RUSSIAN JOURNAL OF EARTH SCIENCES VOL. 7, ES4001, doi:10.2205/2005ES000181, 2005
[10] The rock samples discussed in this paper were dredged mostly from the slopes of the Markov depression (sites I1032, I1060, I1063, and I1069), and basalts were uplifted from a volcanic plateaus north of it (sites I1026, I1027, and I1072). The dredged material included both fresh and variably altered rock material. The completely metamorphosed rocks were classified into (a) those developing after ultramafic rocks, gabbro, and dolerites and (b) metamorphic rocks.
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Figure 3 |
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Figure 4 |
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Figure 5 |
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Figure 6 |
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Figure 7 |
[12] The dredged dolerite samples are fairly diverse: they include common olivine-clinopyroxene and olivine-free varieties, sometimes with elevated contents of ilmenite (up to 3 vol %), and plagioclase-phyric Ilm- Hbl dolerites. Along with fresh rocks, the dredged material contained altered varieties, whose pyroxenes and hornblende are partly or completely replaced by fibrous actinolite.
[13] Both the fresh and the altered basalts can be aphyric or porphyritic, with phenocrysts of Ol, Ol+Pl, or Pl alone. The groundmass of these basalts is usually ophitic or has an unusual texture with sheaves of recrystallized spherulites, which are clearly seen in glassy varieties. The latter type of the basalts usually bears abundant small ilmenite grains, a feature suggesting that these rocks are volcanic analogues of the second group of the plutonic rocks. The rocks often exhibit traces of high-temperature cataclasis with the development of deformation structures in magmatic minerals (Ol, Opx, Pl, and Hbl)1 and the appearance of small subequant neoblasts. The genesis of such high-temperature cataclasites was likely related to deformations in already solid but still hot rocks, because the minerals show no traces of low-temperature alterations, and the composition of the neoblasts is close to that of the cataclasts.
[14] Rocks of particular interest are the volcanic breccias that were cataclased under such conditions and were found among rocks from sites I1060 and I1063. Boudin-, oval-, and lens-shaped fragments (15-20 cm long) in these rocks consist of cataclased peridotites with relict cumulus textures (harzburgites and lherzolites) that were not affected by low-temperature alterations and are cemented by weakly cataclased fine-grained porphyritic rocks, whose composition varies from gabbronorite-diorite to granodiorite with ilmenite and magmatic kaersutite. It is worth noting that fragments (boudins) of these ultramafics bear practically no traces of low-temperature alterations at contacts with the diorite-granodiorite cement, as is also typical of relations between syngenetic vein derivatives and their host cumulates.
[15] The low-temperature alterations are spread much more broadly: the peridotites are usually extensively serpentinized, and the gabbroids are amphibolized with the development of fibrous actinolite after pyroxenes and pargasite. The rocks are commonly cut by veinlets of carbonate, prehnite and chlorite. In places, the rocks were also affected by low-temperature shearing and brecciation. The thickest of these zones are associated with the development of diverse metamorphic rocks, which often have disseminated-stringer or massive sulfide mineralization [Pushcharovskii et al., 2002].
[16] In order to characterize the typical varieties of the rocks, we conducted their detailed petrological and geochemical examination, whose results are summarized below. The petrography of the rocks is briefly characterized in Table 1, and microprobe analyses of their minerals are presented in Tables 2, 3, 4, 5, 6, and 7.
[18] The composition of olivine in the rocks broadly varies from Fo91-85 in the peridotites and troctolites to Fo26 in the olivine gabbro-diorites (Table 2). A sample of the Fe-Ti-oxide olivine dolerite contains zonal olivine crystals with Fo82 in the cores and Fo62 in their peripheral portions. The composition of the olivine microphenocrysts in olivine plagioclase-phyric basalt I1072/1 corresponds to Fo87, i.e., is close to the composition of olivine in the primitive troctolite. An unusual feature of olivine in the peridotites with relics of cumulate textures is its variable composition: the predominant grains have the composition Fo86, and occasional grains correspond to Fo70 (sample I1063/17, Table 2).
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Figure 8 |
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Figure 9 |
[21] Magmatic kaersutitic brown hornblende occurs in the gabbroids both as a cumulus mineral and as intercumulus grains, and the dolerites contain small kaersutite grains in the groundmass. Similarly to the pyroxenes, the hornblende is often replaced by fibrous actinolite. The composition of the hornblende varies from high-Ti varieties in the hornblende gabbroids to relatively low-Ti varieties in the diorites and quartz diorites (Table 5). According to the results of Pl- Am thermometry [Blundy and Holland, 1990], the rock crystallized at temperatures from 800o C (Fe-Ti-oxide hornblende gabbronorite L1124/13-11) to 650-580o C Fe-Ti-oxide gabbro-diorite and quartz diorite) (unpublished data of S. S. Abramov).
[22] The apatite contains small amounts of Cl (0.09-0.24 wt%). Fluorine was not determined in this mineral because of methodical difficulties.
[23] The oxide minerals (Cr-spinel, titanomagnetite, and ilmenite) are particularly interesting as exhibiting some noteworthy structural features.
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Figure 10 |
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Figure 11 |
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Figure 12 |
[26] The composition of most chromite grains in this thin section and the chemistry of the vermicular aggregates among olivine and pyroxene neoblasts, as well as the composition of chromite in the rim of the zonal crystal, are characterized by very low Fe3+ contents and variable Cr/Al ratios. As in this rock, Cr-spinels in harzburgite I1063/2 can be subdivided into two groups: one close to the predominant type of grains in sample I1063/17, and the other characterized by low Al2O3 contents and high concentrations of Cr2O3.
[27] All of these data evidently indicate that the early Fe-rich chromite with elevated TiO2 and relatively low Al2O3 and MgO contents are xenogenic for these peridotites. They could be preserved as inclusions in other minerals (when captured by growing crystals), but more commonly they reacted with the magmatic melt and were dissolved in it. Likely evidence of this processes is displayed in Figure 12.
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Figure 13 |
[29] As obviously follows from all of these data, the rocks contain minerals belonging to at least two distinct assemblages. Relics of the earlier of them are contained as inclusions ("seeds'') in the typical minerals of the magmatic assemblage itself. These are Fe-rich Cht overgrown by rims of a more magnesian and less titanian variety in the peridotite (sample I1063/17) and Ti- Mag with exsolution lamellae in clinopyroxene from the gabbronorite (sample I1063/1). Moreover, the composition of the olivine significantly varies, and volumetrically predominant magnesian varieties are in places associated with moderately magnesian grains. All of these facts testify to very unusual conditions under which the rocks were produced.
[30] A noteworthy feature of the rocks is the variable composition of the same minerals in them. This can be illustrated most glaringly by the example of the peridotites wit relict cumulus textures: a single sample of these rocks can contain olivine of the composition Fo86 and Fo70 (Table 2, sample I1063/17). The broadest compositional variations in this sample and in I1063/2 are exhibited by their Cr-spinel (Table 6, Figure 10). Some of their grains are Fe-rich, which is atypical of mantle rocks but is characteristic of ultramafic layered complexes [Sharkov et al., 2001]. Moreover, some of the chromite grains contain corroded cores of Fe-rich chromite overgrown by more magnesian outer zones, which contain P and Cl (Figure 11, Table 6).
[31] As can be seen from Table 8, the Fe-Ti-oxide gabbronorites
crystallized, according to Ilm- Mag geothermometry, at average
temperatures of 621
41o C, which is close to the analogous values
obtained by the same method for Fe-Ti-oxide gabbroids in the
Southwest Indian Ridge
[Natland et al., 1991].
[33] Analogous rocks were dredged from the walls of other parts of the Markov depression and neighboring depressions in this MAR segment, which testifies to their broad in-situ occurrence. Zircon grains were obtained from sample I1028/1 using magnetic separation and heavy liquids, and by hand-picking under a binocular magnifier. The grains were classified according to their size, color, and morphology. Approximately 70-75% of each population were utilized to prepare epoxy pellets, which were polished and examined under an electron microscope with a cathode-luminescence detector. The rest of the populations were dated (TIMS) by the U-Pb method [Sharkov et al., 2004a].
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Figure 14 |
[35] The smaller fraction includes euhedral zircon grains with clearly pronounced dipyramid faces (Figure 14c). The predominance of euhedral grains was caused both by the crystallization of this zircon from a melt and by the fragmentation of larger grains during cataclasis. This follows form the fact that the zoning of some grains is not conformable with their morphology (Figure 2c), whose euhedral faceting seems to be secondary and imitates the habit of the larger primary grain.
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Figure 15 |
[37] Our experience in studying zircon from blastomylonitization zones in the Early Proterozoic Pezhostrov Island gabbro-anorthosite massif in the Belomorian Mobile Belt in the White Sea area [Sharkov et al., 1994] indicates that zircon recrystallization coupled with a diminish of its grain sizes does not affect the isotopic characteristics of this mineral, and its age evaluated using these grains is close to the age of zircon crystallization form the melt [Zinger et al., 2001]. This led us to believe that the Mesozoic age obtained for zircon from gabbronorite from the Markov depression corresponds to the crystallization age of this mineral.
[38] Finds of ancient zircon in gabbroids form the axial MAR zone put forth the problem of interpreting the nature of this mineral itself, as to whether it crystallized from a magmatic melt and, hence, its age corresponds to the crystallization age of the rock, or the zircon is a relict and xenogenic mineral that was not decomposed during the melting of the rock. According to currently adopted concepts, the opening of the Central Atlantic started at approximately 170 Ma and continues until now [Pushcharovskii, 2001]. This led Pilot et al. [1998] to hypothesize that zircon could be preserved in the mantle at temperatures of about 1000o C for 150 m.y. without losses of its radiogenic lead. These researchers proposed two possible explanations for the nature of ancient zircon in the gabbroids:
[39] (1) The rifting coupled with the opening of the Atlantic resulted in the tectonic disintegration and imbrication of the continental crust, which was broken into a series of tectonic slabs as a consequence of subhorizontal thrusting in the continental lithosphere. Involved in small-cell convection in the upper mantle, these continental lithospheric fragments were remelted. The convection cells that circulated on both sides of the ridge facilitated the displacement of these semi-molten zircon-bearing rocks toward the spreading axis, where they took part in the formation of the gabbroids.
[40] (2) The continental crustal material has been preserved in the Kane zone since the opening of the Atlantic because of the migration of the transform fault and jumps of the ridge axis. Part of this material later subsided in the axial zone of the ridge and was involved in the origin of the gabbroids.
[41] Thus, these researchers believe that the ancient zircon is of continental genesis, is not related to oceanic magma generation, and, hence, cannot provide information on the age of its host rocks. At the same time, geological, petrological, and isotopic-geochemical data provide no evidence that continental material can occur either in the Kane Fracture Zone or in the Markov depression. Moreover, the Fe-Ti-oxide silicic rocks found in the walls of the depression are typical of the third layer of slow-spreading ridges, such as MAR or the Southwest Indian Ridge, and of some ophiolitic associations, such as Monviso in the Western Alps [Lombardo et al., 2002; Rubatto and Gebaver, 2000]. These rocks typically bear zircon with analogous unusual sectorial zoning, which is uncommon in this mineral from continental rocks.
[42] These facts demonstrate that there is no need to propose any specific mechanism for the genesis of the zircon. Furthermore, according to experimental data, zircon can be easily dissolved in basaltic magmas and can crystallize only from melts oversaturated with silica [Watson and Harrison, 1983], and, thus, the occurrence of this mineral in rocks of the silicic Fe-Ti-oxide association can hardly be accidental. This led us to believe that the zircon crystallized from melts that were emplaced into the oceanic lithosphere, and the zircon age corresponds to the age of its host rocks.
Citation: 2005), Silicic Fe-Ti-oxide series of slow-spreading ridges: petrology, geochemistry, and genesis with reference to the Sierra Leone segment of the Mid-Atlantic Ridge axial zone at 6° N, Russ. J. Earth Sci., 7, ES4001, doi:10.2205/2005ES000181.
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