Metallogenic Setting and Evolution of the Pados ‐ Tundra Cr ‐ Bearing Ultramafic Complex, Kola Peninsula: Evidence from Sm–Nd and U–Pb Isotopes

: The article presents new Sm–Nd and U–Pb geochronological data on rocks of the poorly studied Pados ‐ Tundra Cr ‐ bearing complex. It is part of the Notozero mafic–ultramafic complex (western Kola Peninsula) and occurs at the border of the Paleoproterozoic Lapland Granulite Belt and the Archean Belomorian composite terrain. The Pados ‐ Tundra complex hosts two major zones, the Dunite and Orthopyroxenite Blocks. Dunites are associated with four levels of chromite mineralization. Isotope Sm–Nd studies of dunites, harzburgites, and orthopyroxenites from the central part of the complex have been carried out. The isochron Sm–Nd age on 11 whole ‐ rock samples from a rhythmically layered series of the complex is 2485 ± 38 Ma; the mineral Sm–Nd isochron for harzburgites shows the age of 2475 ± 38 Ma. It corresponds with the time of large ‐ scale rifting that originated in the Fennoscandian Shield. When the rhythmically layered series of the intrusion and its chromite mineralization were formed, hornblendite dykes intruded. The U–Pb and Sm–Nd research has estimated their age at ca. 2080 Ma, which is likely to correspond with the occurrence of the Lapland–Kola Ocean. According to isotope Sm–Nd dating on metamorphic minerals (rutile, amphibole), the age of postmetamorphic cooling of rocks in the complex to 650–600 ° С is 1872 ± 76 Ma. The U–Pb age on rutile from a hornblendite dyke (1804 ± 10 Ma) indicates further cooling to 450–400 °C. The conducted research has determined the early Proterozoic age of rocks in the rhythmically layered series in the Pados ‐ Tundra complex. It is close to the age of the Paleoproterozoic ore magmatic system in the Fennoscandian Shield that developed 2.53–2.40 Ga ago. Later episodes of alterations in rocks are directly related to main metamorphic episodes in the region at the turn of 1.9 Ga. Results of the current study expand the geography of the vast Paleoproterozoic East Scandinavian Large Igneous Province and can be applied for further studies of similar mafic–ultramafic complexes.


Regional Geology
The Fennoscandian Shield has the Archean basement that formed at ca. 3.5-3.2 Ga and consists of the Kola, Karelian, and Norrbotten cratonic cores, each with an associated cover of Paleoproterozoic sedimentary and volcanic rocks [48]. Together with the younger Svecofennian orogenic zone, Fennoscandia represents the largest crustal segment, spanning across the northern part of the East European Craton (Baltica). The Kola and Karelian cratons are considered fragments of either the late Archean supercontinent, sometimes referred to as Kenorland [49][50][51], or a yet unconstrained supercraton [52], which in either case, broke apart at 2.5-2.0 Ga, prior to the amalgamation of the next supercontinent, Columbia (Nuna), at ca. 1.9-1.7 Ga [53]. In the Paleoproterozoic (2.50-1.95 Ga), the Karelian craton experienced a number of intraplate magmatic events [54], some of which were likely parts of large igneous provinces (LIPs) [40]. Paleoproterozoic volcanic and sedimentary volcanic successions were deposited and are now preserved in the Onega structure, Perapohja schist belt, Salla-Kuolajarvi basin, and so forth. Terrigenous sequences containing black shales accumulated in the Onega and Salla-Kuolajarvi Basins between 2.06 and 1.95 Ga [54].
The Pados-Tundra complex occurs in the western part of the Kola Peninsula as a fragment of the Notozero ultramafic complex (Figures 1 and 2) [43][44][45]. The Notozero complex stretches 40 km transversally as a chain of the Khanlaut-Varaka, Pados-Tundra, Chapes-Varaka, Lotmvaara, and other intrusions. According to regional studies, this zone is the southeastern continuation of the Tanaelv Greenstone Belt [55]. It occurs at the border of the Lapland Granulite Belt (LGB) and the Archean Belomorian Belt, which is the northern part of the Karelian Craton ( Figure 2). The LGB is located between the Karelian Craton in the south and the Inari Craton in the north. The Tanaelv Belt [55,56] separates the LGB from the Belomorian complex. A number of geological and petrological studies [56][57][58][59] consider the Tanaelv Belt as a strongly deformed sequence underlying the LGB and tectonically overlying less deformed rocks of the Belomorian Complex in the Karelian Craton [60]. Abundant isotope data on rocks from the eastern Tanaelv suggest that it formed over 2.5-1.85 Ga [26,61]. Isotope data indicate that granulites in the LGB have the similar age of 1.9 Ga [62][63][64]. The rocks of the Belomorian Complex are significantly older (2.86-2.58 Ga) [65,66].

The Pados-Tundra Complex
Geotectonic settings of the Notozero complex and its occurrence at the boundary of the folded belt and the Archean geoblock can be correlated with settings of some chromite-bearing stratiform massifs (Campo Formose [67], Kemi [68,69], etc.). Massifs of the Notozero complex, Mt. Lyavaraka [70] and Falaley on the Kola Peninsula and in the North Karelian ultramafic belt, have been defined as early Proterozoic dunite-harzburgite formations [71].
The Pados-Tundra complex (Figures 2 and 3) takes the central position in the complex, being the largest in scale and best studied. In the 1990s, JSC "Murmansk Geological Prospecting Expedition" carried out a geological survey of the Pados-Tundra complex [43], searching for mineral resources. Geophysical, mining, and drill works (200 m deep boreholes) were carried out. No profound survey has been provided ever since.
At the present-day erosion level, the intrusion is represented by a body extending some 13 km 2 to the northeast ( Figure 3). Host rocks are composed of Archean granite and granite-diorite gneisses with the northeastern strike and the northwestern dip at the angles of 20-45°. The southern and western contacts of the complex have an intrusion pattern, the northern contact has the tectonic pattern, the eastern contact is not studied. In its exocontact, biotite-chlorite-amphibole schists and mylonites (1-5 m thick) are developed over granite and granodiorite gneisses. The complex was subject to the Karelian tectonogenesis, which is associated with the obduction of the LGB over the Belomorian Geoblock, processes of folding and fracturing, and acid magmatism. As a result, the initial form and geological structure of the complex experienced significant alterations [43], that is, turn-up of the flanks, superposed tectonometamorphism, replacement of rocks, and substitution of gentle for steep dip and strike patterns of rock layers. The complex contains three major structural units: an endocontact zone, a rhythmically layered series, and an upper zone (Figures 3 and 4). The 10-20 m thick endocontact zone is composed of amphibolite resulting from the metamorphism of mafic rocks. The rhythmically layered series hosts a number of rocks from dunites to orthopyroxenites and constitutes most of the complex. In total, there are seven rhythms, each of them starting with dunites and ending with orthopyroxenites [43,47]. Each rhythm starts with dunites and is associated levels of syngenetic chromite mineralization ( Figure 4). Rocks of this series contain dykes of meso-and leucocratic gabbro, diorites, and hornblendites (up to 1.5 m thick) that are bedded conformably with their host rocks. However, these dykes are not present in every rhythm, therefore, their syngentic nature is doubted. Gabbro and diorites occur in boreholes in the northwestern part of the complex only. The upper gabbronoritic zone is partly traced in the northeast of the complex; its estimated thickness is ~400 m. The main volume of this zone seems to be overlapped by host rocks as a result of obduction. Intensely amphibolized gabbroiс rocks were found in this upper zone during field works. These rocks seem to be relics of gabbronorites at the end of the Pados-Tundra complex section. The complex is marked by granite dykes [40,43]. They intersect the whole complex from the northeast to the southwest. Intrusion of these granites is linked to the development of tectonic faults of the northwestern strike in the complex. The genesis of these granites is associated with the acid volcanism activity at the turn of 1.9 Ga; however, these granites are not dated [43].
Chromitites occur in the seventh rhythm only as irregular lenses and stock-like bodies. The gabbronorite zone is partly observed within the Maliy Pados ( Figure 3). The setting of the satellite Maliy Pados complex is a disputable issue. Some researchers suggest that it formed at the same time as the main Pados-Tundra body [43]. Others believe that it disintegrated later from a minor block of the Pados-Tundra [72][73][74].

Chromite Ore Bodies
The complex is characterized by a high content of Cr (up to 1.10%), which is concentrated mainly in chrome spinellids. The complex hosts four horizons of bedded disseminated chromite ores and a series of lens-like and columnar bodies of chromite ores [43,[72][73][74]. All of the found chromite occurrences are associated with dunites or serpentinites after dunites ( Figure 4). The first chromite horizon Сr1 with the thickness of 0.04-0.5 m is traced in fragments at the bottom of serpentinites after dunites of the first rhythm. The second, third, and fourth horizons, as well as lenses and columnar bodies, are traced within the northeastern dunite block of the seventh rhythm (Figures 3 and 4).
The second horizon, Сr2, occurs at the footwall of the studied block. At the surface, the contact of this horizon with underlying orthopyroxenites is low and complicated by folded and block forms.
Outcrops of this horizon are assumed to be 1400 m long; there are numerous bodies with the length of 0.1-0.6 m, densely disseminated and with massive chromite mineralization.
The third horizon, Сr3, occurs in the central part of this block and is traced for 320 m by stripping. The horizon hosts 1-5 m long and 0.13-0.45 m thick lens bodies jointed by chromite veinlets for 140 m in the northeastern part of the block. Further south for 180 m, the horizon is divided into separate blocks displaced from 1-2 to 15 m. The chromite horizon has clear contacts with host dunites.
The fourth horizon, Сr4, occurs in the hanging wall of the dunite block and strikes for 85 m. This horizon also has a heterogeneous structure with 0.2-1.75 m thick swells jointed by mineralized veinlets. The contacts of the horizon with host dunites are sharp [43,72,74].
It is assumed that horizons Cr2-Cr4 within the dunite block were altered at the turn of 1.9 Ga as a result of tectono-thermic episodes associated with the obduction of Lapland granulites. As a result, chromite layers broke up, and lenses and columnar bodies occurred. Notably, there are still no geochronological data on these episodes.
According to the density of the chrome-spinellid dissemination, poor and relatively rich ores are identified in the chromite mineralization of the bedded morphological type (Сr2, Сr3, Сr4). Poor ores with the spinellid content of 30-40% have an irregularly grained and disseminated texture with certain densely disseminated areas. Their structure is sideronitic and xenomorphic-granular. Rich ores contain 60-85% of chrome spinellids; they have a massive densely disseminated texture and a hypidiomorphic-granular and idiomorphic-granular structure. The rich ore hosts minor amounts of sulfides represented by fine pentlandite and chalcopyrite phenocrysts.
Discrete steeply dipping lenses and columnar magnetite-chromite bodies are exposed in the western part of the dunite block in the continuation of the third and fourth horizons of disseminated chromite ores. In total, there are 120 such exposed bodies, the area of their cross-section ranges from 33 сm 2 to 1.4 m 2 , and the bodies are traced to the depth of 1 m in sections. Magnetite-chromite bodies are represented by massive densely disseminated and drained ores. It is irregularly grained; larger (0.2-0.8 mm) isometric and rounded crystals of chrome spinellids are surrounded by clustered crystals of chrome magnetites. Magnetite-chromite bodies occur among serpentinized and schistose dunites saturated in fine veinlets of talcite, magnesite, magnetite, and quartz.
Both rich and poor chromites in disseminated ore beds have a heterogeneous composition manifested in their zoned structure. In reflected light, central parts of chrome spinellid grains differ from marginal areas that produce fine rims with a worse reflecting capacity. In transmitted light, most of them are brownish red in contrast to black marginal areas [43].

Samples
Large (up to 20 kg) samples of dunites, harzburgites, and orthopyroxenites and one sample from a hornblendite dyke were selected from a rhythmically layered series of the complex (dunite block, fifth and sixth rhythms mainly). A harzburgite sample (P-1/1) was separately selected from the third rhythm for further extraction of mineral fractions. Sampling points were selected based on several factors, mainly on the clear geological occurrence of a sample regarding the structure of the complex, the absence of visible alterations of the rock as a result of superposed processes, sampling representativeness of the location, easy access to the site, and safe sampling. Gabbronorite was sampled in the upper zone of the Maliy Pados complex for the U-Pb studies. Sampling points are shown in Figure 3.
Two generations of Zr ( Figure 5) and one rutile variety were extracted from a hornblendite sample for U-Pb research. The first generation of Zr was represented by elongated-prismatic light brown crystals and their debris. The crystals were transparent with a glassy luster and their surface was corroded. The average size was 0.21 × 0.105 mm; the elongation coefficient (Cel) was 2. The weight value of an average crystal was 9.3 × 10 −6 g. A distinct zonation and transverse cracks were observed in immersion. Blurred and dark areas occurred along zonal boundaries and cracks. The second Zr variety was represented by grayish yellow debris. The grains were transparent with a glassy luster and their surface was corroded. The average size was 0.175 × 0.14 mm; Cel was 1.3. A distinct zonation and various cracks rimmed by dark areas were observed in immersion. Rutile was represented by a dark red variety. The grains were nontransparent with a glassy luster and their surface was corroded. The average size was 0.175 × 0.14 mm; Cel was 1.3.

Sm-Nd Method
The isotope research was carried out in the Collective Use Center of the Kola Science Center RAS (Apatity, Russia). First, the samples were prepared by crushing, then minerals were separated using heavy liquids and mineral fractions were selected under binocular microscope.
In order to define concentrations of Sm and Nd, the sample was mixed with a compound tracer 149 Sm/ 150 Nd prior to dissolution. It was then diluted with a mixture of HF + HNO3 (or + HClO4) in Teflon sample bottles at a temperature of 100 °C until complete dissolution. Further extraction of Sm and Nd was carried out using standard procedures with two-stage ion-exchange and extractionchromatographic separation using ion-exchange tar "Dowex" 50 x 8 in chromatographic columns employing 2.3 N and 4.5 N HCl as an eluent. The separated Sm and Nd fractions were transferred into a nitrate form, whereupon the samples (preparations) were ready for mass-spectrometric analysis.
The isotope Nd composition and Sm and Nd contents were measured with a 7-channel solidphase mass-spectrometer Finnigan-MAT 262 (RPQ) in a static double-band mode, using Ta + Re filaments. A mean value of 143 Nd/ 144 Nd ratio in a JNdi-1 standard was 0.512081 ± 13 (N = 11) in the test period. An error in 147 Sm/ 144 Nd in ratios was 0.3% (2σ), which is a mean value for 7 measurements in a BСR-2 standard [75]. An error in estimation of isotope Nd composition in an individual analysis was up to 0.01% for minerals with low Sm and Nd contents. The blank intralaboratory contamination was 0.3 ng in Nd and 0.06 ng in Sm. The accuracy of estimation of Sm and Nd contents was ±0.5%.

U-Pb Method
The method proposed by Krogh [80] was used to dissolve samples in strong (48%) hydrofluoric acid at the temperature of 205-210 °C over 1-10 days. In order to dissolve fluorides, the samples were reacted with 3.1 N HCl at the temperature of 130 °C for 8-10 h. To determine the isotope composition of Pb and concentrations of Pb and U, a sample was divided into two aliquots in 3.1 N HCl, then a mixed 208 Pb/ 235 U tracer was added. Pb and U were separated on an AG 1×8, 200-400 mesh anion exchanger in Teflon columns. A laboratory blank for the whole analysis was <0.1-0.08 ng for Pb and 0.01-0.04 ng for U. All isotope determinations were made on Finnigan MAT-262 mass spectrometers (Geological Institute KSC RAS, Apatity, Russia). The Pb isotope composition was analyzed on a secondary-ion multiplier on a Finnigan MAT-262 in an ion-counting mode. Measurements of the Pb isotope composition are accurate to 0.025% when calibrated against NBS SRM-981 and SRM-982 standards. U and Pb concentrations were measured in a single filament mode with H3PO4 and silica gel added. The method described in [81,82] was used. Pb and U concentrations were measured in temperature ranges of 1350-1450 °C and 1450-1550 °C, respectively. Isotope ratios were corrected for mass discrimination during static processing of replicate analyses of the SRM-981 and SRM-982 standards (0.12 ± 0.04% per a.m.u. for the Finnigan MAT-262). Errors in the U-Pb ratios were calculated during the statistical treatment of replicate analyses of the IGFM-87 standard. They were assumed equal to 0.5% for Finnigan MAT-262. Isochrons and sample points were calculated using the PBDAT (version 1.22, United States Geological Survey, Reston, VA, USA) and Isoplot programs (Version 2.05, Berkeley Geochronology Center, Berkeley, CA, USA) [83,84]. Age values were calculated with the conventional decay constants for U [85]. All errors were reported for a 2-sigma level. Corrections for common Pb were made according to [86]. Corrections were also made for the composition of Pb separated from syngenetic plagioclase or microcline, if the admixture of common Pb was >10% of the overall Pb concentration and the 206 Pb/ 204 Pb ratios were <1000.

Dating of the Complex
An Sm-Nd isochron was constructed based on 11 whole-rock samples. It complies with the age of 2485 ± 77 Ma, εNd(T) = +2.0 ± 0.7 ( Figure 6, Table 1). The model Sm-Nd age (TDM) of the analyzed samples is close to 2.7-2.8 Ga. The mineral Sm-Nd isochron on pyroxenes, chrome spinellid, and rock (orthopyroxenite, sample A-1, data from [47]) conforms to the age of ca. 1.87 Ga ( Figure 6, Table 1). The estimated age is close to the time of the metamorphic event at the turn of 1.9 Ga [62].  *: data from [47].

Dating the Hornblende Dykes
The U-Pb and Sm-Nd ages of hornblendite dykes were estimated at ca. 2080 Ma (Figures 8 and  9, Tables 1 and 2). The U-Pb age on rutile from hornblendite dykes is 1804 ± 10 Ma, which corresponds with the postobduction cooling of the rocks in the complex to 450-400 °С.
The intrusive magmatism at 2.53-2.39 Ga was widespread not only in the Fennoscandian Shield, but also in Canada [95][96][97][98] and Zimbabwe [99]. Thus, the conducted research allows the reconstruction of the main stages in evolution of the matter in the Pados-Tudra complex. The intrusion originated at the edge of 2485-2475 Ma, at the turn of the Archean and Paleoproterozoic.
The clearly traced rhythmic layering of rocks in the complex suggests a fluid (initial phase) and crystallization differentiation (next stage) of the melt in the intrusive chamber. This event is temporally constrained by the age of rocks in the rhythmically layered series (2485 ± 77 Ma and 2475 ± 38 Ma) defined by the Sm-Nd method.
Hornblendite dykes intruded when the rhythmically layered series of the intrusion and its chromite mineralization were finally formed. Complex U-Pb and Sm-Nd studies have estimated the age of these dykes at ca. 2080 Ma. Though these dykes occur unconformably to rocks in rhythmic series, the dykes intruded much later than the complex formed, before the Lapland granulites and acid magmatism occurred [43]. In general, the estimated age is close to the age of rocks in the Tanaelv and Kolvitsa Melange (2056-2041 Ma) [62,88]. It marks the origination of the Lapland-Kola Ocean [62] in the late Proterozoic. However, it is possible that isotope systems transformed at the level of minerals and rocks as a result of intense metasomatism during the regional metamorphism at the turn of 1.95-1.90 Ga. It could provide much younger ages of hornblendites. This issue should be clarified by supplementary studies. The successive regional metamorphism at the turn of 1.95-1.90 Ga [54,62], overthrusting of Lapland granulites, and the associated acid magmatism provided global changes in the structure of the complex. As a result, dyke bodies of granites and metasomatically altered rocks and epigenetic formations, as well as talc-magnesite, asbestos, and magnetite ores, were formed in the complex. These processes triggered transformations in Sm-Nd isotope systems in minerals and rocks, restarted the "isotope clocks", and produced metamorphic mineral species. Based on isotope Sm-Nd data on metamorphic minerals, the age of postmetamorphic cooling of the complex rocks to 650-600 °С has been estimated at 1872 ± 76 Ma. The U-Pb age on rutile (1804 ± 10 Ma, Figure 8, Table 2) from a hornblendite dyke indicates further cooling to 450-400 °С [88]. The close Sm-Nd age on whole-rock and metamorphic rutile (1872 ± 76 Ma [43,44]) has been estimated for amphibolized metagabbronorites in the Maliy Pados. The estimated age is interpreted as a time constraint of the postmetamorphic cooling of rocks to the temperature of 650-600 °С. It complies with the temperature of the isotope Sm-Nd system closure in granites from 1.89-Ga-old aluminiferous gneisses in the Yavr-Pados-Nota area [88].
This stage is also marked with intense tectonic activity that produced fractures of the northeastern strike, granitic dykes, and ore occurrences of talc-magnesite, asbestos, and superposed magnetite and sulfide ores.
An almost total absence of clinopyroxene in the composition of the rocks can indicate that these ultramafites could form from the primitive mantle source of the komatiitic composition [100], like in the Chapesvara dunite-orthopyroxenite complex [101], which is also part of the Notozero ultramafic belt and similar with the Pados-Tundra complex in composition and settings of occurrence.

Conclusions
The newly obtained dates indicate an older age for the Pados-Tundra complex than the until not admitted age 2.15 Ga age, which is also the age of the Paleoproterozoic mafic-ultramafic Cu-Ni-Co-Cr-PGE systems in the Fennoscandian Shield. According to results of our study, the Pados-Tundra evolution can be reconstructed in three stages: (i) the emplacement of the layered Cr-bearing intrusion occurred at 2485 ± 77 Ma-2475 ± 38 Ma; (ii) horblendite dykes were intrusive at 2075 ± 59 Ma, which can be related to the origination of the Lapland-Kola Ocean; (iii) the complex was reworked at a tectono-metamorphic event at ca. 1.9 Ga, apparently coeval with the regional overthrust of the LGB [54,62]; the final cooling at 450-400 °C is recorded by U-Pb dating of rutile at 1804 ± 10 Ma. Thus, the Pados-Tundra dunite-harzburgite-orthopyroxenite complex (and, probably, other ultramafic intrusions of the Notozero complex) belongs to the East Scandinavian LIP [12], in which metallogenic specialization [11,12,36] suggests the presence of major deposits of Cu-Ni-PGE raw materials and Cr. Thus, the western part of the Kola Peninsula is promising for chromite raw materials, which is economically valuable for development of the raw material base in the Russian Arctic.