Volume 56, Nº 9 (2018)

Interaction between a melt of kimberlite from the Nyurbinskaya pipe (Yakutia) and natural monocrystalline diamonds was studied experimentally at 0.15 GPa and 1200–1250°C in high-pressure and high-temperature Ar gas “bombs.” The loss of diamond weight with slight surface dissolution of diamonds in a Ca carbonate-bearing kimberlite melt over the course of 2 h (the period of kimberlite transport from upper-mantle diamond-forming chambers to the crustal cumulative centers) is 3–4.5%. In 4 and 7–8 days (under the conditions of crustal cumulative centers), the weight of diamond decreases with remarkable bulk dissolution by 13.5 and 24.5–27.5%, respectively. In the run at 0.15 GPa and 1200°C kimberlite and ilmenite (added) melts interact to produce perovskite melt. Both of the melts, rich in titanium minerals, are immiscible with kimberlite melt and therefore cannot influence the diamond dissolution kinetics in the kimberlite melt. The experimental results suggest that precisely the dissolution processes for thermodynamically metastable diamonds in silicate–carbonate kimberlitic magmas are responsible for the effective decrease in the diamond potential of kimberlite deposits. The paper discusses the physicochemical reasons for the decrease in the kimberlite diamond potential during the chemically active history of diamond genesis: from upper-mantle chambers to the explosive release of diamonds and kimberlite material from cumulative centers to the Earth’s surface. The data on experimental physicochemical studies of the origin, analytical mineralogy of inclusions, and isotope geochemistry of diamonds are correlated.

The paper summarizes newly obtained data on the chemical and isotopic composition of groundwaters in the southern part of the Kuznetsk Basin. The territory is demonstrated to host groundwaters of three types, which were distinguished according to the chemical, isotopic, and gas composition of the waters. The waters of type 1 are widespread in areas of sluggish water exchange in coal-bearing terrigenous rocks, bear mineralization of 0.2–4.7 g/L, have pH 7–10, and N₂–CH₄ and CH₄–N₂ gas composition. Their source is biogenic (coal) with δ¹³C from –17.1 to –4.2‰. The sodic waters of type 2 were found relatively recently and only locally, are hosted in the same rocks but at more hampered water exchange, and have mineralization of 5–25 g/L, pH 7.7–10, and CH₄ gas composition. These waters are noted for an anomalously isotopically heavy isotopic composition of their carbon δ¹³С(\({\text{HCO}}_{3}^{ - }\)) = 0.2–30.9‰ and δ¹³С(CO₂) = –3.2–22.3‰. The source of their carbon is also biogenic, but long-lasting interaction in the water–coal system resulted in significant carbon fractionation, with isotopically lighter carbon fractionated into CH₄ (δ¹³С = –67.3 to ‒43.3‰) and isotopically heavier one preferably coming to CO₂. The waters of type 3 occur only at certain sites, for example, the Tersinskii CO₂ waters. These waters have mineralization of 2–5.5 g/L, pH 6.2–6.7, and CO₂ gas composition. Their δ¹³С(\({\text{HCO}}_{3}^{ - }\)) = –4.3 to –4.1‰ and δ¹³С(CO₂) = –12.3 to –3.9‰ suggest a mixed source of their carbon. This source was mostly deep-seated, not genetically related to waters. All of the sodic waters belong to the infiltration type according to their δ¹⁸O and δD, with the waters of type 2 showing a positive oxygen shift. The sodic waters of all three types are shown not to be in equilibrium with many primary aluminosilicates but are oversaturated with respect to carbonates, montmorillonites, illite, chlorite, and sometimes even albite and microcline (waters of type 2). The waters of all of the three types were generated by the same mechanism: by dissolving aluminosilicate minerals in coal-bearing rocks (with which the waters are not in equilibrium) and the simultaneous precipitation of carbonates. The reasons for certain differences between the waters are their interaction with rocks and coal (waters of type 2), which increases their mineralization and makes their carbon and oxygen isotopic compositions heavier, or the addition of carbon dioxide from an external source (waters of type 3), which acidifies the waters.

A hypothesis was proposed that a considerable amount of gaseous HF is supplied to the atmosphere from the soil cover owing to the reaction \({{{\text{H}}}^{ + }} + {{{\text{F}}}^{ - }} = {\text{HF}}{\text{.}}\) In the upper layer of tundra and taiga soils with pH 4.0–4.5, the partial pressure of HF is (65–20) × 10^–12 atm. Correspondingly, the concentration of dissolved F in the condensates of water vapor is 1.6–11 μg/L, which is comparable with F concentration in rainwater.