-undersaturated olivine cumulates became sulphide-saturated at a post-cumulus stage. As a result, Ni-rich im-
miscible sulphides formed within and migrated through the early olivine-rich cumulate piles to generate poorly-
mineralised plagiodunite. In the troctolite and gabbroic parts of the Dovyren chamber, sulphide immiscibility
likely occurred at lower temperatures, producing Cu-rich sulphide precursors, which gave rise to the platinum
group mineral(PGM-containing)troctolite and low-mineralised PGE-rich anorthosite in the MainReef. The geo-
chemical structure of the YDM demonstrates C-shaped distributions of TiO
O, P
, and incompatible trace
elements,which are 35 fold depleted in the cumulate rocksfrom the inner horizonsof the intrusion with respect
to the relatively thin lower and upper contact zones. In addition, a marked misbalance between estimates of
the average composition of the YDM and that of the proposed olivine-laden parental magmas is established.
This misbalance reects a signicant decit of the YDM in incompatible elements, which argues that 6070%
of basaltic melts had to have been expelled from the Dovyren magma chamber during its consolidation.A possi-
ble scenario of the evolution of the open magma chamber is proposed.
© 2018 Elsevier B.V. All rights reserved.
Layered intrusion
Cu-Ni-PGE mineralisation
Parental magma
Sulphide immiscibility
Open magma chamber
Anomalous mantle source
Lithos 302303 (2018) 242262
Corresponding author at: Faculty of Geology, Moscow State University, Leninskie Gory 1, 119234 Moscow, Russia.
E-mail addresses:, (A. Ariskin).
0024-4937/© 2018 Elsevier B.V. All rights reserved.
Contents lists available at ScienceDirect
journal homepage:
1. Introduction.............................................................. 243
2. Summaryofthegeology,structure,andfertilityoftheSynnyr-DovyrenComplex............................... 245
2.1. Regionalgeology ........................................................ 245
2.2. GeochronologyoftheDICandassociatedvolcanics......................................... 245
2.2.1. AgeoftheDovyrenIntrusiveComplex........................................... 245
2.2.2. Ageofvolcanicsandtimingofoverprintedprocesses.................................... 246
2.3. Sr-Ndisotopiccompositions ................................................... 246
2.4. MaintypesofYDMcumulaterocks ................................................ 246
2.5. Magnesian skarns in ultramacs ................................................. 247
2.6. Cu-Ni-PGEmineralisationwithintheDIC.............................................. 248
2.6.1. SyngeneticCu-Nimineralisation ............................................. 248
2.6.2. EpigeneticCu-Nimineralisation.............................................. 248
2.6.3. PGE-richsulphide-poormineralisationinanorthosite.................................... 248
2.6.4. RecentlydiscoveredPGM-containingtroctolite....................................... 249
3. DetailedstructureoftheYoko-Dovyrenmassif .............................................. 249
3.1. TheBolshoi-Tsentralnyicross-section............................................... 249
3.1.1. Chilledrocksatthelowercontact............................................. 249
3.1.2. Basalplagioperidotite .................................................. 249
3.1.3. Plagiodunite ...................................................... 250
3.1.4. Dunitezone....................................................... 250
3.1.5. Highlycontaminateddunite ............................................... 251
3.1.6. AdcumulateandCpx-bearingtroctolite .......................................... 251
3.1.7. Zoneofolivinegabbro.................................................. 251
3.1.8. Olivinenoritetogabbronoriteandnear-roofrocks..................................... 252
3.2. StructureoftheYDMattheNEandSWmargins .......................................... 252
3.2.1. TheSchkolnyisection .................................................. 252
3.2.2. TheYokosection .................................................... 252
4. Underlyingsillsandassociatedvolcanics................................................. 252
4.1. Mac-ultramacsills....................................................... 252
4.2. Volcanicsequences ....................................................... 252
4.2.1. High-Tibasaltsandquartz-feldsparporphyry ....................................... 253
4.2.2. Low-Tivolcanics .................................................... 254
5. Traceelementgeochemistry....................................................... 254
.................................................................... 254
5.1.1. ThecentreoftheYDM.................................................. 2555.1.2. TheSchkolnyisection .................................................. 255
5.1.3. Low-Tiandhigh-Tivolcanics............................................... 255
6. Discussion............................................................... 255
6.1. ParentalmagmasoftheDovyrenIntrusiveComplex......................................... 255
6.2. Sulphideimmiscibilityintheproposedmagmas .......................................... 256
6.3. Fingerprintsofanopenmagmachamber.............................................. 257
6.3.1. Misbalanceofincompatiblecomponents.......................................... 257
6.3.2. Signicance of the Al O
2 3
-MgOdiagram........................................ 257
6.4. FormationandevolutionoftheDovyrenmagmachamber...................................... 258
6.4.1. The rststage...................................................... 259
6.4.2. Thesecondstage .................................................... 259
6.4.3. Thethirdstage ......................... ............................ 260
6.4.4. Thefourthstage..................................................... 260
6.5. Themantlesource........................................................ 260
7. Conclusions.............................................................. 26
Acknowledgements............................................................. 261
References................................................................. 261
1. Introduction
Sedimentary sequences within the foldbelts surrounding the
southern margin of the Siberian Craton host a number of Precambrian
large layered intrusions and mac-ultramac complexes referred to as
the Cu-Ni-PGE East Siberian metallogenic province (ESMP; Polyakov
et al., 2013). The ~728 Ma Dovyren Intrusive Complex (hereafter DIC)
is located ~60 km NE of Lake Baikal and represents the eastern branch
of the ESMP (Fig. 1). The main components of the DIC include the
Yoko-Dovyren mac-ultramac massif (YDM), underlying ultramac
sills, and associated dykes or sill-like bodies of gabbronorite, both
below and above the YDM (Grudinin, 1963, 1965;Gurulev, 1965,
1983;Kislov, 1998;Konnikov, 1986;Yaroshevskii et al., 1982). The
main intrusive body (also referred to in the Russian literature as
the Dovyren intrusionor simply Dovyren) is composed of a su ccession
of ultramacandmac rocks ranging from plagioperidotite and
dunite to troctolite, olivine gabbro, and gabbronorite. It forms a 26 km
long lens-shaped ridge, which is up to ~3.5 km wide in its central part
(Fig. 2).
The DIC is one of the best-exposed layered complexes in the world
(see below) and, over the last 50 years, it has continued to be a refer-
ence site for Russian petrologists dealing with the intra-chamber
differentiation of ultramac magmas and the Cu-Ni-PGE fertility of
layered intrusions. This is because of the relatively simple geological
setting and uncomplicated internal structure of the YDM (Fig. 2),
excellent exposure of complete cross-sections from the bottom to
the roof, the occurrence of most known types of Cu-Ni sulphide and
PGE-rich mineralisation, as well as Cu-Ni sulphide ores near the bottom
5.1.1. ThecentreoftheYDM.................................................. 254
5.1.2. TheSchkolnyisection.................................................. 255
243A. Ariskin et al. / Lithos 302303 (2018) 242262
(e.g., Ariskin et al., 2009, 2016;Balykin et al., 1986;Bolikhovskaya et al.,
2007;Kislov, 1998;Konnikov, 1986;Yaroshevskii et al., 1982, 2006).
The DIC is spatially and temporally associated with Inyaptuk-Synnyr
volcanics (Manuilova and Zarubin, 1981), which include high-Ti basalts
and low-Tibasalts to basaltic andesite. The latter are geochemically and
isotopically similar to the DIC (Ariskin et al., 2013b, 2015a), suggesting
that the low-Ti volcanics are genetically related to the intrusive rocks
(Kislov, 1998). The association between the DIC rocks and the low-Ti
Fig. 1. Major tectonic features of the Siberian Craton and its southern margin. The outline of the Siberian craton includes the Neoproterozoic and older basement (after Parfenov et al.,
2010). BP outerBaikal-Patom Foldbelt, EA East Angara Fold bel t. Tectonic Collages: CS Circum-Siberia (Proterozoic), YT Yenisey- Transbaikal (Vendian through Early Ordovician),
AL Altay (Vendian through Ordovician), MO Mongol-Okhotsk (Devonian through Late Jurassic). Late Proterozoic and Cambrian superterranes: AR Argun-Idermeg, TM Tuva-Mongolia.
The Late Proterozoic intrusions of the southern margin of the Siberian Craton (the East Siberian metallogenic province; Polyakov et al., 2013) are shown by red circles and include:
KKingash, VK Verkhnii (Upper) Kingash, YDM Yoko-Dovyren massif, and other small mactoultramac bodies hosting Cu-Ni-PGE deposits and sulphide ore occurrences.
Fig. 2. Photograph of the YDMfacing northwest (top) and a schematic geological map of the southwestern termination of the Olokit trough, modied after Ariskin et al. (2013b,2015a).
244 A. Ariskin et al. / Lithos 302303 (2018) 242262
volcanics is referred to as the Synnyr-Dovyren Complex (Ariskin et al.,
2009), which may be considered as a Late Riphean magmatic province
active during the nal stages of the Precambrian geological evolution
of the southern margin of the Siberian Craton (Kislov, 1998;Konnikov,
1986;Rytsk et al., 2007). This is of particular signicance in the context
of reconstructions of the geodynamic history of the Siberian Craton
during the break-up of Rodinia, which suggest that in the Late Riphean
the Siberian Craton was adjacent to northern Laurentia (e.g., Li et al.,
2008;Metelkin et al., 2009;Pisarevsky et al., 2008). The Synnyr-
Dovyren Complex is contemporaneous to the Franklin event, which
formed the large igneous province in Arctic Canada at 723 ± 4.2 Ma
(Ernst et al., 2016;Ernst and Bleeker, 2010;Heaman et al., 1992).
The proposed spatial and temporal association of the large igneous
provinces (LIPs) in Southern Siberia and Arctic Canada may indicate
their similar metallogenic potential, because many intrusive bodies
from both provinces contain signicant Ni-Cu-PGE mineralisation
(Jowitt and Ernst, 2013;Polyakov et al., 2013). The Synnyr-Dovyren
Complex and surrounding area was considered as a fertile Ni-Cu-PGE
province long before its recognition as the eastern part of the ESMP.
Since the late 1950s, Cu-Ni sulphide ores have been discovered within
the Yoko-Dovyren (in 1959), Chaya (in 1962), and Gasan-Dyakit (in
1964) massifs (see a review of the Northern Baikal nickel reserves by
Kislov, 2010). In 20062016, we carried out eight expeditions to the
DIC aimed at creating detailed studies of the internal structure of the
YDM and its associated sills and volcanics. As a result, a total of 2500
samples have been collected, comprising three complete cross-
sections of the YDM (through the central part and at the margins), sev-
eral cross-sections of the underlying sills, a representative collection of
dykes and Cu-Ni-sulphide ores, and several cross-sections through the
PGE-rich anorthosites and other mineralised sulphide horizons.
Here we summarise the published information of the DIC (Section 2)
and present new data on the structure of the YDM (Section 3) and asso-
ciated sills and volcanics (Section 4), focussing on the geochemistry
of the volcanic-plutonic complex (Section 5)andnewdataonthe
low-sulphide and PGE-enriched rocks. The results allow us to reassess
the internal structure and various petrological aspects of the YDM, in-
cluding Cu-Ni-PGE fertility, the compositions of parental magmas, the
history of sulphide immiscibility in the cumulate rocks, as well as evi-
dence for an open-system behavior of the Dovyren magma chamber
and the nature of an ancient mantle source.
2. Summary of the geology, structure, and fertility of the
Synnyr-Dovyren Complex
2.1. Regional geology
The area northeast of Lake Baikal is part of the Baikal Fold Region
within the eastern Central Asian Fold belt (Parfenov et al., 2010), which
comprises the outerBaikal-Patom Belt at the southern margin of the
Siberian Craton and the Baikal-Muya Belt at the eastern margin of the Pro-
terozoic Circum-Siberia folded area (see inset in Fig. 1). Within the Baikal-
Patom Belt, the basement rocks are exposed within marginal inliers of the
Siberian Craton as blocks of Archean gneisses and Paleo-Proterozoic
anorogenic complexes. The Riphean units of the Baikal-Patom Belt com-
prise several large troughs (e.g., the Patom and Olokit-Bodaibo zones)
and uplifted areas (e.g., the Mama Zone; Konnikov, 1986;Rytsk et al.,
2007). Geodynamic reconstructions relate the formation of the Olokit-
Bodaibo trough to the evolution of the Baikal-Patom paleobasin between
the southern margin of the Siberian Craton and the Baikal-Muya paleo-
arc during the Proterozoic (Konnikov et al., 1999). More specically,
Middle to Late Riphean mature terrigenous sediments, which include tur-
bidites, carbonaceous shales, carbonates, and volcanics, form the 57km
thick Olokit Complex (Rytsk et al., 2002). The structural units within the
Olokit Complex comprise sharply asymmetric troughs separated by
blocks of elevated basement rocks. The basement rocks below the com-
plex are metarhyolites dated at 1863 ± 5 Ma (Neimark et al., 1990).
Late Riphean collisional processes along the southern margin of the
SiberianCraton were accompanied by large-scale sub-horizontal reloca-
tions of individual units towards the northeast direction (in the current
orientation) and resulted in the formation of the ~150 km long and up
to 1215 km wide Synnyr rift. The Late Riphean rocks of the Synnyr
rift (~2.5 km thick) form the top of the Olokit Complex; the uppermost
black shale-dominated part of the complex hosts the DIC.
The 3.5 km thick YDM is the main part of the DIC; it forms a ridge
comprised of Mt. Yoko (SW) and the much larger Mt. Dovyren (Fig. 2).
The YDM is a lens-shaped body with a minimum thickness of ~700 m
at its SW termination and is generally concordant with the host sedi-
mentary units and overlying volcanics. The host sediments dip nearly
vertically around the YDM, thus exposing the entire cross section of
the intrusive and volcanic-sedimentary units. The overlying volcanics
include the rocks of the Inyaptuk and Synnyr suites, which form the up-
permost units of the Synnyr Rift (Kislov, 1998;Rytsk et al., 2002). The
Inyaptuk suite is composed of picritic and basaltic pillow lavas associ-
ated with sub-volcanic bodies of trachydacites and rhyolites, whereas
the Synnyr suite includes basalts, basaltic andesite, and andesite. These
rocks compose the Synnyr Ridge and Mt. Inyaptuk (Fig. 2).
2.2. Geochronology of the DIC and associated volcanics
Hereinafter, we follow rock type specications presented in
Appendices A and B, where analytical methods and whole-rock and
mineral compositions are given.
2.2.1. Age of the Dovyren Intrusive Complex
Amelin et al. (1996) suggested that the YDM formed at 673 ± 22 Ma
based on the Sm-Nd isotopic compositions of mono-mineralic fractions
separated from a gabbronorite from the upper portion of the YDM.
These authors also obtained an age of 707 ± 40 Ma for a gabbronorite
from a sill beneath the YDM, and a Rb-Sr age of 713 ± 7 Ma from a
biotite; however, the Sm-Nd isochron age was considered more reliable.
Recent data on U-Pb baddeleyite dating of a pegmatoidal gabbronorite
from the YDM yielded an age of 724.7 ± 2.5 Ma (Ernst et al., 2016).
Given the uncertainty related to the timing of YDM emplacement,
Ariskin et al. (2013b) performeddetailedU-Pbdatingofzirconfrom11
samples, including the YDM rocks (3 gabbronorites and a recrystallised
hornfels near the roof), a sulphide-rich gabbronorite dyke near its
lower contact, a 200 m thick sill beneath the YDM (the Camel Sill,
5 samples), and an albite hornfels representing low-temperature contact
metamorphic facies within the host rocks. Combined together, these data
yielded 728.4 ± 3.4 Ma (MSWD = 1.8, n = 99) as the age of the DIC
(Fig. 3). This age is consistent with the results of Ernst et al. (2016) and
~55 Ma older than the Sm-Nd isochron age of Amelin et al. (1996).
Fig. 3. Geochronolog y of the Dovyren intrusive rocks and associated vol canics, after
Ariskin et al. (2013b).
245A. Ariskin et al. / Lithos 302303 (2018) 242262
2.2.2. Age of volcanics and timing of overprinted processes
Ariskin et al. (2013b) performed U-Pb zircon dating on two samples
of quartz porphyritic rhyolite from an ~50 m thickdyke cutting through
the black shale at the bottom of the Inyaptuk suite, and on a sample of
an agglomerate quartz porphyritic tuff overlying the 250 m thick
sequence of Inyaptuk basalts. Zircons from the tuff yielded an age
of 721 ± 7 Ma (n = 9, MSWD = 1.3), overlapping with the age of the
DIC. However, the U-Pb system in some of the zircons from the rhyolite
dyke was found to be disturbed. The range observed in the 22 individual
analyses was interpreted to reect the presence of zircon populations
corresponding to two discrete events. The age of the rst population is
729 ± 14 Ma (MSWD = 0.74, n = 8) overlapping with the age of the
DIC, whereas the second population is much younger at 667 ± 14 Ma
(MSWD = 1.9, n = 13), likely corresponding to the timing of the hy-
drothermal alteration that affected the entire Synnyr-Dovyren Complex
(Kislov, 1998; see below). The exact age of the low-Ti Synnyr volcanics
remains unconstrained; however, given their close structural associa-
tion and compositional similarity with the YDM, it is believed that
they are related to the YDM (Ariskin et al., 2009).
The timing of hydrothermal alteration was assessed by Rb-Sr geo-
chronology carried out on leachates derived from partially- to totally-
serpentinised peridotites from the Camel Sill (Appendix A). Thesesam-
ples are characterised by very high Rb/Sr ratios (5.87.1), atypical for
mac and ultramac rocks; the high Rb/Sr values were interpreted to
originate during serpentinisation of these rocks. The results suggest
that the age of the overprinted process is 659 ± 5 Ma (MSWD = 1.3),
which is consistent with the youngest U-Pb age recorded by some of
the zircons from the rhyolite dyke, as discussed above.
2.3. Sr-Nd isotopic compositions
Radiogenic isotopic compositions of Sr, Nd, and Pb have been
analysed in 31 samples from the DIC and associated volcanics (Ariskin
et al., 2015a). The examined samples include 14 mac and ultramac
rocks from the YDM, eight samples from the underlying peridotitic
sills, two samples of gabbronorite sampled in the vicinity of the lower
YDM contact, a contact albite hornfels, three high-Ti basalts, two low-
Ti basaltic andesites, and a low-grade metamorphosed tuffaceous silt
from the exocontact of a rhyolite dyke.
The isotopic composition of the high-Ti basalts is similar to the
mid-ocean ridge basalt (MORB) source at the time of emplacement,
t= 728 Ma (Ariskin et al., 2013b). Unaltered plagiodunite, gabbronorite,
and contact gabbronorite from all three studied sections of the YDM dis-
play a wide range of
Sr(t) ratios (~0.70950.7135) and extremely
anomalous values of ε
(t)(11.5 to 15.5; Fig. 4). This range is consis-
tent with the isotopic composition of olivine gabbronorites from the
ultramac sills and a low-Ti sub-volcanic body underlying the Synnyr
suite.Another low-Ti basaltic andesite has similar ε
less enriched
Sr(t) (~0.7070). This lower value is likely due to
extensive alteration of this rock (Section 2.2.2).
The most important result of these studies is that the maxi-
mum enrichment(
Sr(t) = 0.71387 ± 0.00010 (2σ), ε
16.09 ± 0.06) is found in the lowest marginal rocks of the YDM intru-
sion, including the chilled picrodolerite (Fig. 4). As far as these rocks
represent crystallisation products of the most primitive high-Mg
magmas of the YDM (Section 6.1), their anomalous isotopic composi-
tions can be used to propose an ancient mantle source of the parental
DIC magmas, as it was argued in Ariskin et al. (2015a) and inferred
from the results of this study.
2.4. Main types of YDM cumulate rocks
The YDM was rst described in the 1950s during a geological survey of
the Northern Baikal region (see Kislov, 2010). This was followed by sam-
pling of representative sections through the intrusion (Balykin et al.,
1986;Emov and Potapova, 2003;Grudinin, 1963, 1965;Gurulev, 1965,
1983;Kislov, 1998;Konnikov, 1986;Konnikov et al., 2000;Yaroshevskii
et al., 1982), accompanied by ne-scale mapping of selected areas within
the central YDM (Kislov, 1998;Konnikov et al., 2000;Yaroshevskii et al.,
1982). The best described cross-section, starting from the Bolshoi and
Tsentralnyi creeks in the bottom part of the intrusion (see sections Ia
and Ib in Fig. 2), is shown in detail in Fig. 5. This generalised cross-
section, which also includes our new data (Section 3), is used to highlight
the most important features of the main YDM rock types.
The generalised cross-section through the central YDM (hereafter
the Bolshoi-Tsentralnyi section) consists of (I) a 35 m thick layer of
chilled picrodolerite at the lower contact with hornfels, followed by
(II) a basal unit of a plagioclase peridotite up to 150170 m thick
(Gurulev, 1965; Konnikov, 1986). The plagioperidotite gradually transi-
tions into (III) plagiodunite (4060 m), followed by (IV) a thick layer
of low porosity adcumulate dunite (up to 970 m). The dunite zone tran-
sitions into (V) a layered sequence of adcumulate melano- to
leucotroctolite (~950 m), which contains several layers of
clinopyroxene-bearing (up to 57% Cpx) troctolite or olivine gabbro
(Kislov, 1998) in the upper portion (forming the crest of the Dovyren
ridge). Starting from this horizon, the upper part of the YDM contains
both veins and schlieren of gabbro-pegmatite, granophyre, and
coarse-grained anorthosite (Fig. 6). Along the southeastern slope of
the ridge (~2200 m from the lower contact of the intrusion, Fig. 5),
modal clinopyroxene increases rapidly leading to a transition from
troctolite to (VI) the olivine gabbro zone, ~450 m thick. This is followed
by (VII) interbedding of olivine norite and olivine gabbronorite (up to
600 m), which transitions into (VIII) a sequence of pigeonite gabbro
and quartz-granophyre gabbronorite (~220 m) near the roof of the in-
trusion. A thin unit of ne-grained gabbronorite marks the upper
YDM contact. The transitions between the main units of the intrusion
commonly involve rhythmic intercalation of various rock types. A hori-
zon of poikililic Cpx-containing dunite within the troctolite zone and
layers of olivine gabbro among gabbronorite are typical examples of
such intercalation.
The detailed structure of the lower contact zone is complicated by
the presence of tabular bodies of mactoultramac rocks, which likely
represent multiple, nearly coeval intrusive events at the onset of the
Fig. 4. SrandNdisotopiccompositionsofigneousandsedimentaryrocksfromtheSynnyr-
Dovyren complex. All values are corrected to the DIC age of 728 Ma (Ariskin et al., 2013b).
The volcanic rocks were sampled along traverses III and IV in Fig. 2.Underlying sills
and Serpentinised peridotiteinclude samples from the Camel Sill (Section 4.1). Two
sedimentary rocks (referred to as dolomitesin the legend) represent dolomitised marbles
from the host sedimentary sequence (Amelin et al., 1996). Two samples denoted as
hornfelsinclude albite hornfels 07DV163-1 (Ariskin et al., 2015a) and a metamorphosed
siltstone shale, TST-28 (Amelin et al., 1996). The MORBeld is from Kostitsyn (2004, 2007).
246 A. Ariskin et al. / Lithos 302303 (2018) 242262
main emplacement stage. These bodies (often referred to as sills)havea
thickness of several tens to 200m and are locally separated by country
rocks (Fig. 2). In some places, these sills form an interconnected sequence,
allowing for their consideration as apophyses from the basal zone of the
intrusion, making it difcult to locate the exact position of the lower
YDM contact. A further complication is the occurrence of several large-
scale faults cutting the intrusion across and along its strike (Kislov, 1998).
2.5. Magnesian skarns in ultramacs
A distinguishing feature of the YDM is the occurrence of numerous
xenoliths of magnesian skarns in dunite and troctolite that are
interpreted to be theproducts of in-situ assimilation of marbles and do-
lomites by high-Mg magmas during emplacement (Gurulev, 1983;
Pertsev and Shabynin, 1979). According to Wenzel et al. (2002), rapid
heating of the xenoliths by mac magma resulted in decomposition of
carbonates, releasing CO
and CaO into the magma, thus expanding
the stability eld for Ca-rich clinopyroxene in the surrounding rocks.
The magnesian skarns occur as disturbed leucocratic xenoliths, several
cm to several m in size, or as undisturbed layers up to 100150 m
long (Fig. 6). In the central YDM these xenoliths are most abundant at
adistanceof200400 m below the contact between dunite and
troctolite. Mineralogy of the metamorphic rocks is dominatedby brucite
pseudomorphs after periclase, forsterite, and aluminous spinel; the
Fig. 5. Structure ofthe Bolshoi-Tsentralnyi section in the centreof the Yoko-Dovyren massif.Mineral proportionsrepresent normative contents, calculated from whole-rock compositions
assuming Fe
/ΣFe = 0.05. S and V concentrations are from our X-ray uorescence (XRF) analyses (Appendix B, Sheet Rocks), except for the Main PGE-Reef samples, with bulk rock S
concentrations taken from Tolstykh et al. (2008).
Fig. 6. Xenoliths of Mg-rich skarns within dunite (A); schlieren of anorthosite (B); and a gabbro-pegmatite (C) within the gabbroic part of the YDM.
247A. Ariskin et al. / Lithos 302303 (2018) 242262
ne-grained Fo-Sp assemblages can also be associated with the brucite
skarns, or observed as isolated schlieren (Pertsev and Shabynin, 1979;
Wenzel et al., 2002). It was suggested that these xenoliths mark the
pre-intrusion position of the sedimentary carbonates, with their initial
location remaining almost undisturbed (Konnikov, 1986).
2.6. Cu-Ni-PGE mineralisation within the DIC
The Cu-Ni sulphide mineralisation within the YDM and associated
intrusive rocks ranges from widespread disseminated sulphides to
net-textured and massive pyrrhotite-rich ores at the lower YDM contact
and within the underlying ultramac sills and gabbronorite dykes
(Ariskin et al., 2016;Kislov, 1998). YDM mineralisation was rst
described in 1949; however, it was only in 19591963 that this
area was rst subjected to detailed geological prospecting (Tolstykh
et al., 2008). Limited-scale surveys were conducted in 19761979 and
19861993, leading to the establishment of a subeconomic so-called
Baikal Deposit, which contains a total of 147, 000 tons Ni, 51,000 tons
Cu, and 9500 tons Co (Kislov, 2010). The deposit contains both
syngenetic and epigenetic Cu-Ni sulphide mineralisation.
2.6.1. Syngenetic Cu-Ni mineralisation
The syngenetic mineralisation varies from nely-disseminated in-
terstitial sulphides (0.53 mm in size, commonly b1% sulphide in the
rock) to impregnated(Ariskin et al., 2016)orglobularand net-
texturedores (Fig. 7A), following terminology by Barnes et al. (2017).
All these types are irregularly distributed throughout the entire YDM
and associated sills (Kislov, 1998). The most sulphide-rich occurrences
(up to 30% sulphides) are typical for ultramac bodies below the
lower contact. These ores generally occur where dykes of gabbronorite
cut plagioperidotite and olivine gabbronorite within the sills. Outcrops
of this type of sulphide ore were traced continuously for ~1700 m paral-
lel to the lower contact in the YDM centre; the width of the ore lenses
varies from 8 to 25 m, locally reaching 80 m (Kislov, 1998). Where it
occurs in the gabbronorite, the Cu-Ni mineralisation is observed as
disseminated pyrrhotite-rich sulphides that may produce branchy
blebs up to 20 mm in size.
Another type of syngenetic mineralisation includes poorly-
mineralised rocks with ne Ni-rich (6095% pentlandite) interstitial
sulphides, which occur in a ~150 m horizon within the transition zone
from plagiodunite to adcumulate dunite (Fig. 5 and Ariskin et al.
(2016)). Near the YDM roof, pyrrhotite-rich sulphide mineralisation is
observed in olivine-free gabbronorite and quartz-pigeonite gabbro,
with S concentration reaching as much as 3 wt% S (Fig. 5).
2.6.2. Epigenetic Cu-Ni mineralisation
Epigenetic mineralisation is observed in metasomatized and essen-
tially recrystallised rocks, generally as massive and vein-like brecciated
ores (Fig. 7B), locally occurring within widespread domains of net-
textured sulphides (Kislov, 1998). The largest sulphide lode was discov-
ered in 1959 at the NE YDM contact (the Ozernyi prospect; Denisova,
1961). It extends along the base of the YDM for a distance of 650 m,
and is 0.71.0 m wide. Smaller lodes (1550 m long and 0.21.5 m
thick) are conned to tectonic transgressions demarcated by sills of
plagioperidotite and associated dykes. Drilling has demonstrated that
the sulphide-rich veins dip nearly vertically and extend to depths of
N500 m (Kislov, 2010). The massive ores consist of 7095% pyrrhotite
and contain minor amounts of troilite, pentlandite (725%), and chalco-
pyrite (0.16%). These ores contain up to 2.1 wt% Ni, 0.64 wt% Cu, and
0.14 wt% Co (Kislov, 1998).
2.6.3. PGE-rich sulphide-poor mineralisation in anorthosite
Low-sulphide PGE-rich mineralisation within the YDM was rst
documented within anorthosite near the transition from troctolite to
olivine gabbro (Distler and Stepin, 1993;Konnikov et al., 1994;Kislov
et al., 1995;Orsoev et al., 1995;Fig. 5). Later, a discontinuous zone of
PGE-rich sulphide-poor mineralisation was traced near the basement
of the gabbroic subsection along the YDM strike for over 20 km
(Kislov, 1998). The highest PGE contents (up to 12 ppm) are found in
the central part of the intrusion (the Main Reef) within a 150200 m
zone composed of concordant veins and lenses of coarse-grained and
Fig. 7. Sulphide mineralisation of the Dovyren Intrusive Complex. A Net-textured Cu-Ni sulphide ore from a sill below the YDM (sample 07DV107-1); B Low-mineralised PGE-rich
anorthosite from the Main PGE Reef (07DV146-1); C A massive Po-rich sulphide ore from the Ozernyi prospect at the NE termination of the YDM (sample provided by D.A. Orsoev,
Geological Institute in Ulan-Ude, Buryatia, Russia); D and E Back-scattered electron (BSE) images of Pt-Pd-Ag minerals associated with sulphides: D Moncheite in a thin anorthositic
vein (the Main PGE Reef, sample 07DV146-2); E A composite grain of moncheite (light) and telargpalite (grey) in a low-mineralised troctolite (Section 2.6.4).
248 A. Ariskin et al. / Lithos 302303 (2018) 242262
taxitic troctolite, olivine gabbro, as well as minor leucogabbro and
gabbronorite. This highly heterogeneous zone hosts numerous bodies
of barren (Fig. 6B) and mineralised anorthosite (Fig. 7B), which
occur as large schlieren and lens-like bodies commonly surrounded
by sulphide-free or poorly-mineralised gabbro-pegmatite (Fig. 6C).
Anorthosite bodies are usually a few cm to 1 m thick and extend for 2
to 5 m along the massif strike (rarely N40 m long), forming a discontin-
uous sulphide-poor mineralised zone.
The sulphide assemblages are Cu-rich, including chalcopyrite,
cubanite, bornite, and in rare cases talnakhite, heazlewoodite, and
godlevskite (Konnikov et al., 2000). Pentlandite and pyrrhotite
are minor phases (b1520 vol% in total). Using a combination of
hydroseparation of nely-crushed sulphide-poor anorthosite (Orsoev
et al., 2003;Rudashevsky et al., 2003) andscanning electron microscopy
(SEM) studies of the mineralised samples (Tolstykh et al., 2008),
approximately 80 grains of precious metal minerals were found. The
PGMs range in size from 1 to 2 to 60 μm, with moncheite (Fig. 7D),
potarite, and tetraferroplatinum being predominant. Minor PGMs in-
cluded kotulskite Pd(Te,Bi,Pb), sobolevskite Pd(Bi,Te), and a number
of Ag-minerals (argentite Ag
S, stephanite Ag
, amalgam AgHg).
Additional studies of three anorthosite samples from the Main Reef re-
vealed the presence of 22 PGM grains(~235 μm), including moncheite
, paolovite (Pd,Pt)
(Sn,Te), atokite (Pd,Au)(Sn,As,Sb),
and merenskyite Pd(Te,Bi)
(Ariskin et al., 2016).
Despite the sulphide-poor character of the Main Reef
mineralisation, the amount of PGE-bearing sulphides in the anortho-
site may locally reach up to 35 vol%. Due to such irregular distribu-
tion of sulphides, the total concentration of PGE + Au in the rock
varies in the range 0.312.1 ppm at 0.0060.710 wt% Cu and 0.023
0.430 wt% Ni (Tolstykh et al., 2008). Laser ablation inductively
coupled plasma mass spectrometry (ICP-MS) analysis of sulphide
phases from the PGE-mineralised anorthosite revealed very high
concentrations of Pd in pentlandite (235 ± 84 ppm) which also con-
tains 6.9 ± 6.5 ppm Rh, 1.9 ± 2.0 ppm Ir, 0.93 ± 0.64 ppm Ru, 0.55
± 0.30 ppm Os, 0.30 ± 0.12 ppm Re, 34.8 ± 12.1 ppm Ag, and 25.1
± 15.5 ppm Te (Ariskin et al., 2016).
2.6.4. Recently discovered PGM-containing troctolite
Geochemical and mineralogical evidence for the presence of PGM-
bearing low-sulphide occurrences within troctolite was rst presented
by Ariskin et al. (2015b). A troctolite collected ~120 m above the
dunite/troctolite boundary along the stream of Tsentralnyi Creek
contained Cu-rich sulphides with minor pentlandite, which is highly
enriched in Pd (91 ± 84 ppm, n = 7; max 250 ppm). This is similar to
the average 235 ppm Pd value observed in pentlandite from the Main
PGE Reef. High-resolution SEM analysis revealed the presence of
12 grains of PGMs (mainly Pd bismuth-tellurides, 19μm in size) and
two grains of electrum (up to 8 μm in size). A more detailed sampling
of the lower troctolite horizon revealed a zone of weakly-altered
mesocratic troctolite ~250 m above the dunite/troctolite boundary,
which contained disseminated round silicate-sulphide clusters, several
cm to ~10 cm in diameter.
In thin sections, these heterogeneous schlieren displayed typical
orthomagmatic intercumulus sulphides lling the pore space between
olivine and plagioclase crystals (Ariskin et al., 2015b). Similar to the
Main PGE Reef, chalcopyrite and cubanite accounted for ~7580%
of the total sulphide in the samples. As many as 60 PGM grains (1 to
40 μm in size) were identied in two samples from this new occurrence.
Further analyses of the 30 largest grains by high-resolution SEM re-
vealed the presence of PGM assemblages (Fig. 7E), which are generally
similar to those observed in the Main Reef anorthosite. Tellurides and
bismuthotellurides of Pd, Ag, Pt, and Pb are the most abundant phases
(moncheite, paolovite, atokite, merenskyite, michenerite Pd
kotulskite Pd(Te,Pb,Bi), and telargpalite (Pd,Ag)
Te). In addition, native
alloys/amalgams, such as electrum (Au-Ag), tetraferroplatinum (PtFe),
and potarite (PdHg) are present. Minor PGMs include zvyagintsevite
Pb and taimyrite (Pd,Cu)
Sn. Chemical analyses of pentlandite from
these samples demonstrated high concentrations of all PGE with the no-
table exception of Pt (all values in ppm): Pd 23.1 ± 7.2, Rh 23.1 ±7.2, Ru
7.5 ± 4.7, Os 4.2 ± 3.3, and Ir 6.7 ± 4.0 (Ariskin et al., 2015b).
3. Detailed structure of the Yoko-Dovyren massif
We summarise below the petrographic descriptions of Kislov (1998)
and Ariskin et al. (2009, 2016) and present new data on chemical com-
positions of minerals and rocks from three cross-sections through the
YDM, which are important for the following discussion on the origin
and parental magma compositions of the DIC.
3.1. The Bolshoi-Tsentralnyi cross-section
3.1.1. Chilled rocks at the lower contact
Directly at the lower YDM contact, a ~35 m thick zone is composed
of magnesian gabbronorite and picritic rocks (compositionally olivine
gabbronorite) with variable amounts of cumulus olivine and sub-
ophitic textures of their groundmass. The chilled gabbronorite is
composed of ne-grained brous aggregates of altered plagioclase
and clinopyroxene, bronze to brown mica (mostly phlogopite with
MgO/FeO N2), and rare resorbed olivine crystals. Plagioclase and
clinopyroxene commonly form zoned laths, whereas phlogopite mostly
forms oikocrysts. Rare idiomorphic grains of chromite and apatite are
also present.
Transition from the chilled gabbronorite to the more ultramac
lithologies inward of the intrusive body is marked by the appearance
of subhedral, more-abundant grains of olivine (0.42 mm), resulting
in a porphyritic texture. A large amount of thin plagioclase laths
leads to the ophitic texture of the groundmass (Fig. A1A in Appendix
A). In the Russian literature these rocks were variably described as
ophitic gabbro(Gurulev, 1983)orpicrodolerite(Kislov, 1998), fo-
cussing on the gradual change from aphyric rocks with 1011 wt%
MgO at the very contact, towards olivine-phyric coarse-grained
rocks with 1722 wt% MgO several meters above (Fig. 8, Appendix
B, Sheet Rocks). The textural features indicate that picrodolerite
formed via fast cooling of a high-Mg magma, suggesting that the
rocks are orthocumulates crystallised from olivine-laden YDM par-
ents (Ariskin et al., 2003, 2009). Typical mineralogy of the
picrodolerite is shown in Fig. 9, with details given in Appendix A
and average compositions listed in Appendix B.
3.1.2. Basal plagioperidotite
Olivine-rich ortho- and mesocumulates, often referred to as
plagioclase peridotite(Kislov, 1998)orplagioclase lherzolite
(Tolstykh et al., 2008), contain up to 37 wt% MgO (Appendix B, Sheet
Rocks) and comprise a ~160-m-thick basal unit. In fact, these rocks
represent a diversity of olivine gabbronorite ranging from moderately
(2540 vol% Ol) to highly melanocratic varieties (6575 vol% Ol). This
type of rock rst appears within 510 m from the contact and displays
rapidly cooled ophitic textures, although less magnesian picrodolerite
is also present at that level. The transition from picrodolerite to
plagioperidotite is recorded by a gradual decrease in phlogopite and in-
crease in olivine, which is accompanied by a transformation of the
groundmass texture from ophitic to poikilitic and then hypidiomorphic.
Cumulus olivine compositions vary between samples (Fo74-86),
whereas intercumulus material contains of compositionally-variable
plagioclase (An52-86), magnesian clinopyroxene (mg#86.4 ± 1.3),
othopyroxene (mg#84.4 ± 1.0), and phlogopite (mg#80-86), as docu-
mented in Appendix B. Cr-spinel occurs as inclusions in olivine and in
the intercumulus (see variations in cr# and mg# in Fig. 9).
The chemical compositions of the fresh and moderately-altered
(LOI b5%) plagioperidotite allowed their separation into two major
types (Ariskin et al., 2016): the high-mg# Type I and the low-mg#
Type II. This is clearly seen on the FeO-MgO diagram (Fig. 8B), where
249A. Ariskin et al. / Lithos 302303 (2018) 242262
Type I rocks form a trend towards the composition of olivine ~Fo88
(Fo88-trend), whereas Type II compositions extend towards olivine
~Fo86 (Fo86-trend). Other major elements presented in terms
of MgO display well-dened single trends suggesting that the rocks
reect variable degrees of olivine accumulation in the primitive melts
(e.g. Al
vs. MgO in Fig. 8A). All samples from the chilled margin
plot together with Type I plagioperidotite along the Fo88 trend
(Fig. 8B). In addition to plagioperidotite, the basal unit includes rare
thin veins (several cm) and thicker layers (up to 12m) of more
leucocratic gabbronorite with up to 6070 vol% plagioclase. Fe-Mg
silicates in these rocks commonly have mg# b82-83, indicating that
the rocks represent a more evolved magmatic material formed during
in-situ solidication of the host peridotitic unit.
3.1.3. Plagiodunite
The collective term plagiodunite(or Pl-dunite) is applied to olivine-
rich rocks (8085 vol% Ol, 1012 vol% Pl, b56vol% Cpx+OPx;
Fig. A1D), which compose an horizon up to 60 m in thickness that
overlies the plagioperidotite horizon. A gradual transition between
these two types occurs over an interval of 2030 m, where the amount
of pyroxene oikocrysts decreases up the section. These changes in min-
eralogy are correlated with depletion in Ti, K, P, and other incompatible
elements (Appendix B, Sheet Rocks), suggesting that these changes
are due to decreasing porosity of the Ol-rich cumulates (Ariskin et al.,
2009). Olivine compositions vary between Fo84-87 (Fig. 9), with plagio-
clase (An70-77) becoming the main poikilitic mineral (Fig. A1D).
Occasional clinopyroxene is present in some interstices, whereas
orthopyroxene is observed as thin rims around olivine crystals or as
very rare oikocrysts. Cr-spinel is texturally and compositionally similar
to plagioperidotite, whereas sulphide is rare (Appendix A). Similar to
plagioperidotite, the whole-rock Pl-dunite compositions fall onto the
Fo88 and Fo86 trends (Fig. 8B). Based on a combination of the petro-
graphic features and the major element data, we dene the transition
from the plagiodunite to the overlying Pl-containing dunite (b57% Pl)
as MgO 3940 wt%, calculated on the volatile-free basis.
3.1.4. Dunite zone
In the central YDM, a thick zone of dunite is present (up to 970 m,
Fig. 5). These rocks contain 9097 vol% olivine. The boundary between
plagiodunite and dunite is gradual, with the amount of intercumulus
Fig. 8. Whole rock major element compositions of mac to ultramac rocks from the YDM basal units in the centre of the intrusion (see Appendix B, Sheet Rocks). Chilled rocks
(gabbronorite to picrodolerite), plagioperidotite and pla giodunite, sampled at the distance 0.1320 m from the lower contact, form two distinct trends in the FeO-MgO diagram,
indicating accumulation of olivine ~Fo88 in Type I rocks (Fo88-trend) and ~Fo86 in Type II rocks (Fo86-trend). Except for seven chilled rocks (green stars), other ones are interlayered
in a scale of tens meters. Note that the two types cannot be distinguished from their petrography (Fig. A1B and C) or using other major elements (Al
in Fig. 8A is shown as an
example). Modelled melts represent two compositions calculated at 1285.4 and 1189.7°C (see Section 6.1).
Fig. 9. Variations of mineral compositions along the Bolshoi-Tsentralnyi cross-section. The complete mineral compositions are listed in Appendix B; mg# = MgO/(MgO + FeO),
cr# = Cr/(Cr + Al).
250 A. Ariskin et al. / Lithos 302303 (2018) 242262
decreasing upward. The lowest 2050 m of the dunite zone are com-
posed of layers of poikilitic rocks relatively enriched in the plagioclase
intercumulus (hereinafter dunite I rocks, which are locally similar to
plagiodunite) separated by horizons of hypidiomorphic granular dunite
(Fig. A2A). As the amount of cumulus olivine continues to increase,
panidiomorphic texture develops, displaying linear boundaries of
olivine crystals, which meet at ~120° triple points (Fig. A2B; dunite II).
This is typical of adcumulates and is accompanied by a further decrease
in the proportion of intercumulus phases from 10 to 15 vol% to b5vol%.
At ~450 m above the lower contact, the dunite is effectively a
monomineralic olivine adcumulate with chromite inclusions, rare
plagioclase interstices, and very thin rims of pyroxene. Their chemical
compositions are characterised by MgO N44 wt% and Al
an anhydrous basis (Appendix B, Sheet Rocks).
This olivine adcumulate comprises a texturally homogeneous zone
up to 400 m in thickness (Fig. 5). The entire dunite succession displays
minor olivine compositional variations with a subtle tendency towards
upward increase in Fo contents (Fig. 9; Appendix A). In several cases,
Cpx-lled interstices were found in the lowest porosity olivine cumu-
lates. Such clinopyroxene is generally rich in diopside (Di) (b0.5 wt%
; very high mg#94-97.5 and CaO 25.626.6 wt%; Appendix B,
Sheet CPX). The amount of Di-rich clinopyroxene increases towards
a zone of highly-contaminated dunite and reactive wehrlite, where
such high-Ca pyroxene may coexist with a fassaite-like Cpx. Plagioclase
composition in the lower dunite I (An 69-81) is commonly less calcic
compared to that in the overlying adcumulate dunite II (~An78, see
Appendices A and B).
On the Al
-MgO diagram (Fig. 8A), the dunite I and dunite II com-
positions extend the general trend dened by plagioperidotite and
plagiodunite to lower Al
and higher MgO contents. In the FeO-MgO
diagram (Fig. 8B) these two types of dunite reveal subparallel trends
with respect to the olivine stoichiometry linearisation, with dunite I
rocks spanning the whole compositional range between Type I and
Type II plagiodunite. This may indicate a genetic link between dunite I
and the underlying plagiodunite via a process that combines in situ
crystallisation of the original ortho- to mesocumulates followed by
their compaction (Ariskin et al., 2009). The higher Fo in dunite II likely
reects the localised assimilation of carbonates, which is recorded
in the anomalous mineral compositions (Fig. 9). It is also recorded in
relatively low NiO (less 0.15 wt%) at higher Fo (8890 mol%) in olivine
from dunite and wehrlite closely associated with xenoliths of magne-
sian skarn (see olivine compositions at 974 and 975 m heights in
Appendix B).
3.1.5. Highly contaminated dunite
The impact of carbonate assimilation on dunite II compositions was
assessed by detailed analyses of rocks along a ~200 m long section
starting from ~560 m above plagiodunite. This section includes a
gradual transition from weakly contaminateddunite II through to
highly contaminateddunite, Di-containing reactive wehrlite and
diopsidite, as well as the uncontaminated dunite II adcumulate above
this zone of magnesian skarns. Overall, major element rock and mineral
compositions are similar in contaminated and uncontaminated dunite II
(Appendix B).
Distinctive features of the contaminated dunites are the very high
CaO contents in olivine, which may exceed 1 wt% (Fig. 9), and the
gradual enrichment in the amount of Di-rich clinopyroxene towards
contacts with magnesian skarns. Near the contacts, a fassaite-rich
clinopyroxene with 5.56.6 wt% Al
is common. The amount of
Di-rich clinopyroxene in the highly-contaminated rocks may exceed
2030 vol%, resulting in dunite grading into reactive wehrlite with a
poikilitic texture (Fig. A2C), which locally contains small inclusions of
calcite. In the sampled area, the strongly serpentinized wehrlite horizon
is ~25 m thick. Locally, Di-rich clinopyroxene forms patches and irregu-
lar veins of diopsidite within wehrlite and dunite, giving rise to a taxitic
texture in the contact rocks.
The highly-contaminated dunitealso displays relatively low-cr# and
high-mg# spinel compositions compared to spinels in the uncontami-
nated dunite II (Fig. 9 and Appendix B, Sheet SPIN). The most Cr-
depleted spinel occurs 2050 m above the wehrlite horizon. Schlieren
of even more aluminous chromitites (cr# 0.22-0.23) are abundant
within the contamination zone (Pushkarev et al., 2004). These schlieren
are 1020 cm wide and up to 0.51 m long and contain up to 4060 vol
%of35 mm idiomorphic Al-rich chrome spinel crystals (Fig. A2D).
Above the wehrlite, the amount of Di-rich pyroxene in the contami-
nated rocks decreases together with decreasing Ca and Al in Cpx, until
the mineral chemistry becomes similar to that from uncontaminated
dunite II below the horizon of carbonate xenoliths. This change is
accompanied by a decrease in Fo and CaO contents in olivine (Fig. 9).
The uppermost dunite of the dunite zone is generally similar to that
classied as normalor weakly contaminateddunite II below the con-
taminated horizon.
3.1.6. Adcumulate and Cpx-bearing troctolite
The contact between dunite and troctolite is located ~1200 m above
the lower YDM contact (Fig. 5), with only one mapped outcrop near the
source of Tsentralnyi Creek in the central YDM (prole Ib in Fig. 2). This
contact appears as a transitional zone ~100 m thick, marked by the
rst occurrence of intercalated layers of melanoctratic troctolite and
Pl-bearing dunite. The thickness of these layers varies from tens of cm
to several meters. Above this level, the amount of troctolite (see tex-
tures in Fig. A3), thickness of individual layers, and the amount of pla-
gioclase in troctolite all generally increase, reaching their maximum
next to the upper troctolite boundary. However, dunite-like rocks still
occur through the upper troctolite zone (Fig. 5). In its upper part,
these rocks form several narrow horizons enriched in clinopyroxene
(poikilitic plagiowehrliteafter Konnikov, 1986). The taxitic texture of
the plagiowehrlite is distinctly different (Fig. A3D). It is composed of
8590% olivine and a matrix of Cpx-oikocrysts with rare poikilitic
Most olivine crystals demonstrate a panidiomorphic granular
texture; however, smaller, rounded, likely resorbed olivine grains are
also present within clinopyroxene. The composition of plagioclase is
similar to that of the hosting troctolite. Overall, such dunite-like rocks
are similar to troctolite in their mineral chemistry (Fig. 9; see Appendices
A and B for details). This is evidence for the plagiowehrlite to have ini-
tially been a higher porosity Ol-Pl cumulate depleted in the cumulus pla-
both plagiowehrlite and leucotroctolite to thelowermost anorthosite are
principal components of the troctolite unit, which shows extreme phase
separation in its upper transitional zone (Figs. 5 and 9).
3.1.7. Zone of olivine gabbro
The gabbroic part of the YDM above the troctolite zone is ~1200 m
thick and can be subdivided into three units. The rst unit is ~430 m
thick and begins directly above the main PGE-rich horizon as an
interbedding of rare layers of Cpx-containing troctolite and more abun-
dant olivine gabbro (Fig. 5). It is characterised by an upward decrease
in the amount of olivine and general enrichment in clinopyroxene,
with the Cpx maximum observed at ~300 m above the irregular contact
with troctolite. Another distinctive feature of the olivine gabbro is
the occurrence of numerous schlieren of anorthosite and veins of
leucogabbro and gabbro-pegmatite with a taxitic texture. On average,
the amount of the anorthosite schlieren decreases upwards, whereas
veins of gabbro-pegmatite become more abundant.
The textural position of clinopyroxene in the olivine gabbro is vari-
able. In the lowermost Cpx-poor rocks this mineral occurs mostly as
an unevenly distributed intercumulus material (Fig. A4B), whereas in
the Cpx-rich varieties it is present as relatively large sub-idiomorphic
poikilitic grains with rims that occupy the interstices of olivine and
plagioclase crystals. Unlike ultramac YDM rocks, olivine gabbro is
characterised by a nonmonotonic upward decrease in the mg# of
251A. Ariskin et al. / Lithos 302303 (2018) 242262
mac minerals and the An content of plagioclase (Fig. 9). A distinctive
feature of the olivine gabbro and other gabbroic rocks in the centre of
the YDM is the absence of Cr-Al spinel. Only few small spinel grains
(mg#13-30, cr#72-79) were found in plagioclase from the lowermost
olivine gabbro.
3.1.8. Olivine norite to gabbronorite and near-roof rocks
Above the olivine gabbro, there is an ~530 m thick zone, which con-
tains intercalated horizons of olivine norite and olivine gabbronorite.
These units demonstrate a complex layering dominated by olivine
gabbronorite. Both types of rocks display gabbroic textures (Fig. A4AC).
The mineral assemblage includes ovoid olivine grains (Fo76-79),
short tabular plagioclase crystals (An81), and prisms of clino- and
orthopyroxene. The occurrence of large oikocrysts and small subhedral
grains of orthopyroxene is typical; rare phlogopite is also present.
Clinopyroxene has mg#~81 and orthopyroxene mg#~79.
At a distance ~250 m below the upper YDM contact, olivine norite
and gabbronorite transitioninto a sequence of leucocratic quartz gabbro
and granophyric gabbronorite (Figs. 5 and 9). These rocks include
both intercumulus and sub-idiomorphic grains of inverted pigeonite
as their common feature. The low-Ca clinopyroxene is distinguished
by a characteristic texture originated via subsolidus decomposition
of the original solid solution into lamellae of high-Ca clinopyroxene
hosted by orthopyroxene (Fig. A4D). These rocks are dominated by
tabular grains of zoned plagioclase (An70-76), prisms and oikocrysts
of hypersthene (mg#70-71), and subhedral grains of low-Ca to high-
Ca clinopyroxene (mg#75-76). Amphibole and brown phlogopite ll in-
terstices between these cumulate grains. Quartz-orthoclase granophyre,
ilmenite, and apatite are also common. Schlieren of granophyre and
gabbro-pegmatite are widespread throughout the near-roof zone, com-
monly cutting across gabbroic units.
Closer to the upper contact, the quartz-pigeonite gabbro becomes
texturally more ophitic. Theuppermost ~30 m of theYDM are composed
of the ne-grained rocks classied as upper chilled gabbronorite; they
are more evolved here than at the lower contact.
3.2. Structure of the YDM at the NE and SW margins
3.2.1. The Schkolnyi section
The Schkolnyi section (~1345 m thick; prole II in Fig. 2)represents
the northeastern termination of the YDM. Due to limited exposure in
this area,this prole is a combined cross-sectionconstructed usingsam-
ples collected froma number of exploration trenches (up to 200 m long)
at the NW slope of Mt. Dovyren, and samples from rocky outcrops at the
SE slope of Mt. Dovyren (the upper YDM). The sampling interval along
the Schkolnyi section range from 10 to 15 to 2030 m. Overall, the
Schkolnyi section is composed of two rock types: ~2/3 are ultramacs
(mostly melanotroctolite) and the remainder includes leucocratic gab-
broic rocks (Fig. A5). Based on detailed mineralogy and rock texture
(Appendix A, Section A4.1), the Schkolnyi section is subdivided
into eight zones, including: (1) lower chilled gabbronorite (~5 m)
(2) plagioperidotite (~100 m) (3) plagiodunite (~80 m) (4) inter-
calation of plagiodunite and melanotroctolite (~350 m) (5) melano-
to mesotroctolite (~390 m) (6) Ol-containing leucogabbro and
gabbro-pegmatite (~230 m) (7) Ol-free gabbronorite (~110 m)
(8) quartz-pigeonite and granophyre gabbro replaced by the upper
chilled gabbronorite at the upper contact (~80 m).
3.2.2. The Yoko section
The Yoko cross-section at the SW termination of the YDM was com-
bined from 32 samples collected along two sub-parallel traverses across
the intrusive body (~2200 m thick; see traverses IIIa and IIIb in Fig. 2).
Most of the rocks are veryfresh and display unaltered original magmatic
textures. No single outcrop of the lower YDM contact with hosting rocks
was found in this area.However, a gabbronorite dyke ~2030 m below a
suggested contact location (see sample 07DV220-1in Appendix B, Sheet
Rocks) was proposed to be a compositional proxy of a liquid portion of
a parental Ol-laden magma in the YDM area (Ariskin et al., 2015a).
Similar to the Shkolnyi prole, the Yoko section can be divided into
mac and ultramac parts; however, the relative proportions of these
types are very different (Fig. A6 in Appendix A). The amount of Yoko
gabbroic rocks is higher, with the rst olivine gabbro occurring already
within the lower third of the intrusive body. The sampled cross-section
starts from a 150 m thick unit of sulphide-bearing plagiodunite,
followed by a thick zone of intercalated troctolite and olivine gabbro.
At ~1350 m the cumulate succession is dominated by leucogabbro
and gabbronorite, similar to the upper intrusion zones in other YDM
There are two distinctive mineralogical features in the Yoko section:
1) unlike other cross-sections, Ol-rich gabbro from Mt. Yoko contains
Cr-rich spinel as inclusions in olivine and intercumulus phases;
2) only the uppermost gabbronorite represents Ol-free rocks that are
similar to quartz-pigeonite gabbro from other cross-sections of the
YDM (sample 07DV230, Appendix B). The most magnesian olivine
Fo86 was found in a sulphide-poor plagiodunite from the basal zone
of the Yoko section. This composition is similar to the most magnesian
olivine from the Schkolnyi area. There is no higher-Fo olivine, as
the one documented in the Bolshoi-Tsentralnyi section. This may indi-
cate that both NE and SW parts of the intrusion are composed of
crystallisation products of a parental magma that is more evolved
than that in the centre of the intrusion.
4. Underlying sills and associated volcanics
In addition to the YDM, the DIC includes a number of mac-
ultramac sills 10 m to 200 m thick, which are generally sub-parallel
to the lower YDM contact (Gurulev, 1965, 1983;Kislov, 1998). Geolog-
ical mapping has demonstrated that these sills are separated from
the bottom of the YDM by beds of hornfels and siltstones. However,
some of the bodies may be considered as apophyses from the intrusion
(Fig. 2). Numerous dykes of leucocratic gabbronorite and olivine
gabbronorite are closely associated with sills, generally cutting both
the ultramac sills and the host rocks (Fig. 10). This mac-ultramacas-
sociation below the YDM seems to be part of a much larger magma
plumbing system, which supplied magma to the main Dovyren cham-
ber (Ariskin et al., 2009, 2015a).
4.1. Mac-ultramacsills
Similar to the YDM, the underlying sills dip almost vertically,
allowing for sampling of their complete cross-sections. Most sills
exhibit contrasting layering, which involves leucocratic Ol-bearing
gabbronorite, olivine gabbronorite, plagiodunite, and altered peridotite.
The gabbroic units appear either as concordant layers up to tens of
meters thick, or as cutting dykes (Fig. 10). The largest sampled Camel
Sill is located under the central YDM and is ~200 m thick. The lower-
most ~55 m of this body is made up of melanocratic olivine-rich
gabbronorite, which is chemically similar to the olivine gabbronorite
and plagioperidotite from the bottom part of the YDM (see Appendix
B, Sheet Rocks;Ariskin et al., 2013b, 2015a). The middle part (~50 m
thick) consists of Ol-bearing mesocratic to Ol-free leucocratic
gabbronorite, which gives way to overlying plagiodunite (~40 m) and
extensively serpentinized peridotite (50 m). The occurrence of the
most high-Mg ultramacrocksabovemac units within this sill, com-
bined with a limited exposure, leaves open the possibility that the
Camel Sill represents a combination of separate magma pulses with
variable amounts of transported olivine.
4.2. Volcanic sequences
Volcanic rocks of the Synnyr-Dovyren complex include high-Ti ba-
salts of the Inyaptuk suite and overlying low-Ti sequences composed
252 A. Ariskin et al. / Lithos 302303 (2018) 242262
of sills and dykes of gabbrodiabase, siliceous tuffs and basalts, andbasal-
tic andesite (the so-called Synnyr suite described in Kislov, 1998;
Fig. 11). The contact between high-Ti and low-Ti suites is demarcated
by a horizon of coarse-clastic tectonic breccia (Fig. 11C), whichprobably
formed during tectonic events responsible for the dramatic overturn of
the overall volcanic-plutonic sequence. The total thickness of the high-
Ti suite is ~250 m (Fig. 2, traverse IVa); however, NE of the intrusion
a much thicker sequence of high-Ti massive basaltic ows has been
described (Manuilova and Zarubin, 1981;Fig. 11D).
4.2.1. High-Ti basalts and quartz-feldspar porphyry
The high-Ti volcanics of the Inyaptuk suite include subaphyric
basalts and, probably, more primitive Cpx-porphyritic basalts (Ariskin
et al., 2015a). Due to extensive alteration, the original mineralogy of
the subaphyric basalts is inferred from the composition of replacing
minerals (chlorite after clinopyroxene) and crystal morphology. In
addition, the subaphyric texture involves relics of rare plagioclase
phenocrysts (b5%). Their groundmass is dominated by small laths of
plagioclase and abundant ilmenite (~10%), reecting the plagioclase-
saturated nature of these basalts. The whole-rock composition of the
subaphyric basalts is less magnesian compared to the Cpx-porphyritic
basalts (Appendix B, Sheet Rocks). The porphyritic texture of the latter
is manifested by abundant (~30%) large (up to 45 mm) grains of mostly
unaltered clinopyroxene. Rare chlorite pseudomorphs after likely
Ol-phenocrysts are also present and locally form intergrowths with
clinopyroxene. Some of the porphyritic rocks also contain abundant
Fig. 10. An ultramac sill in thermally-metamorphosed sediments below the bottom of the central YDM (Magnetitovyi Creek). Total thickness is ~80 m. Note the occurrence of a more
leucocratic gabbronorite dyke, cutting across the sill.
Fig. 11. Volcanic rocks of the Synnyr-Dovyren volcano-plutonic complex. A The contactbetween a rhyolite dyke and low-Ti basaltsof the Inyaptuk suitein the middle stream of Morenny
Creek (seetraverse IVa in Fig. 2);B The uppermostsill of a gabbro-diabase underlyinga sequence of low-Tibasaltic ows, whichcompose the top of Mt. Soldatat Synnyr Ridge; C A block of
a coarse-clastic tectonic breccia from the tectonicboundary between the uppermost high-Ti basaltsand overlying volcanic-sedimentary sequence (the Synnyr suite); D A sequenceof the
high-Ti basalts that compose the widespread basaltic covers NE of the Dovyren area.
253A. Ariskin et al. / Lithos 302303 (2018) 242262
sericite pseudomorphs after plagioclase phenocrysts, suggesting that
primary mineral assemblages ranged from Ol + Cpx to Cpx + Ol + Pl.
Another important feature of the Inyaptuk suite is the presence of
quartz-feldspar rhyolite dykes (up to 50 m thick), which cut sedimentary
rocks and basalt ows near the bottom of the volcanic succession
(Fig. 11A). These are leucocratic porphyritic rocks with a microspherulitic
felsic groundmass including accessory titanomagnetite. Relic pheno-
crysts of quartz and plagioclase make up 1015 vol% of the rock. Quartz
phenocrysts have mostly ovoid shape and are generally overgrown by a
microcrystalline quartz-feldspar aggregate. The rocks contain abundant
grains of zircon, which were used to date the felsic volcanics and the
timing of overprinting processes.
4.2.2. Low-Ti volcanics
The low-Ti Synnyr (sub)volcanic suite includes meso- to leucocratic
gabbrodiabase and gabbronorite forming several meters thick bodies
on the SW slopes of the Synnyr Ridge and basalt to basaltic andesite
ows both along the ridge crest and on the NE slopes of Synnyr Ridge
(Figs. 2 and 11B). The largest (~200 m thick) sill concordant with
hosting volcano-sedimentary rocks was sampled at the SW slope
of Mt. Soldat (summit 2232 m in Fig. 2;Ariskin et al., 2015a). It is com-
posed of ne-grained granophyric gabbronorite with a relict gabbro-
ophitic texture. Similar to other subvolcanic rocks, it consists of
completely saussuritized lath-like plagioclase and uralitized pyroxene.
The original orthopyroxene is replaced by secondary chlorite, whereas
clinopyroxene is replaced by actinolite. The interstices between the
relic plagioclase and pyroxene crystals contain subgraphic quartz-
albite intergrowths. Ilmenite plates are also common.The predominant
accessory minerals include Cu-Fe sulphides, titanite, and rutile.
Subaphyric and aphyric basaltic andesites were sampled among
low-Ti ows at the summit of Mt. Soldat. These are massive rocks
with a relic micro-intersertus texture dened by microlites of
saussuritized plagioclase and amphibolized pyroxene. There are also
nely dispersed sulphides, including both pentlandite and Ni-bearing
pyrrhotite, and rutile. The whole-rock compositions of the low-Ti
volcanics are generally similar to the chilled gabbronorite from the
roof and bottom of the YDM in the Schkolnyi section, and mac dykes
below the lower contact of the YDM in the Schkolnyi and Yoko areas
(compare the gabbrodiabase 07DV183-1 and the basaltic andesite
07DV192-1 with samples S25-1 to S25-6 or 07DV220-1 in Appendix B).
As discussed below, the low-Ti volcanics are likely related to the YDM
(Ariskin et al., 2015a).
5. Trace element geochemistry
Concentrations of incompatible elements in most intrusive rocks
and representative volcanics of the Synnyr-Dovyren Complex are
given in Appendix B (Sheet Rocks) and partly published by Ariskin
et al. (2015a). These data are summarised on mantle-normalised dia-
grams (Fig. 12). Overall, the compositions of the YDM rocks have
broadly similar mantle-normalised spectra, best seen in similar varia-
tions of REE, LILE, and HFSE, including a distinct Nb-Ta minimum and
a minor depletion in Zr and Ti, and high enrichment in Pb. A more de-
tailed comparison reveals some differences in whole rock geochemistry
between the Bolshoi-Tsentralnyi and Schkolnyi cross-sections.
5.1.1. The centre of the YDM
Geochemical spectra of the chilled rocks and plagioperidotite in
Fig. 12A represent samples from both Fo88and Fo86-trendstrends
(Fig. 8B). All the near-contact rocks are geochemically similar, with
only subtle variations of incompatible element concentrations due to
minor magma fractionation, consistent with the narrow compositional
range of melts in equilibrium with olivine Fo88 to Fo86. These rocks
are also similar to the ne-grained olivine gabbronorite near the lower
contact of the Camel Sill (sample DV35-2; Appendix B, Sheet Rocks).
The plagiodunite is more depleted in incompatible elements compared
to the YDM contact rocks. All patterns for dunite, troctolite, and olivine
gabbro are characterised by very low concentrations of incompatible el-
ements (Fig. 12B, C and D). The lowest concentrations of incompatible
elements are typical for the troctolite. The geochemical patterns are
Fig. 12.Mantle-normalised incompatible traceelement contentsof samples fromthe YDM and the underlying sills.Plots A to D show samples fromthe Bolshoi-Tsentralnyi cro ss-section: A
Chilled rocks, plagioperidotite and plagiodunite; B Dunite zone; C Troctolite; D Gabbroic rocks. Plots E to H show samples from the Shkolnyi cross-section: E Basal chilled rocks and
plagioperodotite; F Plagiodunite; G Tro ctolite; H Upper chil led rocks. The compos ition of an endocontact rock from the Camel Sill is shown in plots A and E for co mparison. All
compositions were normalised to a primitive mantle from Sun and McDonough (1989).
254 A. Ariskin et al. / Lithos 302303 (2018) 242262
consistent with petrographic observations indicating an inward de-
crease in the porosity of olivine and olivine-plagioclase cumulates,
which is a measure of the amount of the intercumulus melt.
5.1.2. The Schkolnyi section
Geochemical patterns of rocks from the NE termination of the YDM
resemble those in the centre of the intrusion. Maximum concentrations
of incompatible elements are typical for chilled gabbronorite and
plagioperidotite from the lower contact zone, and quartz-pigeonite
gabbro to chilled gabbronorite from the roof (Fig. 12E, H). The quartz-
pigeonite gabbro has higher concentrations of REE and slightly
lower abundances of LILE and Sr compared to the chilled rocks
from the bottom. Most samples of melanotroctolite and many olivine
gabbro samples display the lowest incompatible element concentra-
tions (Fig. 12G, H). Unlike the Bolshoi-Tsentralnyi cross-section,
plagioperidotite from the Schkolnyi Section is readily distinguished
from the chilled rocks due to depletion in incompatible elements
(Fig. 12E). Another difference from the centre is the presence of positive
anomalies of K and Sr in chilled rocks and plagiodunite (compare
Figs. 12 A, B and E, F). Overall, the entire succession of mac to ultra-
mac rocks from the Schkolnyi Section displays higher concentrations
of incompatible elements than in the central YDM. This is additional
evidence for a more-evolved composition of the parental magma in-
truded in the NE margin of the YDM chamber.
5.1.3. Low-Ti and high-Ti volcanics
Ariskin et al. (2015a) have demonstrated that the geochemical
spectrum of the low-Ti basalts of the Synnyr suite is similar to that
of the YDM chilled rocks, gabbro and gabbronorite. The data shown
in Fig. 13A support the hypothesis that Synnyr volcanics are likely
comagmatic with YDM magmas, particularly the Schkolnyi upper con-
tact rocks, even though the Nb-Ta minimum in the Synnyr gabbro-
diabase and basaltic andesite is not as pronounced as in the YDM con-
tact facies. Compared to the low-Ti volcanics, the high-Ti basalts of the
Inyaptuk suite display different mantle-normalised patterns, which
lack both Nb-Ta and Sr minima (Fig. 13B). This is consistent with the
Sr-Nd isotopic differences shown in Fig. 4.
6. Discussion
In this section, we focus on modelled characteristics of the Dovyren
parental magmas and possible formation mechanisms of the layered
YDM structure. The cumulate succession is shown to be related to the
history of sulphide saturation in the Dovyren magma chamber. Finally,
a probable mantle source of the Dovyren magmas is discussed.
6.1. Parental magmas of the Dovyren Intrusive Complex
Ariskin et al. (2003) presented probable parameters of the Dovyren
magma based on COMAGMAT modelling of equilibrium crystallisation
for 10 ultramac compositions, which represent the bottom of the in-
trusion and the underlying sills. As a result, a sub-eutectic parental
magma (Ol + Pl ± pyroxene) was proposed, which had a temperature
of 11801190 °C and contained ~40% of transported olivine. The
modelled melt contained ~54 wt% SiO
and ~7.5 wt% MgO, being in
equilibrium with olivine Fo84.6 and plagioclase An80.5 (Table 1). How-
ever, Bolikhovskaya et al. (2007) pointed out that the dunite zone of the
YDM does not contain any record of cotectic Ol-Pl crystallisation.
Instead, intercumulus plagioclase in dunite is much less anorthitic
than that in the overlying troctolite (Fig. 9). Further studies of the
lower contact zone indicated that most contact rocks should be
interpreted as originally Ol-rich cumulates, crystallised from a magne-
sian parental magma in equilibrium with olivine ~Fo87 (Ariskin et al.,
2009). In this case, the COMAGMAT calculations predict that upon em-
placement the original magmatic melt would contain 910% wt% MgO
at ~12401270 °C.
Based on the FeO vs. MgO relationships for the basal YDM ultra-
macs (Fig. 8B), it is possible to calculate probable parental magma
compositions more accurately (Ariskin et al., 2016). As mentioned
above, there are two distinct petrochemical trends: the Fo88-trend
includes all samples of chilled gabbronorite to picrodolerite and a con-
tinuum of plagioperidotite to plagiodunite, whereas the Fo86-trend is
formed by the rest of the plagioperidotite and plagiodunite composi-
tions (Fig. 8B). Since both groups of rocks occur in the same cumulate
succession, we interpret these observations as evidence for the early
stages of YDM formation, including almost coeval emplacement of a
compositional range of olivine-laden magmas carrying variable amounts
of intratelluric olivine Fo86 to Fo88.
To estimate an initial temperature of the most primitive magma
(Fo88 + melt), one can utilize the results of COMAGMAT-5 calculations
simulating equilibrium crystallisation of the primitive DV30-2
picrodolerite sampled ~1.4 m above the lower contact of the YDM
within the chilled zone (Ariskin et al., 2016;Table 1). This is because
the composition of this rock belongs to the Fo88-trend in Fig. 8B. As
follows from data in Table 1, the equilibrium state of a modelled mixture
of 37.4 wt% olivine Fo88 and 62.6 wt% magmatic melt is consistent
with the 1285.4 °C equilibrium temperature. The modelled initial melt
contains ~11 wt% MgO; this is very similar to the composition of a
chilled gabbronorite sampled at a 10 cm distance from the lower con-
tact (see CGNcolumn in Table 1).
The results of this modelling characterise compositions of two most
primitive phases of the Dovyren magma, whereas, the calculated value
37.4 wt% olivine is attributed to a primitive olivine orthocumulate con-
taining relatively large amounts of intercumulus melt. The contact cu-
mulate is assumed to crystallise at the same initial temperature as a
parental magma composed of the same magmatic melt and unknown
amount of olivine Fo88. Assuming that sorting and accumulation of
olivine suspended in the parent melt could affect the bulk composition
of cumulus mixtures during their solidication near the lower contact,
one can state that the amount of olivine in the parental magma could
not exceed 37%. This is an upper limit, which is consistent with the
bulk MgO composition that is below 24 wt%. At this stage, we cannot
provide accurate estimates of the modal and chemical composition of
the most primitive Dovyren magma. However, the occurrence of chilled
Fig. 13. Mantle-norma lised patterns of incompatible eleme nts in low-Ti volcanics
compared to characteristics of the roof gabbronorite from the YDM and high-Ti basalts.
The low-Ti compositions in both plots represent the Synnyr gabbro-diabase 07DV183-1
and basaltic andesite 07DV192-1, including comp arisons: A Uppermost gabbronorite
from the Schkolnyi Section; B High-Ti basalts 07DV323-1, 327-6, 327-7, and AA30 (see
compositions in Appendix B, Sheet Rocks). Normalis ation to a primitive mantle from
Sun and McDonough (1989).
255A. Ariskin et al. / Lithos 302303 (2018) 242262
rocks and rare dykes of 1720 wt% MgO (see picrodolerites DV30-1,
07DV100-1, and dyke T1-7 in Appendix B) reects a probable picritic
parental melt.
Further crystallisation of the DV30-2 cumulate system generates
more-evolved melts. As the temperature decreases to 1189.7 °C, the
equilibrium state corresponds to a cotectic mixture of 44.8 wt% olivine
Fo86, 2.1 wt% plagioclase An78.6, a small amount of Fe-Ni sulphide
liquid, in equilibrium with a melt containing ~7.6 wt% MgO. The melt
composition closely resembles that of the 07DV220-1 gabbro-diabase,
sampled below the lower YDM contact in the Yoko area (Table 1). The
calculated melt is also similar to the original YDM meltcalculated by
Ariskin et al. (2003), despite the fact that the authors utilized other
ultramac compositions and another magma crystallisation model
(COMAGMAT-3.5) instead of the COMAGMAT-5.2 used in this study.
We suggest that the modelled melt composition, given as the
1189.7 °Ccolumn in Table 1, may be considered to be a close approx-
imation of the magmatic parental melt consistent with olivine Fo86
in the more-evolved Ol-laden Dovyren magma (Fig. 8B). Thus,
this modelling resolves the apparent contradiction between the rst
COMAGMAT-3.5 calculations and the present cumulate structure of
the YDM centre (Bolikhovskaya et al., 2007). The earliest estimate
of the Dovyren magma (Ariskin et al., 2003) should be treated now
as recovering the second relatively-evolved intrusion-forming mag-
matic melt.
Overall, the results suggest a diversity of olivine-laden to multiply-
saturated magmas, which formed the YDM. The most primitive and
higher-temperature magma was saturated with olivine (+spinel),
whereas its lower temperature derivatives represent Ol-Pl cotectics.
Based on the structure of the YDM (Figs. 5 and 9), it is possible to sug-
gest that the central part of the intrusion was formed predominantly
by high temperature Fo88-magma, whereas its peripheral zones
and ultramac apophyses crystallised from the lower temperature
Fo86-magma. This inference is supported by the presence of a large
volume of dunite in the centre, which is absent in the Schkolnyi
and Yoko sections, and by the presence of less magnesian olivine com-
positions observed within the cross-sections in the marginal areas
(Figs. 5, 9 and A5/A6 in Appendix A). Some of the mac-to-ultramac
sills beneath the main intrusive body might be formed by even more
evolved Ol + Pl-saturated magmas; however, this hypothesis needs
further investigation.
6.2. Sulphide immiscibility in the proposed magmas
Understanding the sulphide saturation history of mac-to-ultramac
igneous systems includes responding to the question: Is the amount of
magmatic sulphur sufcient for reaching sulphide immiscibility at early
stages of magma evolution?Ripley and Li (2013) have considered how
sulphur capacity varies in open and closed magmatic systems evolving
by fractional crystallisation. They emphasised that in open magma con-
duits, where efcient sequestration of sulphide from a large magma
volume is possible, an external source of sulphur may not be required to
form magmatic sulphide mineralisation. Conversely, in closed magma
chambers, sizeable sulphide-rich magmatic mineralisation is unlikely
to form without assimilation of sulphur from the host rocks, because
20 to 40% magmacrystallisation is required before the onset of sulphide
immiscibility, with the consequent sequestration of nickel and other
metals of interest in the early forming silicate phases. However, they
also concluded that mantle-derived magmas crystallised as closed sys-
tems have the more likely potential to generate Cu and PGE sulphide-
poor reefs.
Sulphidemineralisationin the YDM and the compositions of the pro-
posed parental magmas provide insights into sulphide saturation pro-
cesses in mac systems. Comparing the modelled S concentrations
with estimates of sulphide solubility in melts from our area of interest
(Table 1), it is possible to quantify the onset of sulphide immiscibility
in the crystallising DV30-2 cumulate. At the initial temperature of
1285.4 °C the S concentration is 20% lower than SCSS, reecting
the sulphide-unsaturated character of the most primitive Dovyren
magma. At the temperature of 1189.7 °C the COMAGMAT-5 modelling
predicts 650 ppm S = SCSS in the melt. This is consistent with
sulphide-saturated conditions and theoccurrence of ~0.04 wt% sulphide
among crystallising olivine and plagioclase (Table 1). In fact, sulphide
Table 1
Primitive chills and parameters of parental magmas proposed for the Dovyren Intrusive Complex.
Melt components, wt% This study Ariskin et al. (2003),n=10
CGN DV30-2 DV30-2 melts 07DV220-1 1185 °C S.D.
1285.4 °C 1189.7 °C
53.20 47.51 52.14 54.27 53.14 54.13 0.80
0.65 0.39 0.63 0.74 0.61 0.78 0.10
14.15 8.79 14.12 15.34 15.52 15.24 0.42
–– 0.86 0.84 ––
FeO 9.38 9.83 8.16 6.62 7.98 8.19 0.57
MnO 0.22 0.17 0.16 0.14 0.15 0.15 0.02
MgO 10.64 24.42 10.94 7.62 8.12 7.51 0.33
CaO 9.31 6.34 10.00 11.05 10.99 11.33 0.87
O 1.28 1.08 1.74 1.97 1.94 1.65 0.47
O 0.79 0.62 1.00 1.17 1.26 0.88 0.25
0.11 0.08 0.13 0.16 0.12 0.14 0.09
0.10 0.57 ––0.08 ––
NiO 0.032 0.142 0.034 0.017 0.015 ––
S 0.13 0.05 0.080 0.065 0.07 ––
SCSS –– 0.099 0.065 ––
Modelled characteristics of the proposed magmas
Ol, wt% –– 37.4 44.8 ––
Pl, wt% –– – 2.1 ––
Sulphide, wt% –– – 0.044 ––
Fo, mole% –– 88 86 84.6 ±1.0
An, mole% –– – 78.6 80.5 ±4.5
CGN and DV30-2 are chilled gabbronorite (07DV100-3b-I) and picrodolerite sampled at the distance of 10 cm and 1.4 m from the direct contact with contact hornfels; 07DV220-1
represents a comagmatic gabbro-diabase dyke below the lower contact of the YDM in the Yoko area (Appendix B). Columns 1285.4 °Cand 1189.7 °Care melt compositions calculated
at 1285.4 °Cand 1189.7°C, as a result of modelling equilibrium crystallisation of DV30-2 using the COMAGMAT-5program (Ariskin et al., 2013a). The 1185°Ccomposition is an initial
magmatic melt calculated for 10 samples of plagioperidotite (mostly from sills) using COMAGMAT-3.5 (Ariskin et al., 2003). S.D. is standard deviation (1σ).
256 A. Ariskin et al. / Lithos 302303 (2018) 242262
immiscibility in the DV30-2 system started at T~124C, i.e. 3040 °C
below the magma emplacement temperature in equilibrium with
olivine ~Fo87 (Ariskin et al., 2016).
These results suggest that the olivine cumulate piles in the centre of
the YDM attained sulphide immiscibility during crystallisation of
the intercumulus melt at a post-cumulusstage. As a result, a Ni-rich im-
miscible sulphide liquid is assumed to have originated directly in the
cumulate pile, followed by its downwards migration to form a poorly-
mineralised plagiodunite horizon (Ariskin et al., 2016). This conclusion
is supported by eld observationsand chemical rock analyses, which in-
dicate that a number of basal cumulates are oversaturated with sul-
phides in the intercumulus melt, particularly within plagiodunite and
To explain such a local enrichment (which could not be generated in
situ due to the very low crystallisationproportion of sulphides), one
should suggest a downward transport of the immiscible sulphide liquid
and its concentration within the cumulate piles. On the contrary, less
primitive cumulates from the Schkolnyi and Yoko sections crystallised
from sulphide-saturated magmatic melts. The modelling-based conclu-
sions may explain the absence of essential sulphide accumulation in the
most primitive ultramac rocks from the central YDMand a widespread
distribution of Cu-Ni sulphide ores in the NE and SW terminations of the
intrusion. This is also consistent with the occurrence of patchy to net-
textured Cu-Ni sulphide ores in relatively evolved olivine gabbronorite,
exposed asdykes and sills alongand underneath the lower YDM contact
(Ariskin et al., 2016;Kislov, 1998).
In addition, the results of the sulphide immiscibility calculations are
consistent with the conclusion of Ripley and Li (2013) that mantle-
derived magmas have the potential to generate Cu-sulphide-rich PGE
reefs in closed magma systems. In fact, even small portions of the
intercumulus sulphide liquid can extract most PGE and other highly-
chalcophile elements from the melt (Mungall and Brenan, 2014). We
suggest that in the troctolite and gabbroic parts of the Dovyren chamber,
sulphide immiscibility occurred at the temperatures around 1200 °C,
generating Cu-rich sulphides, which formed the sulphide-poor PGE-
rich troctolite and PGE-rich anorthosite (Figs. 5 and 7). However, we
argue that PGE-rich sulphide-poor mineralisation in the YDM troctolite
and anorthosite formed in the Dovyren magma chamber as it evolved
as an open system, at least in the initial tomiddle stages of its evolution.
6.3. Fingerprints of an open magma chamber
Following from geochemical data on chilled rocks from the lower
contact, the modelled porosity of olivine ± plagioclase cumulates, and
a narrow range of mineral compositions throughout the YDM, Ariskin
et al. (2003) have concluded that there was no large-scale magma frac-
tionation in the Dovyren chamber. Instead, they suggested that the
main formation mechanism of the contrasting YDM structure was the
compaction of the original olivine-rich cumulate piles, accompanied by
upward migration of the intercumulus melt through the porous space
of partly-crystallised cumulates. However, it is unclear whether these
melts accumulate and crystallise in the upper part of the magma cham-
ber, or whether they migrate out of the heterogeneous magmatic reser-
voir, leaving behind a crystalline residue of low-porosity ultramacand
gabbroic cumulates. In the rst case, one can argue for a closed magma
chamber, with the average intrusion composition consistent with that
of the basal picrodolerite/plagioperidotite. The second scenario suggests
the depleted character of the average weighted composition of the
Dovyren intrusion compared to that of the contact rocks as a proxy for
the composition of probable parental melts. Such misbalance should
rst be reected in the concentrations of components partitioning into
the melt, including incompatible major and trace elements.
6.3.1. Misbalance of incompatible components
Fig. 14 presents variations in SiO
, TiO
, and four trace elements in
the Dovyren rocks along cross-sections in the central part of the
intrusion and in the Schkolnyi area. Zr, Y, and Nb were selected as in-
compatible elements; they are relatively immobile at the secondary
alteration. Rb is chosen to assess the behavior of potentially mobile
elements. Rocks from both sampled cross-sections display a C-shaped
chemostratigraphy, with elevated whole rock concentrations being
towards the lower and upper contacts, and a marked depletion within
the inner parts of the intrusion. The maximum incompatible element
concentrations of the lower contact picrodolerite are very similar
to those observed in olivine gabbronorite from ultramac sills and associ-
ated dykes. This supports previous interpretations that the contact YDM
rocks approximate the parental magma composition (Ariskin et al.,
2009). Overall, these geochemical patterns illustrate the ultra-depleted
character of the intrusion, because the dunite, troctolite, and most of
the gabbroic rocks are 35 fold depleted in incompatible elements as
compared to the contact facies, particularly chilled gabbronorite.
The calculated weighted average compositions of the YDMare listed
in Table 2. Columns 13 represent published estimates for the major
rock-forming elements (Bolikhovskaya et al., 2007;Konnikov, 1986;
Yaroshevskii et al., 1982). The values in columns 4 (Aver-1)and
7(Aver-2) for the Bolshoi-Tsentralnyi and Schkolnyi sections, respec-
tively, are based on major element and trace element analytical data
summarised in this study (Appendix B, Sheet Rocks). Thus, our esti-
mates from this work are consistent with the ones from previous
studies, indicating a high-MgO YDM composition: ~2729 wt% MgO
and around 44 wt% SiO
. Due to somewhat elevated Al
(~10 wt%)
at present MgO, it could be treated as a troctolite-likecomposition
(Ariskin et al., 2009). Comparison of data given in columns 4 and 7 evi-
dences a slightly more magnesian composition at the centre of the
intrusion than that at the NE termination. This is consistent with the
absence of adcumulate dunite in the Schkolnyi section (Fig. A5 in
Appendix A) and the above conclusion that the marginal parts of the in-
trusion were formed during crystallisation of a lower temperature,
more evolved magma (Table 1). This inference is also supported by
higher concentrations of TiO
O, P
, and other incompatible ele-
ments in the Schkolnyi Section (Fig. 14B; Appendix B, Sheet Rocks).
In addition, Table 2 includes two compositions denoted as Proxy-1
and Proxy-2in columns 4 and 8, respectively. These compositions are
considered to be a close approximation of olivine-laden Dovyren
magmas, assuming that those should have MgO and mg# similar
to the weighted average YDM compositions (following from a
closed magma chamberscenario). Proxy-1 represents the 07DV132-3
plagioperidotite from the bottom zone of the YDM, which belongs to
the Fo88-trend in Fig. 8B. Proxy-2 corresponds to a mixture of 68.8%
olivine gabbronorite T1-7 (from a dyke sampled 50 m underneath the
lower contact of the YDM in the Schkolnyi area) and 31.2% olivine
Fo86. Both sample compositions are listed in Appendix B (Sheet
Rocks). Columns 6 and 9, denoted as Av/Pr, represent ratios of the
Aver-1and Aver-2concentrations to those of the two proxies above.
These calculations indicate that both average YDM compositions are
25 fold depleted with respect to the proposed Dovyren magmas,
with the cross-section in the centre being more depleted than that in
the Schkolnyi area. The depleted character of the YDM suggests that at
least ~50 to 70% of intercumulus gabbronorite melts were expelled
from the magma chamber during chamber formation and solidication.
6.3.2. Signicance of the Al
-MgO diagram
Additional evidence for the open magma chamber is given in Fig. 15,
which presents a projection of the YDM whole-rock compositions onto
the Al
-MgO plane. There are two well-dened trends in theter-
nary diagram, which yield insight into the origin of the Dovyren rocks.
The rst trend represents nearly concurrent tie-lines between observed
olivine compositions (Fo84-88) and a chilled YDM facies, displaying a
linear sequence of picrodolerite, plagioperidotite, plagiodunite, and
the lowest dunite. Figurative points of these rocks are located in the
order of decreasing olivine cumulate porosity, so that more magnesian
rocks contain less intercumulus. One can suggest that the maximum
257A. Ariskin et al. / Lithos 302303 (2018) 242262
porosity is recorded in the composition of the chilled gabbronorite,
which has the lowest MgO due to a minimum amount of suspended
olivine (Fig. 15). Such linear compositional relationships are typical
only for a basal zone ~250 m thick (Figs. 5 and 9).
Another distinct trend includes a diversity of the Dovyren troctolite
(Fig. 15), with whole-rock compositions following a tie-line connecting
those of pure olivine and plagioclase (An84-87). This observation
argues that the Dovyren troctolite should be considered as a binary
adcumulate mixture of olivine and plagioclase (plus traces of spinel),
which is almost free of Px-containing intercumulus material. The
same conclusion is true for the adcumulate dunite, because the compo-
sitions of rocks from the dunite zone fall close to pure olivine in the
ternary diagram (Fig. 15). These distinct trends support the inference
that the original characteristics of the olivine-laden magmas are re-
corded in the variable ultramac compositions from the basal zone,
whereas the adcumulate nature of dunite and troctolite reects a
large-scale expulsion of intercumulus melt from the magma chamber.
Whole-rock compositions, which fall in between two major trends,
represent cumulate gabbroic and anorthositic rocks from the upper
Assuming the open character of the Dovyren magma chamber, it
is possible to explain the unusual troctolite-likeweighted average
compositions of the YDM rocks (Table 2). We have demonstrated that
the YDM parental magmas were saturated in olivine + spinel and
carried a variable but substantial amount of olivine crystals. However,
the average YDM composition (Table 2) plots away from the initial
olivine melt tie-line, which connects the original olivine composition
(~Fo88) and probable magmatic melt (Fig. 15). This shift towards
the olivine-plagioclase tie-line (i.e., a lower SiO
and higher Al
compared to the olivine melt trend) is direct evidence for the loss of
asignicant amount of primitive and residual melts from the Dovyren
magma chamber, giving rise to the increased proportion of olivine and
plagioclase in the average troctolite-likeYDM composition.
6.4. Formation and evolution of the Dovyren magma chamber
In addition to the compaction hypothesisdiscussed above, it is pos-
sible to speculate whether the observed large-scale depletion in incom-
patible elements may be due to a ma gma-staging system that evolved in
the upper part of the Dovyren magma chamber, where large amounts of
olivine gabbronorite magmas passed through leaving behind cumulate
piles as a melt-depleted precursor of Ol-rich to gabbroic adcumulates.
The proposed mechanism is similar to the one that explains the origin
of olivine-plagioclase adcumulus aggregates (the so-called allivalites)
found as xenoliths in volcanic systems of the Kamchatka Peninsula
in Russia (Plechov et al., 2008). Both scenarios are important for the de-
velopment of a petrologic-geological model of the formation and evolu-
tion of the Dovyren magma chamber. Below, we propose such hybrid
mechanism, including four stages whereby magma emplacement and
early crystal sedimentation proceeded simultaneously, followed by the
formation of adcumulates during compaction and in situ crystallisation
of the original cumulate piles.
Fig. 14. Chemostratig raphy of SiO
, and selected trace e lements in the YDM. Cross-sections of the YD M: A Bolshoi-Tsentralnyi; B Schkolnyi. Th e green lines represent two
approximations of olivine-rich parental magmas from Table 2: A Proxy-1, B Proxy-2.
258 A. Ariskin et al. / Lithos 302303 (2018) 242262
6.4.1. The rst stage
The formation of the Dovyren chamber may be considered as a result
of numerous pulses of olivine-laden geochemically similar picritic
magmas (1720 wt% MgO), spanning a temperature range of 100 °C,
approximately from 1290 °C (Fo88) to 1190 °C (Fo86). The magma em-
placement process was fast enough to escape complete solidication of
each pulse, thus accommodating the continued growth of the magma
chamber. Large-scale sedimentation of the olivine crystals started si-
multaneously with the lling and growth of the magma chamber, thus
giving rise to an original Ol-rich cumulate pile and an overlying layer
of crystal-depleted gabbronoritic magma. The uppermost macpartof
the growing Dovyren chamber was not stagnant; it evolved as an
open magma owing system. Occurrence of autoliths of plagiolherzolite
in the upper part of the YDM (D.A. Orsoev, pers. comm) may be consid-
ered as a record of the rst stage process.
During chamber formation, assimilation of the country rocks could
take place; however,we still cannot evaluate this process quantitatively.
Preliminary estimates from isotope studies suggest that this interaction
had minor effect on the bulk magma composition (Ariskin et al., 2015a).
It is also unclear whether carbonates acted as the country rock that
hosted the original (much smaller) magma chamber. It is only
possible to suggest that the carbonate horizons collapsed and partly
dissolved during the later stages of the magma chamber formation
(Wenzel et al., 2002), accommodating the continuing growth of the
magma chamber accompanied by the proposed large-scale differentia-
tion. The undisturbed structure of the marginal part of the YDM in the
Schkolnyi area, where carbonate xenoliths are absent, supports this
inference (Fig. 14B).
6.4.2. The second stage
At the time when the magma chamber attained its nal size and
geometry, it had already developed a layered heterogeneous structure,
with most olivine crystals accumulated in its lower ultramacpart,
promoting the formation of a relatively crystal-depleted non-stagnant
upper gabbroicpart. After a critical mass of olivine cumulates formed,
compaction of the cumulate pile started, accompanied by additional
adcumulus growth of olivine crystals and generation of in situ crystallised
poikilitic plagioclase with minor clinopyroxene. Filter-pressing of the
Table 2
Average weighted compositions and proposed proxiesof the most primitive olivine orthocumulates of the Yoko-Dovyren massif.
Oxide, wt% Previous studies Bolshoi-Tsentralnyi section, n = 141 Schkolnyi section, n = 89
n = 256 n = 114 Aver-1 Proxy-1 Av/Pr Aver-2 Proxy-2 Av/Pr
44.54 45.00 43.92 43.87 45.82 0.957 44.18 46.41 0.952
0.09 0.21 0.11 0.10 0.29 0.354 0.16 0.32 0.503
10.64 10.27 9.72 9.74 7.31 1.334 10.12 7.66 1.321
FeO 10.05 10.95 10.53 9.94 10.36 0.959 10.31 11.09 0.930
MnO 0.14 0.17 0.16 0.17 0.932 0.17 0.18 0.927
MgO 26.57 24.31 27.88 28.50 29.23 0.975 26.98 26.94 1.001
CaO 7.35 6.87 6.99 6.66 5.06 1.314 6.73 5.88 1.145
O 0.54 0.78 0.59 0.59 0.48 1.217 0.78 0.54 1.430
O 0.07 0.28 0.07 0.12 0.52 0.225 0.24 0.82 0.294
0.01 0.02 0.02 0.06 0.292 0.03 0.06 0.512
––0.31 0.69 0.441 0.31 0.10 3.173
mg# 0.825 0.805 0.825 0.836 0.834 0.824 0.812
S, wt% ––0.055 ––0.085 ––
Trace elements, ppm
Zr 9.37 35.8 0.262 14.5 34.71 0.417
Y 2.32 7.69 0.302 4.20 7.98 0.527
Nb 0.485 2.04 0.237 0.863 2.37 0.365
Rb 2.24 17.5 0.128 6.45 33.4 0.193
Ba 50.3 228 0.221 123 217 0.567
Sr 95.7 85.6 1.118 140.8 77.1 1.828
Sc 11.9 21.8 0.546 18.0 18.6 0.966
V 42.9 106 0.404 81.9 123.3 0.664
La 1.23 7.42 0.166 2.86 5.83 0.490
Eu 0.104 0.424 0.245 0.255 0.385 0.664
Lu 0.027 0.121 0.222 0.069 0.147 0.466
The average YDM compositions are normalised to a volatile-free basis. Previous studies: 1 Yaroshevskii et al. (1982), calculated as the average of 256 rocks sampled through the entire
massif; 2 Konnikov (1986);3Bolikhovskaya et al. (2007), calculated for a combined cross-section 2950 m thick. Assumed original olivine cumulates (slightlyenriched in olivine with re-
spect to a probable olivine-laden (picritic) parentalmagma: Proxy-1 is plagioperidotite 07DV132-3 from the bottom zone of theYDM in the centre (the Fo88-group in Fig. 8B); Proxy-2
corresponds to a mixtureof 68.8% olivine gabbronorite T1-7 and 31.2% olivine Fo86.The whole-rock compositions aregiven in Appendix B (Sheet Rocks). Ratiosemphasised using black
colour in columns 6 and 9 demonstrate a relative enrichment of the average YDM in plagioclase-elements, such as Al, Ca, Na, and Sr.
Fig. 15. The Al
-MgO diagram for the whole-rock compositions from the central
YDM (the Bolshoi -Tsentralnyi se ction). Modelled melts repres ent two composit ions
calculated at 12 85.4 and 1189.7 °C (see Table 1 and Fig. 8B). The average YDM
composition is given in Table 2 as Aver-1.
259A. Ariskin et al. / Lithos 302303 (2018) 242262
residual melt out of the lower crystallisation zone was most likely a trig-
ger for the upward ow of the porous melt. The onset of this stage is re-
corded in the occurrence of plagiodunite, whose composition displays
much lower Ca/Al ratios as compared to the chilled picrodolerite. In ad-
dition, plagiodunite shows a subtle enrichment in Sr and Eu (Fig. 12),
which is expected for heterogeneous systems slightly enriched in
cumulus plagioclase. This is an important argument for the cotectic
nature of poikilitic plagioclase crystallised in situ, with the residual
melt leaving the zone of initial adcumulus growth.
The magma-staging behavior of the upper part of the chamber con-
tinued, as demonstrated by the presence of both dykes and sill-like bod-
ies of gabbronorite above the roof of the YDM. Despite some localised
upper reversal signatures (Latypov, 2015;seeVinFig. 5) in the upper-
most part of the YDM, the absence of a widespread roof succession
(similar to the Upper Marginal Series of the Skaergaard intrusion)
may be considered as indirect evidence for the high efciency of the
magma owing process. This is because under the force of the emplaced
magma, horizons of roof rocks crystallised downward collapsed. This in-
ference is supported by the occurrence of large blocks of host country
rocks and recrystallised hornfels inside the gabbronorite (Ariskin et al.,
2013b;Kislov, 1998).
6.4.3. The third stage
Due to continuing growth of the cumulate pile, the upper crystal ac-
cumulation front reached the highest levels of the magma chamber,
while its lower part was essentially solidied. This part of the Dovyren
history is most speculative and questionable. In fact, one can suggest
accumulation of the residual melts percolated from below at the
upper boundary of the compacting cumulate pile. Assuming coupled
effects of crystal compaction and continuing in situ crystallisation, the
mac melt residue had a subcotectic temperature (1180 °C or slightly
lower), consistent with olivine Fo84-85 and plagioclase (An83-86), as
observed in troctolites (Fig. 9). This slightly evolved melt could mix
with the owing magma. One cannot exclude that texturally inhomoge-
neous Cpx-containing troctolite in the upper part of the troctolite zone
could originate at an initial stage of such mixing. Additional argument
for the continuing melt expulsion from the upper part of the magma
chamber follows fromgeochemical features of the olivine gabbronorite,
whereby some of these rocks are more depleted in incompatible ele-
ments than the plagiodunite near the base of the YDM (Fig. 12D).
6.4.4. The fourth stage
The fourth stage characterises concluding magma transport phe-
nomena in a residual and smaller magma reservoir, still including
both cumulate piles and relatively liquid sub-chamber. Due to the fact
that the efciency of magma ow decreased, the chamber nally closed.
Continuing heat loss, both from below and above, induced fractionation
within the crystal-depleted melt, with most evolved rocks crystallising
near the roof of the chamber (Fig. 9). From the other side, upward per-
colation of the residual melts from crystallising adcumulates was
continuing at the same time, thus favouring mixing of the primitive
magma with upper fractionated products. It is interesting that in a
gabbronorite approximately 100 m below the upper contact, we found
two distinct clusters of olivine compositions ranging from Fo68 to
Fo79 (see data for the sample 09DV503-7 in Appendix B). This supports
the proposed magma mixing hypothesis, as such contrasting olivine
compositions are non-typical for most Dovyren rocks.
6.5. The mantle source
There is still uncertainty with respect to the nature of a probable
mantle source of the YDM parental magmas, which are geochemically
similar to the low-Ti basalts of the Synnyr suite, as well as being ex-
tremely enriched in radiogenic Sr and depleted in ε
.Amelin et al.
(1996) proposed a model which explains such anomalous characteris-
tics as a result of melting of a metasomatized mantle peridotite shortly
after subduction of sediments derived from an ancient (~2.42.8 Gyr)
upper continental crust into a depleted lherzolitic mantle. Ariskin
et al. (2015a) considered a different model, which is based on
reconstructing the time-dependent evolution of ε
(t) for the YDM
rocks until the inverse ε
-trend (consistent with Sm/Nd 0.221 in a
mantle protolith) intersects that of primitive mantle evolution with
the initial mantle ratio of Sm/Nd = 0.350 (Kostitsyn, 2004). This recon-
struction suggests that the anomalous mantle protolith formed at the
Meso-Neoarchean boundary at ~2.8 Gyr, remained isolated from mag-
matic events for ~2 Gyr, and then reactivated at 728 Ma.
Geochemical similarities between the YDM rocks and those of gran-
ulites and granitoids from the southern margin of the Siberian craton
suggest that the anomalous mantle could have formed within a subduc-
tion zone that existed at ~2.8 Gyr along themargin of the craton (Ariskin
et al., 2015a). In alternative, the crust-like Sm/Nd 0.221 ratio in
the proposed mantle protolith may argue for contamination by a prim-
itive mantle (komatiite-like?) magma of even more ancient crustal ma-
terials. This is consistent with the slightly shifted oxygen isotope
composition of olivine from the YDM (δ
O = 5.8 ± 0.1;Fomin
et al., 2013).
7. Conclusions
(1) The Yoko-Dovyren layered massif, associated mac-ultramac
sills, and gabbronoritic dykes represent the intrusive constituent
of the Upper Riphean Synnyr-Dovyren volcano-plutonic com-
plex. Comparison of complete cross-sections of the thickest
part of the largest intrusive body with those from this NE and
SW terminations of the magmatic body reveal key differences
in the lateral YDM architecture.The marginal domains are largely
dominated by melanotroctolite to troctolite with relatively thin
zones of plagiodunite, whereas the core mainly comprises a
thick dunite zone. These differences, highlighted with petrochem-
ical reconstructions, mineral chemistry, and COMAGMAT-5
calculations, indicate that the temperatures of the emplacing
olivine-laden parental magmas in the central and peripheral
parts of the intrusion ranged across 100 °C, approximately from
1290 °C (~11 wt% MgO, olivine Fo88) to 1190 °C (~8 wt% MgO,
olivine Fo86).
(2) Based on the present thermodynamic modelling, the high-MgO
magma in the centre was S-undersaturated, whereas its deriva-
tives became S-saturated at the temperature of 1240 °C or
below. This is consistent with geological observations that most
sulphide-rich Cu-Ni ores were discovered in the sills and apoph-
yses of plagioperidotite underneath the YDM, as well as at the
periphery of the intrusion, where ultramac rocks crystallised
from a relatively evolved olivine gabbronorite magma.
(3) Because of the S-undersaturated character of the Fo88-magma,
post-cumulus sulphide immiscibility could occur in crystallising
primitive cumulates. As a result, a Ni-rich immiscible sulphide
liquid is assumed to have originated directly in the cumulate
pile, followed by its downwards migration to form a poorly-
mineralised plagiodunite horizon. In the troctolite and gabbroic
parts of the Dovyren chamber, sulphide immiscibility could
occur at lower temperatures (probably around 1200 °C), gener-
ating Cu-rich sulphides, which gave rise to the sulphide-poor
PGE-rich troctolite and PGE-rich anorthosite.
(4) The C-shaped distribution of TiO
O, P
, and incompatible
trace elements along cross-sections of the YDM reects the de-
pleted geochemistry of most Dovyren rocks compared to rela-
tively thin contact zones, as well as associated ultramacsills
and gabbronorite dykes.Accounting for a marked misbalance be-
tween estimates of the average weighted compositionsof the in-
trusion and proxies of the parental magmas, these data suggest
that roughly 5070% of complementary basaltic melts had to be
expelled from the YDM during its consolidation.
260 A. Ariskin et al. / Lithos 302303 (2018) 242262
(5) Two possible scenarios for the evolution of an open magmatic
system are considered. One hypothesis takes into account that
the formation of a thick sequence of Ol-rich cumulate pile should
be accompanied by compaction and crystallisation, giving rise
to upward migration and inltration of the intercumulus melts.
The second hypothesis suggests that the Dovyren chamber is a
magma-staging system, through which large amounts of olivine
gabbronorite magmas have passed, leaving behind a complemen-
tary melt-depleted succession of dunite, troctolite, and gabbroic
cumulates. Here we propose a hybrid mechanism whereby these
processes proceeded simultaneously. Our current model suggests
that only minor fractionation of the Dovyren parental magmas oc-
curred at the early to middle stages of solidication. However, one
cannot exclude the possibility that at a nal stage of evolution the
magma chamber became closed, thus favouring in situ magma
fractionation within the upper portion of the residual heteroge-
neous reservoir. This is consistent with the most evolved mineral
compositions observed in olivine-free gabbronorite and quartz-
pigeonite gabbro from the uppermost YDM.
(6) Reconstructions of the time-dependent evolution of ε
YDM rocks suggest for the Dovyren magmas an anomalous mantle
protolith formed at the Meso-Neoarchean boundary at ~2.8 Gyr. It
remained isolated from magmatic events for ~2 Gyr, and then
reactivated at 728 Ma.
Supplementary data to this article can be found online at https://doi.
We gratefully acknowledge thoughtful reviews by Steve Barnes, Rais
Latypov and two anonymous reviewers, as well as constructive com-
ments from the editor Andrew Kerr. We acknowledge support from
AngloAmerican, BHP Billiton, Votorantim Metais, and the Australian
Research Council through funding to CODES (University of Tasmania,
Hobart, Australia) at the initial stage of this research (AMIRA project
P962, 20072010); support from the Russian Science Foundation (RSF,
grant No. 16-17-10129, 20162018); and support from the University
of Tasmania through Visiting Scholarships to AAA in 2011 and 2014.
MLF acknowledges support from the Australia n Research Council through
the Future Fellowship Scheme (FT110100241) and Foundation Project 2a
of the Centre of Excellence for Core to Crust Fluid Systems. We thank
Roland Maas (School of Earth Sciences, the University of Melbourne),
Sebastian Meffre, Sarah Gilbert, and Paul Olin (University of Tasmania)
for assistance with analytical work. Kirill Bychkov, Ian Woolword,
Ludmila Zhitova, Dima Kamenetsky, Alexey Lygin, Jonas Mota e Silva,
and Dmitry Orsoev assisted during eldwork at the YDM in 2007. We also
thank Evgeny Koptev-Dvornikov (Moscow State University, Russia) for
his help with description of thin-sections, Masha Anosova and Kostya
Ryazantsev (Vernadsky Institute, Moscow) for their assistance with sam-
ple preparation, and Kirill Bychkov for his work on the COMAGMAT-5
model. The authors would like to particularly acknowledge the con-
tributions of late Dr. Eduard Konnikov who worked on this project in
20072011. We thank Candace S. O'Connor for careful editing of the
submitted manuscript. This is contribution 1045 from the ARC Centre of
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