Chemistry and Crystallization Conditions of Minerals in Metasomatized Oceanic Lithosphere and Basaltic Rocks of Govorov Guyot, Magellan Seamounts, Pacific Ocean

Abstract

Studies have provided first detailed data on the chemistry of rock-forming, minor, and accessory minerals of Govorov Guyot volcanic rocks (basalts, trachybasalts, basaltic trachyandesites, and trachyandesites). Some basalt samples bear pargasitic amphibole and clinopyroxene xenocrysts, mantle vein fragments in xenoliths, as well as wehrlite xenoliths, which are remnants of metasomatized oceanic lithosphere. Amphiboles make up a continuous series from pargasite –Mg-hastingsite in wehrlite xenoliths and xenocrysts to Mg-hastingsite–kaersutite end-members in phenocrysts and microlites of basaltic rocks. The discussed issues include the trace element chemistry of Ti-amphibole and clinopyroxene phenocrysts; fractionation of OIB melts; and P-T equilibration of minerals during the formation of mantle and basaltic rocks. Pargasitic amphibole may have crystallized at P-T conditions (2.5–0.6 GPa, 1170–980 °C) corresponding to the spinel facies of peridotite at different depths (73–21 km) in hydrous (6.0–4.5 wt% H₂O) silicate mafic melts that percolated through peridotites of the oceanic lithosphere. Ti-amphibole in basaltic rocks crystallized at 1.2–0.4 GPa (40–15 km), and 1060–910 °C from melts containing 8.6–2.6 wt% H₂O. As the high-temperature (~1100 °C) basaltic magmas reached chambers at the oceanic crust level (7 to 3 km), the Ti-bearing amphiboles of xenocrysts and phenocrysts became replaced by Ti-magnetite- and/or rhönite-bearing mineral assemblages.

1. Introduction

The arc-shaped chain of Magellan Seamounts (MS) at 10° to 18° N and 150° to 160° E borders the Southern Marcus-Wake trail (Dutton Ridge, Hemler, Himu and other guyots and seamounts) in the north and northwest, the Great Caroline and Marshall Islands in the southeast and east, and the Mariana Trench system in the west (Figure 1a). The MS area was studied in many research cruises after the discovery of abundant Co-rich and Fe–Mn crusts and nodules on the guyots surface, which store economic amounts of many trace elements. Since the 1980s, the work has been run by teams from several countries on research vessels (R/V) that belonged to different academic and industrial institutions: the Soviet (Russian) Academy of Sciences; the USSR Ministry of Geology; the Hawaiian Institute of Geophysics; the Smithsonian Institute of Oceanography; JSC Yuzhmorgeologiya, etc. This part of the Pacific between the Ita-Mai-Tai and Fedorov Guyots, near the Mariana Trench, and in the Ita-Mai-Tai Guyot was studied in the course of the Ocean Drilling Program (ODP) and the Deep Sea Drilling Project (DSDP) [1,2,3,4,5].

The team of JSC Yuzhmorgeologiya (Gelendzhik, Russia) has contributed substantially to the research of MS structures and has collected a great amount of geophysical and oceanographic data for the two past decades during yearly cruises since 1998, including medium- and large-scale multibeam echo-sounding bathymetry, photometric, telemetric, and geoacoustic profiling, dredging, and shallow drilling of volcanic, sedimentary rocks and Co-rich Fe-Mn crusts ([6,7,8,9], and references therein).

So far, no data on mineralogy, chemistry, and crystallization conditions of minerals from the Govorov Guyot volcanics have been reported. We present the first data of this kind for minerals from amphibole (Amp)-bearing volcanics and mantle rocks (xenoliths and xenocrysts with pargasitic Amp, which are remnants of oceanic lithosphere (wehrlite peridotite and fragments of metasomatic veins) sampled in 2016–2017 from Govorov Guyot during cruises on the research vessel (R/V) Gelendzhik as part of Project 06-16 run by JSC Yuzhmorgeologiya. The studies have provided constraints on the chemistry of rock-forming, minor and accessory minerals of basalts, trachybasalts, basaltic trachyandesites, and trachyandesites. The discussed issues include the mineral chemistry of pargasitic Amp, Ti-Amp, clinopyroxene (Cpx), olivine, plagioclase (Pl), oxyspinel group Fe and Ti oxides, nepheline, rhönite, phlogopite, some accessories, and volcanic glasses; trace element chemistry of Ti-Amp and Cpx phenocrysts and fractionation of ocean island basalt (OIB) melts; P-T equilibration of minerals with silicate mafic melts during formation of basaltic rocks, wehrlite, and metasomatic Amp-Cpx mantle veins of the oceanic lithosphere beneath Govorov Guyot.

2. Magellan Seamounts and Govorov Guyot: Geological Setting

The MS chain developed upon the oceanic crust within the Ogasawara fracture zone which was identified according to geomagnetic reversals [10,11,12,13]. The limited amount of data on the ages and composition of the oceanic crust beneath the MS area was collected from ODP and DSDP sites, as well as from samples of guyot volcanic rocks [14,15,16,17,18,19,20,21]. The lower stage and the largest edifices of the middle stage of MS guyots are composed of OIB-type volcanics ([6,7,8,9], http://guyot.ocean.ru accessed on January 2022). According to OIB formation models, the Pacific intraplate volcanism resulted from partial melting of metasomatized oceanic lithosphere exposed to the effect of fluids and melts enriched in volatiles (e.g., H₂O, CO₂, etc.) and alkalis in the region of the South Pacific Thermal and Isotopic Anomaly (SOPITA) [21,22,23].

The structure of the lower volcanic stage remains poorly known because dredging is impossible deeper than 3000–3500 m where slopes near the guyot bases lack Co-rich Fe–Mn crusts and the bedrocks are buried under sediments. It is impossible to estimate the duration of volcanic events from the available radiometric (K-Ar and ⁴⁰Ar-³⁹Ar) ages of samples from the guyot slopes and tops. The basement of the guyots may be composed of Upper Jurassic or Lower Cretaceous tholeiitic NMORB, similar to the surrounding oceanic crust [6,7]. The rocks of the middle stage are exposed in steep (>20°) slopes in the middle and top parts of guyot edifices within 3000–3500 m isobaths. They are mostly subalkalic or alkalic basaltic lavas, which are often covered by palagonitized hyaloclastic deposits, and more rarely, dikes and sills. The upper stage consists of lava flows, lava breccias, tuffs of subalkalic–alkalic basaltic or more felsic volcanics, as well as volcaniclastic and sedimentary rocks. Carbonates include reefal (bioherm, oolite, or bioclastic), shallow-marine or pelagic limestones. Destroyed volcanic and sedimentary rocks affected by seawater alteration are cemented with carbonate coccolith–foraminifera material and form so-called edaphogenic breccias. They coexist with turbidites, gravel, sandstone, and dense clay. Bathial-pelagic soft coccolith–foraminifera carbonates over the guyots fill depressions on the tops of guyots or cover their slopes. The tops and slopes of many guyots are encrusted with Co-rich Fe–Mn.

Govorov Guyot was discovered in 1987 during a cruise of R/V Morskoi Geolog (USSR Ministry of Geology) and named after Russian geologist I.N. Govorov (1920–1997). The guyot is the largest in the MS chain, with its 190 × 180 km base, a trapezium-shaped main body with 70–90 km sides, and a 79 × 53 km flat top (Figure 1b). The main guyot edifice is extended with two satellites oriented in the southwest and southeast directions and large offshoots in the south and northeast. The northern slope is cut by a ~25 km long canyon, and another canyon, 15 km wide and up to 700 m deep, separates the southwest satellite from the northern slopes of Gordin Guyot (Figure 1b,c). The southeast satellite has a ~20 km long top (Figure 1b,d). The two satellites may be separate guyots or parts of the main edifice detached along faults. The slopes of the main edifice dip at 4–8 to 25°; the northern and northeastern slopes are shallower (10–15°) than the western and eastern ones till the 4500 m depth. Locally there are breaks in slope with steps tens to hundreds of meters high and 3–7 km to 30 km long, small terraces, swells or depressions. The main guyot edifice and its satellites are composed mostly of volcanic rocks, as evidenced by photo video profiling, dredging, and shallow (within 1 m) drilling [7]. The rocks are fine to boulder-size lava and tuff clasts, hyaloclastics, edaphogenic breccias, pillow lavas, and lava flows with columnar jointing, exposed on slopes. Magmas erupted under the seawater, judging by the presence of sporadic hyaloclastic deposits. The volcanic rocks reach thicknesses of ~1800 m on the southwestern slope and up to 2900 m on the northeastern slope. The top and low-angle slopes are locally covered by carbonate sediments from the Aptian to Miocene ages. Pliocene deposits consist of siltstones with thin intercalations of volcaniclastics and poorly cemented carbonates. The surfaces free from soft sediments, mainly along the top edge, are covered continuously by up to 12–15 cm thick Co-rich Fe–Mn crusts or occasionally by Fe–Mn nodules. Large areas of the guyot’s top and slopes are occupied by non-cemented Pliocene-Quaternary coccolith–foraminifera sediments.

Govorov volcanic rocks were sampled at two sites during R/V Gelendzhik cruises: the top, edge, and slopes of the main edifice to depths of <3500 m, and near the top of the southeast satellite (Figure 1c,d).

3. Analytical Methods

Samples of volcanic rocks were analyzed for bulk chemistry by XRF on a Bruker AXS S4 Pioneer wavelength dispersive X-ray fluorescence spectrometer using glass fusion discs (major oxides), as well as methods of titration (FeO, CO₂), and gravimetric analysis (H₂O⁺, H₂O^–, and total sulfur) at the Center for Isotope-Geochemical Studies of the Vinogradov Institute of Geochemistry, Siberian Branch of the Russian Academy of Sciences (IGC SB RAS, Irkutsk, Russia). The minor and trace elements of Ti-Amp and Cpx phenocrysts from volcanic samples, as well as REE compositions of Amp-bearing basaltic rocks, were determined by mass spectrometry with inductively coupled plasma (ICP-MS) after acid digestion of mineral grains, on an Agilent NexION 300D quadrupole mass spectrometer (IGC SB RAS, Irkutsk). The quality of XRF and ICP-MS analyses was checked against USGS standards BHVO-2, and AGV-2.

Minerals and phases were investigated by scanning electron microscopy coupled with energy-dispersive spectrometry (SEM EDS) on a Carl Zeiss LEO-1430VP electron microscope equipped with an INCA Energy 350 analytical system at the Geological Institute (GI SB RAS, Ulan-Ude), and on a Tescan Mira-3 LMU high-resolution electron microscope with an Ultim MAX-40 SDD detector (IGC SB RAS, Irkutsk). The reference samples were quartz, albite, K-feldspar, wollastonite, MgO, Al₂O₃, CaF₂, and NaCl synthetic compounds and metals from sets of Reference standards by Microanalysis Consultants Limited (Oxford Instruments Ltd., Abingdon, UK). Matrix correction was by the XPP algorithm of the INCA Energy and AZtecLive proprietary software. Phases were analyzed by scanning over >10 μm² rectangular spots. The operation conditions were: 20 kV accelerating voltage and 50 s counting time for LEO-1430VP and 20 kV and 30 s for MIRA-3 LMU. At these conditions, the detection limits of major elements were 0.2–0.3 wt% [24].

4. Results

4.1. Classification of Volcanic and Mantle Rocks

Amp-bearing volcanic rocks were sampled on the northeastern slope of the main Govorov Guyot edifice and were most often dredged on the northeastern offshoot (Figure 1c). The rocks of the area differ in texture and seawater alteration degree (Figure 2a–d) but all contain Amp, either as more or less numerous phenocrysts or as microlites. Amphiboles also occur as partly or fully replaced coarse xenocrysts (Figure 2e,f) in samples from different parts of the guyot slope, as well as from the southeast satellite (Figure 1d).

The Govorov volcanic rocks have basaltic to trachyandesitic LOI-free compositions normalized to 100 wt% (Figure 3, Table S1). According to TAS diagram, the analyzed rocks plot in the fields of picrobasalt (08D100-3), basalt (08D133A, 08D96), trachybasalt (08D81, 08MTP03, 08TR02, 08D121-8, 08D102-3), basaltic trachyandesite (08D99-2B, 08MTP03-1, 08D123-3, 08121-5), and trachyandesite (08D123-1/5). The strongly altered sample 08D100-3 is classified as basalt given that its SiO₂ content may be underestimated in the presence of authigenic calcite and other secondary phases. It bears coarse Amp and Cpx xenocrysts, xenoliths of large Cpx + Amp intergrowths, as well as orthopyroxene-free peridotite xenoliths with olivine fully replaced by iddingsite aggregate, disseminated inclusions of the spinel and Amp, which were identified as amphibole–spinel wehrlite (Figure 2g–k).

4.2. Rock-Forming Minerals

Amphiboles are of five generations identifiable on the basis of mineral assemblages and morphology: (1) 10–30 µm acicular microlites quite evenly distributed in the groundmass; (2) abundant euhedral phenocrysts or megacrysts, up to 1–3 mm long; (3) 5–8 mm round xenocrysts, partly or fully replaced in the 08D100-3 basalt, which is free from Amp phenocrysts and microlites; (4) disseminated inclusions in wehrlite xenoliths and in Cpx xenocrysts; (5) xenoliths of large intergrowths with Cpx (Figure 2 and Figure 4).

The amphibole crystal chemical formulas were calculated in AMFORM [25], taking into account Fe²⁺, Fe³⁺, and Al occupancy depending on O²⁻ in W-sites (^wO²⁻) and Ti and Fe³⁺/Fe_tot according to the stoichiometry. The end-member percentages of Amp following the IMA nomenclature of the amphibole supergroup were calculated by the mass-balance method in CRYSTAL [26], from formula cation coefficients (in apfu) for each analysis. In basaltic rocks, the best statistics (sum of square misfits ΣΔX² < 0.05, where ΔX is the misfit between the apfu values of cation coefficients from AMFORM and CRYSTAL) were obtained for calcic amphibole end-members: Mg-hastingsite, NaCa₂(Mg₄Fe³⁺)(Si₆Al₂)O₂₂(OH)₂; ferri-kaersutite, NaCa₂(Mg₃TiFe³⁺)(Si₆Al₂)O₂₂O₂; ferro-ferri-kaersutite, NaCa₂(Fe²⁺₃TiFe³⁺)(Si₆Al₂)O₂₂O₂; and ferro-kaersutite, NaCa₂(Fe²⁺₃TiAl)(Si₆Al₂)O₂₂O₂. The mineral chemistry mostly falls in the field of Ti-bearing Mg-hastingsite, while the ferri-kaersutite and ferro-ferri-kaersutite compositions are less frequent (Figure 5,

Table 1. Major element chemistry (wt%), crystal chemical formulas, and thermobarometric–chemometric parameters of Ti-bearing amphiboles.

Sample	08D96	08D81	08D99-2B	08TR02	08MTP03	08MTP03-1	08D121-5	08D100-3
	(15)	(32)	(85)	(13)	(28)	(9)	(10)	(14)	(4)	(6)	(14)	(11)	(15)	(4)
Column Num.	1	2	3	4	5	6	7	8	9	10	11	12	13	14
SiO2	38.65	38.35	39.02	37.76	38.25	38.49	38.95	40.76	39.36	39.71	39.57	38.84	39.18	38.81
TiO2	7.36	5.15	4.71	4.18	4.49	4.01	6.05	3.40	3.59	3.23	4.20	3.97	3.96	4.77
Al2O3	12.29	13.32	12.86	12.94	12.90	12.56	12.41	15.15	15.21	14.12	14.54	14.32	14.50	13.46
Cr2O3	bd	bd	bd	bd	bd	bd	bd	0.56	0.10	0.24	0.13	bd	bd	bd
FeO tot	14.87	14.48	13.99	17.73	15.42	16.36	12.46	6.93	8.33	8.96	9.74	10.42	11.44	7.77
MnO	bd	0.10	0.20	0.41	0.23	0.27	bd	0.05	0.06	0.14	0.12	0.09	0.09	0.07
MgO	10.49	10.62	11.01	8.74	10.17	9.78	11.98	15.05	13.59	13.42	12.81	12.38	12.16	14.68
CaO	11.47	12.11	11.62	12.07	11.84	11.90	12.14	10.70	10.45	11.31	11.05	10.97	11.17	11.53
Na2O	2.54	2.30	2.40	2.33	2.41	2.43	2.79	2.55	2.29	2.39	2.21	2.11	2.12	2.01
K2O	1.65	1.72	1.83	1.81	1.81	1.75	0.73	1.51	1.85	1.82	2.18	2.25	2.29	1.63
Total	99.63	98.14	97.64	97.96	97.51	97.53	97.50	96.66	94.89	95.35	96.58	95.36	96.90	94.73
	Corrected compositions calculated with AMFORM, Ridolfi et al. [25]
SiO2	38.38	38.50	39.34	37.89	38.59	38.78	39.40	41.46	40.75	40.87	40.33	40.04	39.77	40.30
TiO2	7.31	5.17	4.75	4.20	4.53	4.04	6.12	3.46	3.72	3.33	4.28	4.09	4.02	4.95
Al2O3	12.20	13.37	12.97	12.99	13.02	12.65	12.55	15.41	15.75	14.53	14.82	14.76	14.72	13.97
Cr2O3								0.57	0.10	0.24	0.13
FeO tot	14.76	14.53	14.11	17.79	15.66	16.48	12.60	7.05	8.62	9.22	9.93	10.74	11.61	8.07
Fe2O3	6.08	6.42	5.66	7.02	6.90	6.50	4.96	2.91	5.04	4.33	3.70	4.74	4.96	4.76
FeO	9.29	8.76	9.01	11.48	9.35	10.63	8.13	4.44	4.09	5.33	6.60	6.48	7.15	3.78
MnO		0.10	0.20	0.41	0.23	0.27		0.05	0.06	0.15	0.12	0.09	0.09	0.08
MgO	10.41	10.66	11.10	8.77	10.26	9.86	12.11	15.31	14.07	13.81	13.06	12.76	12.35	15.24
CaO	11.39	12.16	11.72	12.11	11.95	11.99	12.28	10.89	10.82	11.64	11.26	11.30	11.34	11.97
Na2O	2.52	2.31	2.42	2.34	2.43	2.45	2.83	2.60	2.37	2.46	2.25	2.17	2.15	2.08
K2O	1.64	1.72	1.85	1.82	1.83	1.76	0.74	1.54	1.92	1.88	2.23	2.32	2.33	1.70
H2O	0.84	1.09	1.20	1.30	1.23	1.32	1.03	1.33	1.25	1.41	1.19	1.20	1.21	1.10
Total	100.05	100.27	100.21	100.32	100.30	100.25	100.16	99.96	99.99	99.98	99.99	99.97	100.08	99.93
	Crystal chemical formulas, apfu
Si	5.831	5.813	5.921	5.802	5.843	5.894	5.893	6.012	5.944	5.998	5.960	5.935	5.913	5.909
Al IV	2.169	2.187	2.079	2.198	2.157	2.106	2.107	1.988	2.056	2.002	2.040	2.065	2.087	2.091
Ti	0.836	0.587	0.538	0.484	0.515	0.462	0.689	0.377	0.408	0.367	0.475	0.456	0.449	0.546
Al VI	0.016	0.192	0.221	0.146	0.165	0.160	0.106	0.646	0.652	0.511	0.542	0.513	0.492	0.324
Fe3+	0.696	0.729	0.641	0.809	0.786	0.743	0.559	0.317	0.553	0.478	0.411	0.528	0.555	0.525
Fe2+	1.180	1.106	1.134	1.470	1.183	1.351	1.017	0.538	0.499	0.654	0.816	0.803	0.889	0.464
Mg	2.358	2.400	2.491	2.002	2.315	2.233	2.701	3.308	3.059	3.022	2.877	2.819	2.737	3.332
Mn		0.012	0.025	0.053	0.030	0.035		0.006	0.007	0.018	0.015	0.011	0.011	0.009
Cr								0.065	0.012	0.028	0.015
Total	13.085	13.026	13.051	12.963	12.995	12.985	13.072	13.258	13.191	13.078	13.153	13.130	13.133	13.202
Ca	1.854	1.966	1.889	1.987	1.938	1.952	1.968	1.692	1.691	1.830	1.783	1.795	1.806	1.881
Na	0.744	0.676	0.705	0.695	0.714	0.721	0.819	0.730	0.671	0.699	0.644	0.624	0.559	0.593
K	0.317	0.332	0.355	0.356	0.353	0.342	0.142	0.284	0.356	0.351	0.420	0.439	0.441	0.317
OH	0.847	1.101	1.201	1.326	1.241	1.342	1.026	1.291	1.215	1.378	1.175	1.191	1.202	1.074
O	1.153	0.899	0.799	0.674	0.759	0.658	0.974	0.709	0.768	0.617	0.816	0.802	0.798	0.926
Mg#	0.557	0.567	0.584	0.468	0.540	0.516	0.632	0.795	0.744	0.728	0.701	0.679	0.655	0.771
	End-members, mol%
Pargasite								49.04	43.19	30.06	30.77	25.82	21.56	27.41
Edenite								2.76	0	1.30	0.66	0	0	3.57
Mg-Hst	25.28	43.82	49.05	44.39	45.19	50.59	45.63	15.85	14.80	32.88	24.71	30.52	35.45	22.87
Fkrs	42.58	20.34	15.30	7.40	16.19	5.85	24.25	17.28	25.06	14.34	19.49	18.15	14.51	35.15
Fe2Krs	12.8	22.38	25.84	16.88	18.33	20.48	24.38	15.06	16.95	21.43	24.37	25.51	28.48	11.00
Fe3Fe2Krs	19.29	13.46	9.81	31.33	20.29	23.08	5.74
	Thermobarometric-chemometric parameters
T, °C	1084	978	1011	987	1001	966	1013	1149	1066	1069	1055	1018	1009	1037
P, MPa	540	493	899	725	918	436	396	1992	1677	1957	1459	1333	1260	713
Depth, km	19	18	31	25	32	16	14	62	54	61	48	44	42	25
H2O, wt%	2.66	5.25	4.23	4.94	4.68	5.14	4.70	4.87	5.30	5.53	4.68	5.15	5.36	3.82

Column numbers 1–7 refer to Ti-Amp in basaltic rock samples, 8–14 refer to pargasitic Amp in sample 08D100-3: wehrlite xenoliths (8), inclusions in Cpx xenocrysts (9 and 10), Cpx – Amp xenolith, a mantle vein fragment (11), xenocrysts (12 and 13), and intergrowths with phlogopite in a Cpx xenocrysts (14). Numerals in braces denote quantity of SEM EDS analyses used to calculate average values. bd = below SEM EDS detection limit. End-members calculated in CRYSTAL [26]: Pargasite, NaCa₂(Mg₄Al)(Si₆Al₂)O₂₂(OH)₂; Mg-Hst = Mg-hastingsite, NaCa₂(Mg₄Fe³⁺)(Si₆Al₂)O₂₂(OH)₂; Edenite, NaCa₂Mg₅(Si₇Al)O₂₂(OH)₂; Fkrs = ferri-kaersutite, NaCa₂(Mg₃TiFe³⁺)(Si₆Al₂)O₂₂O₂; Fe2Krs = ferro-kaersutite, NaCa₂(Fe₃²⁺TiAl)(Si₆Al₂)O₂₂O₂; Fe2Fe3Krs = ferro-ferri-kaersutite, NaCa₂(Fe₃²⁺TiFe³⁺)(Si₆Al₂)O₂₂O₂. Mg#, Mg/(Mg + Fe²⁺ + Fe³⁺). Thermobarometric and chemometric values refer to average compositions according to equations in [27,28], and Depth from Equation number 33 in [29]. Uncertainties: for P, MPa ± 12%, H₂O wt% ± 14%, and T ± 22 °C [28].

). Microlites and rare phenocrysts of Ti-Amp in 08D96 have the most variable compositions: Ti-bearing Mg-hastingsite in phenocrysts and cores of microlites, and 90–70 mol % kaersutitic end-members (ferri-kaersutite + ferro-ferri-kaersutite + ferro-kaersutite) in the microlite rims. Other samples contain phenocrysts with different relative contents of Mg and Fe in growth zones, with alternating darker Mg-rich and lighter Fe-rich zones in BSE images or, more rarely, with Mg-or Fe-rich cores or rims (Figure 4a–c). Some Ti-Amp in 08D81 and Amp remnant xenocrysts in 08D121-8 contain a Cpx core (Figure 2fe). Coarse Amp in partly replaced xenocrysts from 08D121-5 and 08D123-1/5 have Mg-hastingsite compositions with up to 40 mol% kaersutitic end-members. Most of Amp xenocrysts in 08D121-8 have been fully replaced by assemblages: Cpx + Ti-Mag + Pl + nepheline in the rim, and Cpx + rhönite + Pl + nepheline in the core (Figure 4g,h).

Amphiboles in xenoliths and xenocrysts from basalt 08D100-3 (Figure 2g–k) have higher Mg and Al than those from other volcanic rocks. Pargasitic Amp formulas fit best (ΣΔX² <0.07) for the end-members of pargasite NaCa₂(Mg₄Al)(Si₆Al₂)O₂₂(OH)₂, Mg-hastingsite, ferri-kaersutite, and ferro-kaersutite (Table 1). The pargasite end-member is the highest (53–44 mol%) in wehrlite xenoliths, lower (36–26 mol%) in Cpx–Amp xenoliths, inclusions in Cpx xenocrysts, and <31 mol% in xenocrysts (Figure 6). The pargasite and Mg-hastingsite end-member percentages show an inverse linear correlation (Figure 6a). The percentage of the ferro-kaersutite end-member is the lowest (13–18 mol %, Figure 6b) in pargasitic Amp from wehrlite xenoliths, while that of ferri-kaersutite (7–36 mol%, Figure 6c) and edenite NaCa₂Mg₅(Si₇Al)O₂₂(OH)₂ (up to 12 mol%) end-members in xenoliths and xenocrysts varies irregularly. Only one coarse pargasitic Amp xenocryst at the contact with the basalt groundmass is locally replaced by the Cpx + rhönite + Ti-Mag assemblage. Some pargasitic Amp xenocrysts enclose up to 100–200 µm equant grains of spinel and 20–80 µm pyrrhotites partly or fully replaced by siderite, sometimes in intergrowths with an Al-bearing phase compositionally close to jarosite KFe³⁺₃(SO₄)₂(OH)₆.

Clinopyroxene phenocrysts and microlites occur in all samples. Most of the mineral compositions plot within the diopside field in the Wo–Fs–En diagram, with Mg# most often in the range 0.78–0.65 (Figure 7a; Table S2). Some compositions correspond to hedenbergite, with relatively high Fe in 1–5 µm rims of phenocrysts and in microlites (Mg# < 0.65). The BSE images of phenocrysts display alternating dark- and light-colored zones with different Mg/Fe ratios (Figure 4a,b,d). Ti contents are lower in the cores than in the rims of Cpx phenocrysts from all rocks. They are the highest (up to 0.25 apfu Ti) in Cpx microlites from the rhönite-bearing assemblage replacing Ti-bearing Mg-hastingsite xenocrysts and the lowest (<0.06 apfu Ti) in the early core of Cpx phenocrysts and in Cpx from wehrlite xenoliths from basalt 08D100-3. Some Cpx fragments in wehrlite xenoliths have greenish coloration common to Cr-diopside (Figure 2h). Cr-diopside locally shows elongated exsolution lamelli of spinel orientation across the cleavage planes (Figure 4i), while some domains in large (3 × 5 mm) Cpx xenocrysts enclose pargasitic Amp, spinel, Ti-bearing phlogopite, and Ni- and Co-rich pyrrhotite (Figure 4i–l). Wehrlite-hosted Cpx contains 14–15 wt% MgO (Mg# 0.81–0.84) and 0.3–1.3 wt% Cr₂O₃, as well as Cpx in xenocrysts 12–15 wt% MgO (Mg# 0.74–0.80) and 0.1–0.3 wt% Cr₂O₃. Compared to Cpx from basaltic rocks, that from wehrlite xenoliths and xenocrysts in 08D100-3 contains less Ca (0.74–0.86 apfu), Fe²⁺ (<0.18 apfu), Fe³⁺ (<0.10 apfu), Ti (0.01–0.05 apfu), but more Na (0.05–0.10 apfu) and Al _VI (0.08–0.17 apfu) (Figure 7b–d; Table S2).

Plagioclase is the most abundant rock-forming mineral. Most of phenocryst and microlite compositions plot within the fields of labradorite and andesine in the An–Or–Ab diagram, with <10 mol% Or end-member (Table S2). Labradorite commonly occurs in the cores of phenocrysts, while their rims and microlites are composed of andesine. Rare phenocrysts and microlites of Ca-Na-K feldspar with 10–30 mol% Or occur also in 08D96, 08TR02, 08MTP03, 08D99-2B, 08D123-3, and 08D123-1/5. Thin (2–3 µm) rims of this composition in microlites correspond to anorthoclase, often with 1–3 wt% BaO and 1–1.3 wt% SrO. Feldspar in all samples contains up to 1 wt% Fe₂O₃.

Olivine in volcanic rocks extremely rarely survives million-years long pervasive seawater alteration and is most often fully replaced by an aggregate of iddingsite or saponite (a mixture of goethite, smectite, and interstratified phyllosilicates). Fresh olivine phenocryst fragments and microlites have been found only in trachybasalt 08D133A. The Mg# ranges are 0.86–0.75 in the cores and 0.75 to 0.69 in thin Fe-rich rims of phenocrysts and in microlites (Table S2). The Mg-rich zones of phenocrysts contain up to 0.4 wt% NiO.

Oxyspinel Group Fe and Ti oxides occur as irregularly shaped <0.1 mm grains in the matrix of all rock samples. Total iron in their formulas was divided into Fe²⁺ and Fe³⁺ according to stoichiometry, and the end-member percentages were found for each analysis by mass balance calculations in CRYSTAL [26], with ΣΔX² < 0.005 errors in apfu values of cations (Table S2). Cr-spinel compositions were calculated for five end-members, with highly variable percentages: Hercynite (FeAl₂O₄, Hy), 2.5–34 mol%; Spinel (MgAl₂O₄, Sp), 0–60 mol%; Chromite–Magnesiochromite (Fe²⁺_0.5Mg_0.5Cr₂O₄, Chr–Mg), 1.2–51 mol%; Ulvöspinel–Qandilite (FeMgTiO₄, Usp–Qan), 1.5–49 mol%; and Magnetite (Fe²⁺Fe³⁺₂O₄, Mag), 5.5–46 mol%. Ti-magnetite contains up to 5–7 wt% MgO in all samples, and its chemistry calculations fit the best four end-members: 4–24 mol% Hy, 31–71 mol% Ulvöspinel Fe₂TiO₄, up to 50 mol% Mag, and 2.5–37 mol% Magnesioferrite MgFe³⁺₂O₄.

Sample 08D96 encloses sporadic 400–500 µm resorbed xenocrysts of Cr-Sp containing 40–45 wt% Cr₂O₃, with Cr-bearing Ti-Mag rims. This and some other samples bear zoned or unzoned spinel microlites, both in the groundmass and as inclusions in phenocrysts of Cpx and olivine replaced by iddingsite aggregate. The contents of Cr₂O₃ in the spinel microlites are 25–30 wt% in the cores and 1–3 wt% in the rims of microlites; the respective contents of Al₂O₃, FeO, and TiO₂ are 22–26 wt%, 20–30 wt%, and 1–2 wt% (cores) and 5–6 wt%, 60–65 wt%, and 20–22 wt% (rims). Thus, the rim compositions of spinel microlites correspond to Cr-bearing Ti-Mag (Figure 8a,b). The most highly aluminous and magnesian spinels with Cr₂O₃ < 8 wt% (Cr# < 0.1) and TiO₂ < 2 wt% contents (Figure 8c,d) are common to the disseminated variety in the wehrlite xenoliths, as well as in pargasitic Amp and Cpx xenocrysts from 08D100-3 (Figure 4i–l). The spinel composition in mantle xenoliths and xenocrysts is 38–78 mol% Sp, 13–38 mol% Hy, 1.5–15 mol% Chr–Mag, 0–7 mol% Usp–Mag, 4–17 mol% Mag (Table S2). Cr-spinel lamelli exolved from wehrlite-hosted Cr-diopside (Figure 4i) contain 15–20 wt% Cr₂O₃ (up to 0.2 Cr#).

4.3. Minor Minerals, Accessories, and Glasses

The compositions of main accessories, rare minerals, and glasses are summarized in Table S3.

The matrix of samples 08D96 and 08D99-2B encloses Ilmenite with 1.2–4.6 wt% MgO; some Mg-rich ilmenite grains in 08D99-2B bear up to 0.9 wt% Nb₂O₅. Few ≤50–70 µm inclusions of Mg-rich Ilmenite with 6–8 wt% MgO were also found in wehrlite xenoliths and Cpx xenocrysts from basalt 08D100-3. The mineral chemistry was calculated for the Ilmenite FeTiO₃, Geikelite FeTiO₃, and Hematite Fe₂O₃ end-members (Table S3). The end-member percentages of Geikelite and Hematite increase, respectively, from 9.1–20 and 0–3.4 mol % in Mg-rich ilmenite from 08D96 and 08D99-2B to 26–29 and 4.6–5.9 mol% in that from basalt 08D100-3 (Table S3).

Nepheline was found to be filling interstices between microlites or entering aggregates that replace pargasitic Amp xenocrysts in 08D121-5 and 08D121-8. The mineral bears 5–10 mol% of kalsilite and has narrow ranges of element contents, including 0.02–0.03 apfu Ca and 0.01–0.04 apfu Fe³⁺. It occurs as microlites in assemblages with Ti-bearing Cpx + Pl ± (olivine, nepheline) and remnant glass, which replace Amp xenocrysts (Figure 2e,f and Figure 4g,h). All compositions fall within the field of rhönite in the Fe²⁺ vs. Mg diagram, at X_Mg, Mg/(Mg + Fe²⁺) = 0.5–0.7. The contents of Ti are from 1.8 to 2.6 apfu Ti (Table S3).

Biotite forms acicular microlites in 08D96 and 08D99-2B. Most of biotites are Ti-rich (0.2–0.6 apfu) with variable Fe/Mg ratios and Mg# = 0.3–0.6; some microlites contain 1–2 wt% F. The total of Na and K cations is 0.7–1.1 apfu, at K/Na = 3–8 (Table S3). Biotite in 08D96 bears 0.8–2.6 wt% BaO.

Few inclusions of Ti-bearing phlogopite (Mg# = 0.7–0.85, 0.10–0.35 apfu Ti), occasionally intergrown with pargasitic Amp, occur in Cpx xenocrysts (08D81 and 08D100-3) and in wehrlite xenoliths (08D100-3, Figure 4l); phlogopite contains up to 0.6–0.8 wt% BaO (Table S3).

Fluorapatite occurs as phenocrysts (rarely, Figure 4b) or microlites (more frequently). It has a stoichiometric composition, with 2–6 wt% F and 0.2–0.6 wt% Cl. Some microlites are relatively enriched with SrO (1–2 wt%) in 08D99-2B and 08D121-8 or with Ce₂O₃ (1–2 wt%) and La₂O₃ (0.7–1.2 wt%) in 08D81, 08TR02, 08MTP03, and 08D123-1/5 (Table S3). Titanite with Ce, Nd, Y, Nb, and Zr ≤ 3 wt% in total coexists with Cpx in agglomerations of Ti-Amp phenocrysts in sample 08D99-2B (Figure 2b and Figure 4f). Few rocks enclose fine grains (3–5 μm) of monazite-Ce and monazite-La (08MTP03), and Sr-enriched barite (08D133A, 08D100-3 and 08D121-8). Phillipsite often fills vugs and small veins in the groundmass. Analcime is restricted to vugs in 08MTP03 (Figure 2a) and 08TR02. Stoichiometric authigenic calcite, free from impurities, fills vugs and veins in samples 08D81, 08D133A, and 08D99-2B (Figure 2b,d).

Volcanic glasses rarely elude seawater alteration and are restricted to trachybasalt samples 08D133A and 08D102-3 (Table S3). They contain numerous micrometer-size inclusions of dendritic Ti-Mag and fill up to 30 × 70 µm interstices between groundmass minerals. Glass from sample 08D102-3 plots in the fields of basalt, trachybasalt, and basanite in the TAS diagram (Figure 3); it is enriched in FeO (8–12 wt%) and P₂O₅ (3–5 wt%), and occasionally contains 0.3–0.6 wt% SO₃, 0.2–0.3 wt% Cl, and up to 1 wt% F. Glass in 08D133A has higher contents of SiO₂ and Na₂O + K₂O, which correspond to tephri-phonolite, foidite, and phonolite, but contain less FeO (5–10 wt%), P₂O₅ (1–1.7 wt%) impurities, and may bear 10–15 wt% H₂O, as estimated from the deficit of totals.

4.4. Trace Element Chemistry of Ti-Amphibole and Clinopyroxene Phenocrysts

Trace elements were determined in monofractions of Ti-Amp and Cpx phenocrysts selected under binocular from crushed samples 08D81, 08TR02, 08D99-2B, 08MTP03, and 08MTP03-1 (Table 2). Amphiboles have higher element contents than Cpx (especially LREE, Sr, and Ba) and show prominent Nb-Ta and Zr-Hf depletion in the PM-normalized patterns (Figure 9a,c).

The PM-normalized patterns of Ti-Amp show Ba enrichment and depletion in Th, U, Nb, Ta, Zr, Hf, and Y, and occasionally in Pb (Figure 9a), with the Nb-Ta and Zr-Hf depletion decreasing and REE enrichment (especially LREE) increasing (Figure 9c) in the rock series 08D81 → (08D99-2B, 08MTP03) → (08TR02, 08MTP03-1). Compared to these, the PM-normalized patterns of Cpx phenocrysts (Figure 9b) have a different shape and lower contents of trace elements, with Pb enrichment and Sr depletion, and REE (especially LREE) increasing in the rock series 08D99-2B → 08MTP03 → 08MTP03-1 (Figure 9d). Clinopyroxene from basalt 08D96 distant to the location of Amp-bearing basaltic rocks (Figure 1c) stands out against the others in moderate MREE enrichment in the chondrite-normalized REE pattern.

Table 2. ICP-MS trace element compositions (ppm) of Ti-amphibole and clinopyroxene phenocrysts.

	Ti-Amphibole					Clinopyroxene
	08D81	08D99-2B	08TR02	08MTP03	08MTP03-1	08D96	08D99-2B	08MTP03	08MTP03-1
Li	3.7	6.3	5.8	8.6	7.1	10	9.8	9.6	7.4
Be	1.1	1.5	1.9	1.5	2.1	0.83	0.69	1.0	0.92
Rb	16	14	16	15	16	7.0	6.9	10	10
Cs	0.13	0.11	0.06	0.05	0.21	0.17	0.16	0.11	0.14
Ba	752	488	964	782	916	144	60	92	65
Sr	786	535	836	750	810	200	100	145	139
Zr	41	154	168	122	215	111	76	137	112
Hf	1.6	5.2	5.9	4.2	7.1	4.4	2.7	4.8	4.2
Ta	0.04	0.45	1.52	0.39	1.2	0.4	0.25	0.64	0.57
Nb	0.86	8.2	34	8.8	26	4.4	2.9	9.3	8.5
Ni	93	76	61	84	86	151	82	75	115
Co	64	46	59	56	60	92	21	27	32
Cr	69	182	69	110	111	2295	562	378	637
V	43	113	149	85	149	180	112	184	196
Ga	17	16	21	19	24	13	8.0	10	10
Ge	0.13	0.37	0.51	0.30	0.51	0.95	0.62	1.2	1.2
Sc	63	37	31	40	40	62	50	57	70
Cu	13	17	17	21	19	22	11	17	19
Zn	68	86	135	104	148	73	33	55	49
Mo	0.08	0.14	0.27	0.32	0.33	0.66	0.08	0.19	0.38
Sn	bd	0.11	0.47	bd	0.23	0.49	0.22	0.68	0.52
Sb	bd	0.02	0.03	bd	0.01	0.12	0.03	0.09	0.25
W	bd	0.20	0.71	0.02	0.06	0.39	0.03	0.05	0.07
Pb	2.5	10	3.5	6.7	3.7	7.4	1.8	3.1	5.8
Th	0.50	1.2	1.2	1.0	2.1	0.58	0.82	1.5	2.9
U	0.17	0.17	0.23	0.15	0.26	0.12	0.12	0.12	0.12
Y	25	23	32	28	37	17	13	18	17
La	18	31	51	39	71	10	10	21	33
Ce	52	80	127	103	160	31	27	54	101
Pr	7.5	10	16	13	20	4.7	3.8	6.4	7.4
Nd	31	35	52	43	61	19	15	21	23
Sm	7.3	7.8	11	9.4	12	5.8	3.6	5.1	5.2
Eu	2.3	2.2	3.2	2.8	3.5	1.8	1.1	1.4	1.5
Gd	7.1	7.2	10	8.6	11	5.7	3.5	4.7	5.1
Tb	1.0	0.94	1.3	1.1	1.5	0.78	0.5	0.67	0.68
Dy	5.5	5.3	7.2	6.3	8.3	4.3	2.9	3.8	3.9
Ho	1.0	0.95	1.3	1.2	1.5	0.73	0.5	0.72	0.71
Er	2.5	2.5	3.4	3.1	3.9	1.8	1.4	1.9	1.9
Tm	0.32	0.33	0.46	0.40	0.51	0.22	0.19	0.27	0.26
Yb	1.9	2.0	2.9	2.4	3.3	1.3	1.2	1.7	1.6
Lu	0.27	0.31	0.44	0.36	0.49	0.18	0.17	0.27	0.24

bd = concentration below ICP-MS detection limit.

Table 2. ICP-MS trace element compositions (ppm) of Ti-amphibole and clinopyroxene phenocrysts.

	Ti-Amphibole					Clinopyroxene
	08D81	08D99-2B	08TR02	08MTP03	08MTP03-1	08D96	08D99-2B	08MTP03	08MTP03-1
Li	3.7	6.3	5.8	8.6	7.1	10	9.8	9.6	7.4
Be	1.1	1.5	1.9	1.5	2.1	0.83	0.69	1.0	0.92
Rb	16	14	16	15	16	7.0	6.9	10	10
Cs	0.13	0.11	0.06	0.05	0.21	0.17	0.16	0.11	0.14
Ba	752	488	964	782	916	144	60	92	65
Sr	786	535	836	750	810	200	100	145	139
Zr	41	154	168	122	215	111	76	137	112
Hf	1.6	5.2	5.9	4.2	7.1	4.4	2.7	4.8	4.2
Ta	0.04	0.45	1.52	0.39	1.2	0.4	0.25	0.64	0.57
Nb	0.86	8.2	34	8.8	26	4.4	2.9	9.3	8.5
Ni	93	76	61	84	86	151	82	75	115
Co	64	46	59	56	60	92	21	27	32
Cr	69	182	69	110	111	2295	562	378	637
V	43	113	149	85	149	180	112	184	196
Ga	17	16	21	19	24	13	8.0	10	10
Ge	0.13	0.37	0.51	0.30	0.51	0.95	0.62	1.2	1.2
Sc	63	37	31	40	40	62	50	57	70
Cu	13	17	17	21	19	22	11	17	19
Zn	68	86	135	104	148	73	33	55	49
Mo	0.08	0.14	0.27	0.32	0.33	0.66	0.08	0.19	0.38
Sn	bd	0.11	0.47	bd	0.23	0.49	0.22	0.68	0.52
Sb	bd	0.02	0.03	bd	0.01	0.12	0.03	0.09	0.25
W	bd	0.20	0.71	0.02	0.06	0.39	0.03	0.05	0.07
Pb	2.5	10	3.5	6.7	3.7	7.4	1.8	3.1	5.8
Th	0.50	1.2	1.2	1.0	2.1	0.58	0.82	1.5	2.9
U	0.17	0.17	0.23	0.15	0.26	0.12	0.12	0.12	0.12
Y	25	23	32	28	37	17	13	18	17
La	18	31	51	39	71	10	10	21	33
Ce	52	80	127	103	160	31	27	54	101
Pr	7.5	10	16	13	20	4.7	3.8	6.4	7.4
Nd	31	35	52	43	61	19	15	21	23
Sm	7.3	7.8	11	9.4	12	5.8	3.6	5.1	5.2
Eu	2.3	2.2	3.2	2.8	3.5	1.8	1.1	1.4	1.5
Gd	7.1	7.2	10	8.6	11	5.7	3.5	4.7	5.1
Tb	1.0	0.94	1.3	1.1	1.5	0.78	0.5	0.67	0.68
Dy	5.5	5.3	7.2	6.3	8.3	4.3	2.9	3.8	3.9
Ho	1.0	0.95	1.3	1.2	1.5	0.73	0.5	0.72	0.71
Er	2.5	2.5	3.4	3.1	3.9	1.8	1.4	1.9	1.9
Tm	0.32	0.33	0.46	0.40	0.51	0.22	0.19	0.27	0.26
Yb	1.9	2.0	2.9	2.4	3.3	1.3	1.2	1.7	1.6
Lu	0.27	0.31	0.44	0.36	0.49	0.18	0.17	0.27	0.24

bd = concentration below ICP-MS detection limit.

5. Discussion

5.1. Fractionation of OIB Melts

Basaltic rocks with Ti-Amp phenocrysts and microlites are restricted to a relatively small area (~20 km²) of Govorov Guyot, while Amp xenocrysts occur in volcanics from different parts of the guyot and its southeast satellite (Figure 1c,d). According to our geochronological data, the ⁴⁰Ar-³⁹Ar ages for the Govorov volcanics are the oldest for basalt 08D100-3 (121.2 ± 2.8 Ma) and the youngest for Amp-bearing rocks: trachybasalt 08MTP03 (106.4 ± 2 Ma), basaltic trachyandesite 08D99-2B (101.9 ± 1.5 Ma), and basalt 08D96 (98.5 ± 1.4 Ma) ([20,21], and unpublished data). Judging by the Ar–Ar ages and similarity in textures and mineralogy, the Amp-bearing rocks formed by eruptions of compositionally uniform basaltic magmas in the 106–99 Ma interval from closely spaced volcanoes on the northeastern slope of Govorov Guyot (Figure 1c). We modeled the fractionation of mafic silicate melts in Amp-bearing magmas with reference to the trace element compositions of Ti-Amp and Cpx phenocrysts.

According to experimental evidence, calcic Amp crystallizes as the first liquidus phase from hydrous trachybasaltic melts at pressures from 0.6 to 2.6 GPa [33], which is confirmed by rock textures in our case (Figure 2a–d). The partitioning of trace elements (including REEs) between Amp and melt (^Amp/LiquidD_REE) depends on melt composition, temperature, and pressure. With increasing pressure at a constant temperature, LREEs in Amp become progressively more incompatible while ^Amp/LiquidD_REE decreases [33].

The REE patterns of H₂O-rich basaltic melts equilibrated with Ti-Amp were modeled with reference to ^Amp/LiquidD_REE for basanitic (500 MPa, 1000 °C, from Adam and Green [34]), trachybasaltic (600 MPa, 970 °C, after [33]), and basaltic trachyandesite (700 MPa, 1010 °C, from Nandedkar et al. [35]) melt compositions. At 500–700 MPa, the chondrite-normalized REE patterns of the melts equilibrated with Ti-Amp were similar to those for 08D81, 08D99-2B, 08MTP03, and 08MTP03-1 (Figure 9e–g). The respective modeling for Cpx was performed using empirical ^Cpx/LiquidD_REE calibrated for the Lousville trail magmatic system at several Pacific seamounts [36]. The chondrite-normalized REE patterns of the melts were also similar to those of Amp-rich basaltic rocks (Figure 9h). As predicted by REE modeling, the Ti-Amp phenocrysts may have crystallized within 500–700 MPa.

The PM-normalized patterns of Ti-Amp in trachybasalt 08D81 show the lowest LREE and prominent depletion in Nb-Ta and Zr-Hf (Figure 9a,c). The rock is enriched in P₂O₅ (0.74 wt%) and encloses numerous microlites and sporadic phenocrysts of fluorapatite containing 2 wt% LREE often intergrown with Ti-Amp and Cpx phenocrysts (Figure 4b). The REE features of Ti-Amp from 08D81 may be due to its crystallization jointly with LREE-bearing fluorapatite. The Nb and Zr depletion of Ti-Amp can also result from crystallization of liquidus phases enriched in these elements. For instance, Cpx in agglomerations with Ti-Amp phenocrysts enclose Nb-bearing ilmenite and titanite with Nb and Zr impurity (sample 08D99-2B, Figure 2b and Figure 4f; Table S2). Thus, the systematic changes in trace element contents (including REE) in Ti-Amp and Cpx phenocrysts (Figure 9c,d) may be due to crystallinity variations (phenocrysts/melt ratio) of the hydrous basaltic melts in different magma sources, as well as to the liquidus crystallization of LREE-bearing fluorapatite and to gravity precipitation of Ti-Amp and Cpx phenocrysts containing inclusions of ilmenite and titanite.

5.2. P-T Equilibration of Minerals

5.2.1. Mantle Rocks

The interactions with silicate melts/fluids percolating through peridotite can lead to the replacement of orthopyroxene and the formation of the wehrlite mineral assemblage (olivine + Cpx) ([39], and references therein). The peridotite exposed to the effect of percolating high-temperature hydrous silicate mafic melts apparently experienced partial melting and metasomatism that led to wehrlitization and to the crystallization of pargasitic Amp in wehrlite peridotite, as well as Amp–Cpx metasomatic mantle veins at different depths of lithosphere [40].

It is impossible to estimate the metasomatic P-T conditions of peridotite wehrlitization because all olivine has been replaced by iddingsite aggregate while no orthopyroxene coexists with Cpx. The crystallization parameters for calcic Amp in silicate mafic melts can be calculated from empirical equations in [27,28] at the following boundary conditions: P (0.4–2.2 GPa), T (800–1130 °C), ΔNNO (from −2.1 to +3.6 log units Ni-NiO buffer), and H₂O (2.8–12.2 wt%). From the pressure values, it is also possible to calculate magma/melt depth (km) using density–depth models (equation number 33 in [29], and in [41]).

According to the calculations, pargasitic Amp in wehrlite xenoliths crystallized at 2.5–1.6 GPa (73–53 km) and 1170–1130 °C (Figure 10). The mafic melts percolating through peridotite in equilibrium with pargasitic Amp contained 5.4–4.5 ± 0.6 wt% H₂O (Figure 11a). As for the metasomatic mantle veins containing intergrown coarse Cpx and pargasitic Amp, whose fragments were entrapped by ascending basaltic magmas and survived as xenoliths or xenocrysts in volcanic rocks (Figure 2i–k), they likely crystallized at lower P and T, from mafic melts containing up to 6 wt% H₂O: 2.1–1.6 GPa, 1100–1050 °C, and 65–50 km for inclusions in coarse Cpx xenocrysts; 1.6–1.3 GPa, 1070–1040 °C, and 52–44 km for the Cpx – Amp xenolith (metasomatic vein fragment); 1.5–1.1 GPa, 1040–980 °C, and 49–38 km for pargasitic Amp xenocrysts. The lowest estimates of pressure 700–600 MPa and depth 25–21 km at 1060–1020 °C were obtained for inclusions of pargasitic Amp intergrown with Ti-phlogopite in Cpx xenocrysts (Figure 10 and Figure 11a).

Pargasitic Amp in mantle rocks could form either concurrently with or after the wehrlitization of peridotite. The P-T estimates for the crystallization of average pargasitic Amp composition in wehrlite xenoliths (1.99 GPa and 1149 °C, Table 1) are close to the pressure inferred from other wehrlite minerals: T-dependent barometers for Cpx (1.87 GPa, after Nimis and Taylor [42]) and Cr-in-Cpx (2.04 GPa, after Sudholz et al. [43], but the temperature is higher than that according to the P-dependent Cpx thermometer (1002 °C, after [42]).

Spinel in wehrlite contains more magnesian (Mg# = 0.70–0.64) and less chromium (Cr# ≤ 0.10) than that from peridotites of the oceanic and continental lithosphere (Figure 8c). As estimated according to Equation F,% = 10Ln(Cr#) + 24 for peridotite melt fraction [32], wehrlite has no evident signatures of melting. The unusually low value of F < 1% is inconsistent with the texture of wehrlite xenoliths composed of euhedral olivine enclosing round spinel grains (Figure 2h and Figure 4i,k). This texture apparently records the crystallization of mantle minerals in a melt after the partial melting of peridotite. For this reason, Cr contents in spinel, which has undergone significant composition changes, make poor reference in the calculations of peridotite melting degrees.

The crystallization temperatures of pargasitic Amp in wehrlite xenoliths are around the pargasite dehydration solidus (1100–1130 °C), sourced from Green et al. [44] (Figure 10). The 30–40 °C excess above the solidus may be due to overestimated P-T values obtained by equations in [27,28], or to changes in the pargasite stability field as a function of its composition, water contents in the lithosphere, and other factors [45]. The Govorov Guyot metasomatized mantle rocks formed at P-T conditions corresponding to the spinel facies of peridotite. The findings of pargasitic Amp in wehrlite xenoliths and xenocrysts, as well as inclusions of Ti-phlogopite in Cpx xenocrysts, provide reliable evidence for metasomatism of the oceanic lithosphere beneath the guyot. The metasomatic agents (silicate melts and/or fluids enriched in volatiles) that affected the oceanic lithosphere and led to its partial melting may have originated either from a plume (SOPITA hotspot) or from the lithosphere–asthenosphere boundary [21].

5.2.2. Basaltic Rocks

According to equations in [27,28], most of phenocrysts containing >40 mol% Mg-hastingsite end-member in Amp-bearing basaltic rocks formed at 1060 to 910 °C within large pressure and depth ranges: 1.2 to 0.4 GPa and 40 to 15 km, respectively (Figure 10). Rim zones of kaersutitic microlites in 08D96 basalt, with the lowest percentages of Mg-hastingsite (<20 mol%), crystallized at higher temperatures than Ti-Amp phenocrysts and the core zones of microlites: 1100–1020 °C against 1050–980 °C, respectively. The crystallization temperatures of Ti-Amp in microlites are also higher than those for phenocrysts in trachybasalt 08D81: 1040–1020 °C against 1020–930 °C. As pressure decreased, the H₂O contents reduced from 8.6 ± 1.2 wt% to 2.6 ± 0.4 wt% while Ti-Amp was crystallizing from mafic silicate melts during the formation of basaltic rocks (Figure 11b). The highest and lowest values are, respectively, for the cores of phenocrysts and microlites; the latter formed at 4–2 wt% lower H₂O in the equilibrium melts.

The equation for ΔNNO in [28] is mostly valid for magmatic calcic Amp that crystallized in the crust and upper mantle because of its behavior and uncertainty at deep mantle conditions (>700 MPa) are hard to predict. Therefore, it was possible to evaluate fO₂ during Ti-Amp crystallization only in basaltic rocks. It was inferred to be above the Ni-NiO buffer (Figure 11c), with the lowest fO₂ values of about Ni-NiO and QFM buffers for xenocrysts in 08D121-5 and microlites in 08D96.

The P-T conditions of Cpx crystallization were calculated by the method of Putirka [46] using the equation for temperature and the updated pressure determination according to Neave and Putirka [47]. The calculations were performed with average compositions of Cpx phenocrysts and microlites, assuming that melts equilibrated with Cpx had LOI-free compositions. Phenocrysts of Mg-rich Cpx (Mg# = 0.7–0.8) in all Amp-bearing basaltic rocks crystallized at 1190–1080 °C and 1050–850 MPa. Microlites formed at 700–400 MPa and at temperatures 30–50 °C lower than for phenocrysts. The P-T estimates from Cpx are very approximate, as neither the difference between the compositions of melts (primary and residual) during the formation of phenocrysts and microlites, nor the effect of pervasive seawater alteration on the bulk-rock compositions were taken into account.

The large variations in P-T and chemometric parameters during the crystallization of Ti-Amp and Cpx may have resulted from magma mixing or degassing of basaltic melts. The local temperature variations associated with degassing of hydrous melts and with significant H₂O decrease during rapid magma ascent may be responsible for Mg–Fe zonation in the Ti-Amp and Cpx phenocrysts.

The P-T estimates of partial or complete replacement of Ti-bearing Amp xenocrysts (pargasitic Amp) or phenocrysts (Ti-bearing Mg-hastingsite) can be estimated from experimental data. According to the liquidus–solidus diagram for basaltic melt, the assemblage of Cpx + (Ti-Mag or rhönite) + Pl + olivine substitutes for Ti-Amp is in a small P-T region at 1120–1060 °C and 70–10 MPa [48]. Such P-T conditions may arise as high-temperature magmas rise to intermediate chambers at depths from ~3 km (70 MPa) to ~1 km (10 MPa). Ti-bearing Amp in xenocrysts or phenocrysts could elude alteration only if basaltic magmas rose very rapidly from the ~3 km depth level and erupted on the guyot surface.

6. Conclusions

The volcanic rocks of Govorov Guyot (basalts, trachybasalts, basaltic trachyandesites, and trachyandesite) consist of rock-forming (Ti-Amp, Cpx, Pl, olivine fully or partially replaced by iddingsite aggregate, oxyspinel group Fe and Ti oxides), minor and accessory (nepheline, rhönite, biotite, fluorapatite, etc.) minerals. Some rock samples bear coarse pargasitic Amp and Cpx xenocrysts with ± Ti-phlogopite inclusions, as well as xenoliths of metasomatic mantle vein fragments (large Cpx + Amp intergrowths), orthopyroxene-free wehrlite xenoliths with fully replaced olivine, disseminated inclusions of spinel and pargasitic Amp, which represent remnant metasomatized oceanic lithosphere beneath Govorov Guyot.

Ti-bearing amphiboles make up a continuous series from pargasite – Mg-hastingsite in wehrlite xenoliths and xenocrysts to Mg-hastingsite–kaersutite (ferri-kaersutite, ferro–ferri-kaersutite, and ferro-kaersutite) end-members in phenocrysts and microlites of basaltic rocks.

The crystallization conditions of Amp in basaltic melts, wehrlite peridotite, and metasomatic mantle veins were reconstructed using equations from [27,28]. The inferred crystallization conditions for pargasitic Amp were: 2.5–1.6 GPa (73–53 km depth) and 1170–1130 °C in wehrlite; 1.6–0.6 GPa (50–21 km) and 1070–980 °C in metasomatic mantle veins. Pargasitic Amp may have crystallized at P-T conditions corresponding to the spinel facies of peridotite at different depths, from hydrous (6.0–4.5 wt% H₂O) mafic silicate melts that percolated through metasomatized peridotite of the oceanic lithosphere beneath Govorov Guyot.

The systematic changes in trace element contents of Ti-Amp and Cpx phenocrysts from basaltic rocks may be due to variations of the phenocrysts/melt ratio in different hydrous magma sources, as well as to the liquidus crystallization of LREE-bearing fluorapatite and to gravity precipitation of Ti-Amp and Cpx phenocrysts containing inclusions of ilmenite and titanite. The mafic melts lost H₂O (decrease from 8.6 to 2.6 wt%) as they moved to lower-pressure conditions, while Ti-Amp phenocrysts and microlites crystallized from the parent melts of basaltic rocks at 1.2–0.4 GPa (40–15 km), and 1060–910 °C.