The Kovdozero and Pados-Tundra Complexes, Kola Peninsula, Russia: Comparable Geochemistry and Age

Abstract

Geochemical whole-rock variations in the Kovdozero complex in the Lapland–Belomorian Belt (LBB) are compared with those observed in the Pados-Tundra layered complex in the Serpentinite Belt (SB) in the complementary structure in the Fennoscandian Shield. A great variety of coronitic associations exists in the entire LBB–SB system. The Kovdozero complex largely consists of more evolved products of crystallization. Our results of U–Pb dating (zircon and baddeleyite) give the dates of 2514 ± 5 and 2478 ± 6 Ma, leading to the revised age ~2.5 Ga for the Kovdozero complex. It is thus considered to be coeval with Pados-Tundra, Perchatka, and gabbro–anorthosite associations of the Belomorian province in the White Sea region. The variation trends are generally extensive, continuous and close to linear at Kovdozero, which point to crystallization of chonolithic bodies of the complex from a single portion of melt, in separate reservoirs that likely communicated to develop as a whole in the connected system. The extreme degree of differentiation of derivatives of the initial komatiitic magma occurred in the large-scale plume. It led to the development of shallowly emplaced complexes grading from dunitic rocks and associated chromitites with Ru–Os–Ir mineralization at Pados-Tundra (the center) to leucocratic gabbroic rocks at Kovdozero, and likely to gabbro–anorthosite rocks of the Belomorian province (the periphery); these are considered the final products in the megastructure. The εNd(T) values are slightly negative at Kovdozero: −0.43 and −0.60. They imply some degree of crustal contamination of the initial magma. The generalized date of 2.5 Ga likely represents the age of the coronitic complexes of ultrabasic–basic rocks that crystallized from portions of komatiite-derived melts in hypabyssal settings of the LBB–SB megastructure in the eastern Fennoscandian Shield.

1. Introduction

Many hypabyssal ultrabasic–basic complexes of Paleoproterozoic age are distributed across the eastern Fennocandian Shield, both in Russia and in Finland. From the border, the Serpentinite Belt (SB) of complexes (Figure 1) can be traced northeasterly for close to 100 km (e.g., [1,2,3,4] and references therein) and westward into Finland as the Tulppio Belt (e.g., [5,6]). The Pados-Tundra complex is the largest and one of the most thoroughly investigated representatives of the SB complexes [7,8,9,10]. The Lapland–Belomorian Belt (LBB) extends southeasterly toward the White Sea along a trend that parallels Kandalaksha Bay (Figure 1). Various complexes have been described ([11,12,13,14,15,16,17], and references therein). Many of the SB and LBB complexes display a coronitic texture. Whereas some of the authors just cited have favored a metamorphic origin, detailed work on several of the SB and LBB complexes shows that the texture reflects a combination of magmatic and deuteric (autometasomatic) processes caused by crystallization in evolving conditions of temperature and oxygen fugacity in a shallow setting [16,17,18].

In this article, we wish to compare the composition of the two major representatives of the belts mentioned above, Pados-Tundra and Kovdozero. Both complexes contain a comparable spectrum of ultrabasic rocks. In which way do they differ? How do those two complexes compare in terms of age? Why is the coronitic texture so well developed, and what is the significance of a garnet-group mineral involved in a coronitic texture? The geochemical and isotopic data provide new insight into the nature of the SB–LBB megastructure.

2. The Geological Context

The Kovdozero complex, one of the most important constituents of the LBB–SB megastructure, is located along the northern shore of Lake Kovdozero in the southern sector of the LBB structure, about 30 to 55 km west of the northern extent of Kandalaksha Bay on the White Sea (Figure 1 and Figure 2a). The area is marked by large chonolithic bodies of ultrabasic–basic rocks interspersed with blocks of gneissic rocks of the Archean Lower Keret’sky series (Figure 2a). There are rocky exposures, some with a well-developed parting (Figure 3a,b). The ultrabasic–basic rocks are stratiform, relatively fresh and typically undeformed, with deuteric modifications of variable extent.

The layered successions consist of variants of lherzolite; olivine- and plagioclase-bearing websterite; plagioclase orthopyroxenite; norite; locally olivine-bearing gabbro; gabbronorite; and gabbro. Some rocks contain minor amounts of quartz and K-feldspar in the mode. Fine-grained or “taxitic” textures (i.e., with a heterogeneous grain-size) are relatively common, especially closer to the northern contact. Amphibolite formed at the expense of gabbro is rare.

Northwest of Kovdozero, the complexes Perchatka, Rogomu, Yanisvaara and Tepsi are also members of the LBB—SB system ([18], and references therein). The complementary SB–TB structure to its north includes the hypabyssal or subvolcanic complexes of Pados-Tundra, Khanlauta, Chapesvara, Lotmvara and Lyavaraka.

The geological structure of the Pados-Tundra layered complex (Figure 1) was described previously [7,10]. Paleoproterozoic suites of komatiites and komatiitic basalts also occur in the southeastern Fennoscandian Shield, namely in the Vetrenyi (Windy) Belt [20,21]. The numerous layered intrusions of the SB–TB structure formed during an interval of continental rifting in the shield at ~2.45–2.5 Ga (e.g., [22]).

3. Materials and Methods

In this investigation, thirty samples were collected across the western segment of the Kovdozero complex, in which the major part of the complex is exposed (Figure 2b). These rocks were analyzed using XRF and ICP–MS techniques; results are listed in Supplementary Tables S1 and S2. They are used in the next section for a comparison with the Pados-Tundra complex, the largest complex in the complementary SB structure that exists to the north of the LBB structure (Figure 1). Twenty-one representative compositions from the Pados-Tundra complex, sampled previously along traverses shown in [7,10], are listed in Supplementary Tables S3 and S4.

The composition of minerals was acquired at the Analytical Center for Multi-Elemental and Isotope Studies, Institute of Geology and Mineralogy, Siberian Branch of Russian Academy of Sciences in Novosibirsk. A JEOL JXA-8230 instrument (JEOL Ltd., Akishima, Tokyo, Japan) was used in wavelength-dispersion spectrometry (WDS) mode. An accelerating voltage of 20 kV and a probe current of 20–50 nA were used. We employed Kα analytical lines for all elements except for Cr, where the Kβ₁ line was used because of peak overlap. Periods of measurements at the peaks were 20 or 10 s. The superposition of the TiKβ₁ line on the VKα line and of the VKβ line on the CrKα line was accommodated. The beam diameter was ~1 μm. Natural specimens of olivine (Mg, Si, Fe and Ni) and chromiferous or manganiferous garnet (Ca, Cr and Mn) were used as standards for olivine. A natural specimen of magnesian chromite (for Cr, Fe, Mg and Al), manganiferous garnet (Mn), ilmenite (Ti) and synthetic oxides NiFe₂O₄ (Ni), ZnFe₂O₄ (Zn), and V₂O₅ (V) were used as standards for chromian spinel. Grains of orthopyroxene and amphibole were analyzed using pyrope (Si, Al, Fe), a glass Ti standard (GL-6), chromiferous garnet (Cr), diopside and pyrope (Mg, Ca), manganiferous garnet (Mn), albite (Na), and orthoclase (K). Clinopyroxene was analyzed with essentially the same set of standards, and diopside was used for Si, Ca, and Mg. The following set of standards was applied in the analysis of plagioclase: orthoclase (Si, Al, K), diopside (Ca), pyrope (Fe), and albite (Na). The data were processed with the ZAF method of corrections. The calculated limits of detection (1σ criterion) are ≤0.01 wt.% for Ti, Cr, Fe, Ni, Ca, Zn, Mn, and K, and 0.02 wt.% for Na and Al.

Whole-rock abundances of major and trace elements were established at the same center by inductively coupled plasma-mass spectrometry (ICP–MS) using a high-resolution Finnigan MAT mass spectrometer (Finnigan MAT Gmbh, Bremen, Germany, model ELEMENT). Analytical details are provided by [23]. Contents of major oxides in the rocks were established by X-ray fluorescence (XRF) using molten tablets. Samples were first dried at 105 °C for 1.5 h, then calcined at 960 °C for 2.5 h and mixed with a flux (66.67% lithium tetraborate, 32.83% lithium metaborate and 0.5% lithium bromide) in a ratio of 1:9. Total weight of the mixture was 5 g. The mixture was melted in a platinum crucible in a Lifumat-2.0-Ox induction furnace. Measurements were performed on an ARL-9900XP X-ray fluorescence spectrometer (Thermo Fisher Scientific Ltd., Waltham, MA, USA). The following set of state-approved standards were used to build calibration relationships and control the correctness of the analysis: MU-1, MU-3, MU-4, SA-1, SChT-1, SChT-2, SDO-1, SDU-1, SG-1A, SG-2, SG-3, SGD-1, SGD-2, SGKh-1, SGKh-5, SGKhM-2, SGKhM-3, SI-1, SI-2, SNS-1, SNS-2, SOP-1, ST-1, plus synthetic mixtures based on components of MgO, SiO₂, Al₂O₃, TiO₂, CaO, CaSO₄, Cr₂O₃, and Fe₂O₃. The analytical accuracy corresponds to the category.

For the U-Pb isotope study, the samples were initially subjected to hydrothermal decomposition in concentrated (48%) HF acid at 205–210 °C for 1 to 10 days [24]. Then 3.1N HCl was added to dissolve the fluorides at a temperature of 130 °C for 8 to 10 h. The sample was divided into two aliquots in 3.1N HCl to measure the isotopic composition of lead and determine the contents of lead and uranium, where a mixed tracer ²⁰⁸Pb + ²³⁵U was used. The isolation of lead and uranium for isotopic studies was carried out on AG anionite 1 × 8, 200–400 mesh, on columns made of fluoroplast. The blank intra-laboratory contamination of the complete analysis was less than 0.1–0.08 ng for lead and 0.01–0.04 ng for uranium. The U-Pb analyses were performed on a seven-channel Finnigan-MAT 262 mass spectrometer (Finnigan MAT Gmbh, Bremen, Germany) in static mode. The isotopic composition of lead was measured using an ion-counting multiplier, with silica gel used as an ion emitter. The errors in determining the isotopic composition of lead were 0.025% according to the SRM-982 standard. The content of lead (1350–1450 °C) and uranium (1450–1550 °C) were measured in a single-tape mode with the addition of H₃PO₄ and silica gel according to the method of [25]. All isotopic ratios were corrected for mass discrimination by studying parallel analyses of the SRM-981 and SRM-982 standards and equal to 0.12 ± 0.04%. The error in the U-Pb ratio was calculated using statistical calculations of parallel IGFM-87 standards and assumed to be 0.5%; where the errors are higher, the actual values are given. The coordinates of the points and the isochron parameters were calculated according to [26]. The ages were calculated using the accepted values of the uranium decay constants [27], with the errors given for ±2σ. According to the Stacey and Kramers model [28], a correction was made for the presence of common lead. Adjustments were also made for the isotopic composition of plagioclase in cases where the presence of common lead is more than 10% of the total amount of lead, and the isotopic ratios of ²⁰⁶Pb/²⁰⁴Pb are less than 1000.

The contents and isotopic compositions of neodymium and samarium were determined in the Laboratory of Geochronology and Isotopic Geochemistry at the Kola Science Centre, Russian Academy of Sciences in Apatity, using a Finnigan-MAT 262 solid-phase mass spectrometer (Finnigan Mat Gmbh, Bremen, Germany). The chemical decomposition of samples mixed with a ¹⁵⁰Nd–¹⁴⁹Sm tracer was performed according to the procedure described in [29]. The measurements were carried out in static mode using a Ta tape as an evaporator and a Re tape as an ionizer (two-tape mode). Before the use, the Ta and Re tapes were degassed in vacuum for 30 min at currents of 4–4.5 A. The Nd and Sm assays were dissolved in 1 µL of 0.2 N HNO₃; then a drop was placed on a degassed tantalum tape and dried at 1.2 A. The drop-application procedure was continued until 1 µL of pipette disappeared, and then the tape was left to warm at a current of 1.2 A for seven minutes. After that, the current was increased until the maroon glow of the tape appeared and it was left in this mode for two seconds. The measurements were carried out by integrating the data into 120–150 scans (12–15 measurement blocks of 10 scans each). The measurement protocol provided an error rate for determining the isotopic composition of Nd in individual analysis of less than ±0.003% with a ratio ¹⁴⁷Sm/¹⁴⁴Nd of ±0.3%. The quality of the isotope analyses was evaluated by repeated measurements of the standard. The average value of the ratio ¹⁴³Nd/¹⁴⁴Nd in the JNdi-1 standard was 0.512065 ± 11 (2σ, n = 3). The error in determining the ratio ¹⁴⁷Sm/¹⁴⁴Nd was 0.3% (2σ), which is the average of seven measurements in the BCR-2 rock sample [30]. Values of the isotope ratios were normalized to ¹⁴⁶Nd/¹⁴⁴Nd = 0.7219 and then recalculated to the reference value of ¹⁴³Nd/¹⁴⁴Nd in JNdi-1, 0.512115 [31]. The average measured value for the JNdi-1 standard over the measurement period was used to determine the coefficient R = (¹⁴³Nd/¹⁴⁴Nd_JNdi(avg.)/0.512115) for all measured and isotope fractionation-adjusted values of ¹⁴³Nd/¹⁴⁴Nd. The idle intra-laboratory contamination by Nd was 0.3 ng and by Sm, 0.06 ng. The accuracy of determining the contents of Sm and Nd was ±0.5%. Modern values of CHUR (¹⁴³Nd/¹⁴⁴Nd = 0.512630, ¹⁴⁷Sm/¹⁴⁴Nd = 0.1960) [32]), DM (¹⁴³Nd/¹⁴⁴Nd = 0.513151, ¹⁴⁷Sm/¹⁴⁴Nd = 0.2136 [33]), and the decay constants of samarium λ₁₄₇ = (6.524 ± 0.024) × 10⁻¹²/year [34] were used to calculate the values of ε_Nd(T) and the model ages T_NdDM.

4. Results and Observations

4.1. Two Kovdozero Samples Are Selected for U–Pb Age Determination

Two Kovdozero samples, KVDOZ-15 and KVDOR-59, were selected for U–Pb dating and a study of Sm–Nd isotopes. The WDS composition of the coexisting minerals is given in Supplementary Table S5.

Sample KVDOZ-15 consists of an unaltered plagioclase-bearing orthopyroxenite (Figure 4a). It contains up to ~80 vol.% orthopyroxene (Wo_2.3–4.5En_70.8–72.2Fs_24.7–26.5; Mg# 72.9–74.5), with ~2–3 vol.% clinopyroxene (Wo_42.5–48.0En_44.0–47.0Fs_8.1–10.5; Mg# 81.7–84.5) and ~15 to 20 vol.% of interstitial zoned plagioclase (Or_0.4–3.3Ab_56.1–95.3An_4.3–42.9) associated with minor amounts of K-feldspar (Or_66.9–90.7Ab_9.3–32.5An_0–1.9). Members of the chromite–hercynite series occur sporadically as small grains (up to ~0.1 mm) having a mean composition (Fe²⁺_0.89Mg_0.08Zn_0.03Mn_0.01Ni_<0.01)_Σ1.01(Cr_0.91Al_0.67Fe³⁺_0.37V_0.02Ti_0.02)_Σ1.99O₄. Also present are rare grains of ilmenite, vanadiferous chromian magnetite (Cr₂O₃ 17.03–17.59 wt.%; 2.83–2.90 wt.% V₂O₃), zircon and baddeleyite in grains 0.2–0.3 mm across.

Three zircon fractions were prepared. The first consists of fragments of translucent dark brown crystals with a glassy luster. The second fraction consists of crystals of prismatic shape (Figure 5a). They are light yellow and translucent, with a glassy luster. The third fraction includes water-clear fragments of roundish light yellow grains (Figure 5b).

Sample KVDOZ-59 (Figure 4b), a gabbronorite, consists of ~50 vol.% orthopyroxene (Wo_0.7–4.5En_58.2–63.2Fs_33.4–40.1; Mg# 59.3–65.0), ~20% clinopyroxene (Wo_43.5–46.9En_39.6–41.7Fs_11.6–15.0; Mg# 73.3–78.2) with exsolution-induced lamellae of orthopyroxene (Wo_0.7–0.9En_58.2–59.2Fs_39.9–41.1; Mg_58.7–59.7), and up to ~30% of zoned plagioclase (Or_0.7–2.3Ab_37.5–72.4An_25.3–61.7) associated with minor perthitic K-feldspar (Or_65.5–84.2Ab_14.7–32.1An_0–2.7). The rock contains zircon (Figure 5c–h) and baddeleyite (Figure 5i–k).

In sample KVDOZ-59, a first fraction of zircon contains transparent dark cognac-colored grains with a glassy luster and a corroded surface. Their mean size is 0.175 × 0.07 mm, with a length-to-width ratio of 2.5. The mean weight of the crystals is 3.4 × 10⁻⁶ g. The second fraction consists of transparent light cognac-colored grains; their mean grain-size is 0.175 × 0.14 mm, with a length-to-width ratio of 1.25. The mean weight is 13.7 × 10⁻⁶ g. Fragments of dark brown grains of baddeleyite are opaque, with a corroded surface and a greasy luster. The mean grain size is 0.175 × 0.07 mm, with a mean weight of 4.7 × 10⁻⁶ g.

4.2. Garnet in Coronitic Associations

Substantial amounts of accessory garnet, invariably almandine–pyrope–(grossular) in composition, appear in the southern portion of the complex near the Puakhta block (Figure 2b). Corona-type textures involving garnet (Figure 6a–c) are present in a sample of olivine-bearing gabbronorite (KVDOZ-63) and gabbro (KVDOZ-65). Compositions of the associated minerals are listed in Supplementary Table S5. Discrete grains or chain-like aggregates of garnet occur in an intimate association with rim-like grains of pargasitic or edenitic amphibole. The garnet–amphibole assemblage formed at the boundary of grains of igneous clinopyroxene in contact with interstitial plagioclase. In the northern part of the complex, e.g., in sample KVDOR-16 (Figure 6d), garnet is present locally as submicrometric grains only, commonly dendritic, or as aggregates or masses formed at the expense of intercumulus grains of plagioclase in association with orthopyroxene grains that have a notably low-Mg composition.

The orthopyroxene and clinopyroxene in KVDOZ-63 correspond to Wo_2.8–4.8En_76.0–77.3Fs_19.1–20.3, Mg# 79.1–79.9 (n = 6), and Wo_36.7–45.6En_46.5–52.6 Fs_7.9–10.6, Mg# 83.2–85.4 (n = 4), respectively (Supplementary Table S5). The olivine is Fo_72.1–72.7 (n = 6). Interstitial grains of plagioclase are zoned: Or_0.3–0.5Ab_30.2–45.5An_54.0–69.5 (n = 12). Accessory grains of spinel (sensu lato) correspond to magnesian–chromian hercynite: (Fe²⁺_0.61–0.72Mg_0.27–0.37Mn_0.005Zn_0.01Ni_{0.002–0.004})(Al_0.88–1.27Cr_0.63–0.93Fe³⁺_0.09–0.16)O₄ (n = 4) or to magnesian aluminous chromite associated with grains of magnesian ilmenite (MgO 3.23–4.54 wt.%; n = 6).

The almandine is enriched in the pyrope component, with MgO contents in the range 9.45–10.36 wt.%. The mean composition of the garnet grains (WDS EMP, n = 12) is as follows: SiO₂ 39.48, TiO₂ 0.01, Al₂O₃ 22.27, Cr₂O₃ 0.03, FeO 20.46, MnO 0.95, MgO 9.88, CaO 6.94, Na₂O 0.05, K₂O 0.01, Total 100.08 wt.%, or Alm_42.7Py_36.8Grs_18.6Sps_2.0. The associated pargasite corresponds to (Na_0.79K_0.18)_Σ0.97Ca_1.84[(Mg_3.09Fe²⁺_0.90Mn_0.01)_Σ4.00Al_1.08](Si_6.07Al_1.93)_Σ8.00O₂₂(OH)₂. The garnet in sample KVDOZ-65 (Figure 6b,c) is compositionally similar: Alm_44.9Py_37.0Grs_16.3Sps_1.8. It is intimately associated with plagioclase (Or_0.4–0.5Ab_35.3–37.0An_62.5–64.3) and an aluminous variant of edenite (Na_0.55K_0.18)_Σ0.73Ca_1.91(Mg_3.14Fe²⁺_0.95Al_0.77Mn_<0.01Ti_0.09Cr_0.02)_Σ4.97(Si_6.51Al_1.49)_Σ8.00O₂₂(OH)₂. Primary grains of clinopyroxene are magnesian diopside: Wo_47.1–48.8En_43.9–45.9Fs_6.8–7.7; Mg# 85.2–86.9.

In plagioclase orthopyroxenite (sample KVDOZ-16), submicrometric grains of garnet have a similar composition, Alm_42.0Py_36.3Grs_20.1Sps_1.7; it is associated with intercumulus grains or pargasite and An_≤57 or their relics (Figure 6d). Interestingly, grains of orthopyroxene generally have Mg# in the range 74.6–75.4, with a notable buildup to Mg# 77.9–78.1 recorded in those grains that host abundant inclusions of symplectitic magnetite (Supplementary Table S5).

4.3. Results of the Geochemical Analyses: A Comparison of the Two Complexes

As noted, plagioclase-bearing lherzolite, orthopyroxenite and websterite as well as olivine-bearing gabbronorite, gabbronorite and gabbro are represented at Kovdozero, whereas the dunite–harzburgite–orthopyroxenite association is present in the Pados-Tundra layered complex. Thus, the geochemical variations are much more extensive in the Kovdozero complex than at Pados-Tundra (Supplementary Tables S1–S4). The Kovdozero suite is compared with rocks representative of the Pados-Tundra complex (new data for this study) using the plots SiO₂—MgO, MgO—NiO, Al₂O₃—CaO and Na₂O—K₂O, all in wt.%, as well as Co—Ni and Rb—Zr, expressed in ppm (Figure 7a–f). The Mg# index, which is the molar ratio (bulk rock) 100MgO/(MgO + FeO_tot + MnO), is examined as a function of Sc (ppm) and of the index (Gd/Yb)_N. In addition, the Kovdozero and Pados-Tundra complexes are compared in plots Ce vs. La and ΣLREE vs. ΣHREE + Y (Figure 8a–d). Their profile in terms of rare-earth elements is shown on a chondrite-normalized plot (Figure 9a).

The comparison reveals the following. (1) The Pados-Tundra dunite displays the greatest enrichment in magnesium among the ultrabasic suites that have been investigated in the region. The bulk-rock contents of Cr₂O₃ generally range from 0.61 to 8.1, with anomalous samples containing up to 18.8 wt.% (Supplementary Table S3). A maximum of 26.7 wt.% is recorded in mineralized dunite in ore zones at Pados-Tundra [10]. The elevated Co in the chromitites correlates sympathetically with Ni; this relationship, also noted at Kovdozero (Figure 7e), is typical of the other complexes investigated in the megastructure. (2) The rocks at Pados-Tundra are strikingly poor in whole-rock contents of CaO, Al₂O₃, alkalis, and depleted in abundances of incompatible and high field-strength elements (Figure 7a–f and Figure 8a–d). (3) The chondrite-normalized spectra based on mean compositions (Figure 9a,b) indicate that suites of the SB structure contain extremely low abundances of HREE; some samples reveal a pronounced negative Eu anomaly. As a group, they are more primitive than the Kovdozero suite as well as other suites of the LBB structure [36]. The REE abundances in the mean spectrum are one order of magnitude lower at Pados-Tundra than in the western Kovdozero (Figure 9a). The Pados-Tundra spectrum is strikingly close to that documented for the Upper Contact Facies, UCF, in the Chapesvara sill. Note that the UCF rock is inferred to approach composition of the Al–undepleted komatiitic parental melt [1].

4.4. Age Determination and εNd Values

The U–Pb dating (Figure 10a,b;

Table 1. U-Pb geochronology of zircon and baddeleyite from the Kovdozero complex, Kola Peninsula.

Sample KVDOZ-15	Sample Weight (mg)	Content (ppm)		Lead Isotope Ratios *			Isotope Ratios and Age (Ma) **			Rho
		Pb	U	206Pb/204Pb	206Pb/207Pb	206Pb/208Pb	207Pb/235U	206Pb/238U	207Pb/206Pb
1	0.5	301	505	2119	6.147	1.892	10.509	0.4601	2480	0.99
2	0.4	159	244	477	5.518	1.655	9.381	0.4102	2397	0.99
3	0.3	62	76	113	3.624	1.272	9.132	0.4001	2513	0.92
Zircon
1	0.5	516	1043	10,817	6.4352	3.5562	8.484 ± 0.057	0.3990 ± 0.0027	2393 ± 1	0.9
2	0.4	246	512	2989	6.4481	3.0999	7.818 ± 0.025	0.3760 ± 0.0012	2355 ± 2	0.9
Baddeleyite
3	0.3	240	1203	4915	6.1252	37.515	2.461 ± 0.022	0.1888 ± 0.0009	2489 ± 3	0.6

* All ratios are adjusted for idle contamination of 0.08 ng for Pb and 0.04 ng for U, and mass discrimination of 0.12 ± 0.04%. ** The correction for the admixture of common lead was determined according to the model of [28].

) of zircon gave the age 2514 ± 5 Ma, based on accumulated values obtained from the three zircon fractions extracted from plagioclase orthopyroxenite KVDOZ-15 in the Gabrish area. A second date, acquired on gabbro KVDOZ-59, is generally consistent but slightly younger: 2478 ± 6 Ma (Figure 10b).

The calculated εNd(T) values at Kovdozero are −0.43 (KVDOZ-15) and −0.60 (KVDOZ-59) (

Table 2. Sm—Nd isotope data on rocks of the Kovdozero layered complex, Kola Peninsula.

Sample	Content (mcg/g)		Isotope Ratios		TNdDM (Ma)	εNd (T = 2514)
	Sm	Nd	147Sm/144Nd	143Nd/144Nd
KVDOZ-15	1.914	10.82	0.1071	0.511138 ± 11	2870	−0.43
KVDOZ-59	1.562	7.94	0.1190	0.511342 ± 9	2903	−0.29

The values of ε_Nd(T) and model ages T_NdDM were calculated based on modern values of CHUR ¹⁴³Nd/¹⁴⁴Nd = 0.512630, ¹⁴⁷Sm/¹⁴⁴Nd = 0.1960 after [32], DM ¹⁴³Nd/¹⁴⁴Nd = 0.513151 and ¹⁴⁷Sm/¹⁴⁴Nd = 0.2136 after [33], using the decay constants of samarium: λ₁₄₇ = (6.524 ± 0.024) × 10⁻¹²/year [34].

). These values are plotted as a function of age determinations in the literature for different groups of basic–ultrabasic complexes in the eastern Fennoscandian Shield (Figure 11).

Characteristics of the Kovdozero and Pados-Tundra complexes are compared in

Table 3. Comparison of characteristics of the Kovdozero and Pados-Tundra complexes.

Characteristic	Kovdozero	Pados-Tundra
Regional structure	Lapland–Belomorian Belt	Serpentinite Belt
Intrusion shape	Connected system of chonolithic layered bodies	Lopolite-like layered intrusion [7]
Isotopic age	2514 ± 5 Ma (zircon U–Pb dating) (this study)	2485 ± 38 Ma (Sm–Nd dating) [9]
Values of εNd(T)	Slightly negative (−0.43; −0.60) (this study)	Positive (2.0) [9]
Sequences of rocks	Peridotite (mainly lherzolite)—olivine- and plagioclase-bearing websterite—plagioclase orthopyroxenite—norite (locally olivine-bearing)—olivine-bearing gabbro—gabbronorite—gabbro (some quartz-bearing).Rocks enriched in Cpx and Pl are abundant	Dunite—harzburgite—orthopyroxenite series (Cpx and Pl are typically absent) [7]
Maximal value of Fo in olivine	Fo87 (our data)	Fo93
Geochemical characteristics	Rocks display trends of relative enrichment in CaO, Al2O3, alkalis, REE, incompatible and high field-strength elements (this study)	Rocks are strikingly poor in CaO, Al2O3, alkalis, and depleted in incompatible and high field-strength elements with extremely low abundances of REE (this study)
Series of spinel-subgroup members	Chromite—spinel—hercynite series (enriched in Al) (our data)	Magnesiochromite—chromite series (extending to chromian magnetite) [7,10]
Ore zones	Low-sulfide zone of Pd—Pt mineralization in mafic rocks (our data)	Chromium and Ru—Os—Ir mineralization in chromitite and dunite [8,10]

.

5. Discussion

5.1. Inferences from the Isotopic Data

The pyroxene of the two samples chosen for zircon–baddeleyite extraction has low-Mg# values and thus indicates an advanced extent of magmatic differentiation. The level of Zr had risen sufficiently to cause zircon and baddeleyite to grow. Plagioclase orthopyroxenite (KVDOZ-15) likely crystallized from a fractionated portion of melt near the margin of a chonolithic body (Figure 2a). Sample (KVDOZ-59) is representative of the most fractionated members of the layered series in the southern portion as the front of crystallization likely moved north to south [18]. It is located relatively close to the Puakhta block (Figure 2b), sampled by [48] for zircon geochronology. Our observations show that autometasomatic effects became more pronounced in this direction owing to an accumulation of volatiles during crystallization. It is in this portion of the complex that garnet appears as an accessory phase in the coronitic associations.

Our ²⁰⁶Pb/²³⁸U vs. ²⁰⁷Pb/²³⁵U data based on the three zircon fractions of KVDOZ-15 and plotted on the discordia line give an upper concordia intersection of 2514 ± 5 Ma, MSWD = 0.52 (Figure 10a). The sample is unaltered (Figure 4a); the date obtained likely corresponds closely to the time of crystallization of the complex. Sample KVDOZ-59 yields a consistent value, 2478 ± 6 Ma, MSWD = 0.20, based on two fractions of zircon and one of baddeleyite; the latter data-point reflects lead loss as a result of metasomatic activity (Figure 10b).

The previous result of U–Pb dating, 2440 ± 10 Ma, pertains to zircon fractions analyzed from a plagioclase-enriched segregation in a gabbronorite (58% Pl, 29% Opx, and 10% Cpx in the mode) in the Puakhta block. The authors [48] noted the high contents of both U and Pb, up to 1000 and 600 ppm, respectively, which could be a reflection of processes at a postmagmatic stage.

The date obtained on the fresher rock at Kovdozero, 2514 ± 5 Ma, conforms well to results reported from the White Sea region in the Belomorian mobile belt, as well as from other suites in the LBB—SB megastructure. A zircon age of 2568 ± 17 Ma was obtained for an anorthositic rock of the Pezhostrovskiy complex in the Chupa Gulf [49]. Results of zircon dating are also consistent from gabbro–anorthosite of the Vorochistoozyorskiy complex: 2505 ± 8 Ma [45]. Similar zircon ages, 2450 ± 10 and 2460 ± 11 Ma, respectively, were obtained from anorthosite of the Kolvitsa complex and gabbronorite–anorthosite rocks of the Voroniy Island in the Kandalaksha Archipelago [50,51]. Results of Sm–Nd dating of the Perchatka complex, 2485 ± 51 Ma, and the Pados-Tundra layered intrusion, 2485 ± 38 Ma [9,14], located in the central LBB and SB, respectively, also agree fully.

The Nd isotopic data acquired for Kovdozero are entirely consistent (Figure 11); they imply a single source for the samples taken in two different parts of the complex. The εNd(T) value observed, −0.43 for KVDOZ-15, agrees satisfactorily with εNd(T) = −0.60 observed for KVDOZ-59. Our results at Kovdozero are consistent with those from the Perchatka complex in the central LBB, for which the ages 2.51–2.48 Ga are reported with variations in εNd(T) values from −0.96 to 1.05 [14]. In contrast, the εNd(T) value is positive (2.0) for the Pados-Tundra complex in the SB [9], as also is the εNd(T) value (+0.5) characteristic of komatiitic suites in the Kuhmo belt, Finland [52]. Results from the gabbro–anorthosite associations in the Belomorian province yield an age of ~2.5 Ga with εNd(T) values in the range 0.3–1.4 [45]. The overall variations in the εNd(T), plotted as function of age, display a field of data-points for the eastern Fennoscandian Shield (Figure 11). In general, the lower εNd(T) values correlate with the younger isotopic ages. Our isotopic data at the western Kovdozero compare well with those plotted close to the age 2.5 Ga with the corresponding εNd values constrained at a range of slightly negative to positive values (Figure 11).

The Pados-Tundra and Kovdozero complexes evidently have a comparable age. The positive ε_Nd(T) value of Pados-Tundra implies that it was derived from a depleted mantle source. In contrast, Kovdozero, with its negative ε_Nd(T) values, indicates derivation from a source contaminated by crust and enriched over time.

5.2. Inferences from the Mineralogical and Geochemical Data

The two complexes are generally compared in Table 3. An advanced degree of magmatic differentiation is well displayed at Kovdozero by the extent of the geochemical variations relative to the significantly more primitive rocks of the Pados-Tundra complex in the SB (Figure 7a–f, Figure 8a–d and Figure 9a). The Pados-Tundra intrusion was emplaced at the junction of the two complementary branches: the SB–TB and LBB appear to have a conduit-like structure. This disposition resembles that of the Monchepluton complex, in which the Dunite Block is the likely conduit [53], located at the conjunction of the two branches, the Nittis–Kumuzhya–Travyanaya and the Sopcha–Nyud–Poaz suites. The conduit structure is expressed by deposition of chromite–magnesiochromite, which cocrystallized with hypermagnesian olivine (Fo₉₆) [53]. In addition, the komatiitic source is directly indicated by occurrences of spinifex-textured clinopyroxene (hypermagnesian) documented at Tepsi in the northern part of LBB structure [54].

Listed here are mineralogical criteria relevant to an overall comparison of the SB and LBB suites. (1) A southward decrease in the maximal Fo contents is recorded, from Fo₉₃ (Pados-Tundra) and Fo₉₂ (Chapesvara), Fo₉₁ (Tepsi), Fo₈₇ at Perchatka and Rogomu, to Fo₈₇ (our data) or Fo₈₆ in the Gabrish area of the Kovdozero complex. (2) The chromite–magnesiochromite deposits and associated Ru–Os–Ir mineralization are hosted by dunitic rocks at Pados-Tundra. There are no significant volumes of chromite ores known in other suites of the megastructure. (3) Compositions of spinel-group minerals become essentially more aluminous (and less chromium-enriched) southward. (4) Modal amounts of clinopyroxene and plagioclase are strikingly greater in complexes in this direction.

The geochemical data provided in the previous section also serve as criteria in the comparison. (1) The Pados-Tundra dunite displays the highest level of magnesium among the ultrabasic suites represented in the megastructure. (2) The rocks at Pados-Tundra are geochemically unevolved. (3) The SB suites contain extremely low abundances of HREE relative to the LBB suites [36]. These data support the idea that the Pados-Tundra–Chapesvara area represents the major magma conduit for the SB–(TB)–LBB megastructure.

5.3. Coronitic Associations

A great variety of corona-type associations and reactions exist in the related LBB complexes at Kovdozero and Perchatka [17,18]. A prominent variant at Kovdozero is composed of the composite succession Ol → Opx → Cpx → calcic Amp → Pl (+ Qz), which corresponds to the normal order of crystallization of a complete series of phases in ultrabasic–basic complexes. Some of the olivine grains host unusual recurrent layers of orthopyroxene. Another pattern involves an olivine rim around a chromite core, which clearly reflects an initial stage of rapid crystallization of a komatiitic melt at Perchatka. Coronitic textures formed by a combination of magmatic and deuteric (autometasomatic) processes in other ultrabasic suites of the SB [16].

The coronitic sequences at Kovdozero are commonly related to late-stage crystallization of interstitial plagioclase [18]. In some coronas, oriented crystallites of phlogopite deposited on orthopyroxene faces are in direct contact with plagioclase as a consequence of buildup in H₂O in the coexisting fluid phase. In the Puakhta area, the intercumulus plagioclase reacted with primary grains of magnesian clinopyroxene (and H₂O) to produce the coronitic association of almandine–pyrope garnet with aluminous calcic amphibole (pargasitic or edenitic; Figure 6a–c). Essentially the same garnet plus pargasite formed deuterically as a result of breakdown of interstitial plagioclase with the involvement of components of primary grains of orthopyroxene enriched in ferrosilite (Figure 6d). The compositional similarity is recorded for different associations of almandine–pyrope garnet in different parts of the Kovdozero complex. This constancy in the garnet compositions implies that the autometasomatic environment was fairly uniform all over a great volume of crystallizing rocks at the final stage of petrogenesis. Interestingly, the related occurrence of almandine and other aluminous phases, formed autometasomatically at the expense of plagioclase and pyroxene, was also noted in the middle portion of the Nadezhda microgabbronorite sill in the Lukkulaisvaara layered intrusion in northern Karelia, Russia [55].

6. Concluding Inferences

• The new U–Pb dates for the Kovdozero complex, 2514 ± 5 and 2478 ± 6 Ma, lead us to propose that the coronitic complexes of ultrabasic–basic rocks in the entire LBB–SB megastructure were emplaced at approximately 2.5 Ga. This age applies to Pados-Tundra (SB), Perchatka (northern LBB) and the gabbro–anorthosite associations of the Belomorian province to the south, in the White Sea region.
• The ultrabasic and basic complexes of the LBB—SB system crystallized from portions of komatiitic magma in a hypabyssal setting. The Pados-Tundra complex, associated with Chapesvara in the SB, appears to reflect a conduit-type center of a giant plume of such magma. Separate shallow reservoirs of melt likely formed above various parts of the plume in the rift setting.
• Southward from Pados-Tundra to Kovdozero, the SB—LBB suites exhibit a general decrease in Fo in olivine and in Mg–Cr contents, with a buildup in Al in spinel-group minerals, an increase in modal amounts of clinopyroxene and plagioclase and increasing levels of the HFSE and other incompatible elements.
• The leucocratic gabbroic rocks at Kovdozero and the gabbro–anorthosite rocks of the Belomorian province seem to be the final products linked to the diapiric megastructure. The εNd(T) values obtained at Kovdozero, −0.43 and −0.60, indicate a certain degree of crustal contamination of the initial magma.
• A great variety of coronas are formed by the combination of magmatic and autometasomatic processes as consequences of instable and rapid crystallization of komatiite-derived melts in the megastructure.