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RUSSIAN JOURNAL OF EARTH SCIENCES, VOL. 5, NO. 1, PAGES 1­29, FEBRUARY 2003

The comp ositional evolution of differentiated liquids from the Skaergaard Layered Series as determined by geo chemical thermometry
A. A. Ariskin
Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences

Abstract. Based on the COMAGMAT-3.65 crystallization model a set of phase equilibria calculations (called geochemical thermometry) have been conducted at P =1 kbar and closed conditions with respect to oxygen for 65 rocks representing the principle units of the Layered Series of the Skaergaard intrusion. It allowed us to define the range of initial temperatures (1145 to 1085 C) and oxygen fugacities (1­1.5 log units above QF M to slightly below QF M ) of the original crystal mush from which the rocks from LZa to UZa crystallized. In parallel, average ma jor-element compositions of residual (interstitial) liquid were calculated demonstrating a trend of continual enrichment of FeO (up to 18 wt.%) and TiO2 (up to 5.5 wt.%) with only minor variations in the SiO2 contents (48 to 50 wt.%). Pro jection of the compositions onto the OLI V -C P X -QT Z diagram provides evidence that most of the Layered Series crystallized on the Ol-P l-C px-oxide cotectic. Systematic differences between the calculated residual liquid compositions for LZa/LZb and LZc to UZa (which are unlikely to reflect fractional crystallization) are within the accuracy of the COMAGMAT model, but may be also indicative of a late-stage process involving migration and re-equilibration of interstitial liquids. Estimated amounts of interstitial melts trapped in the Skaergaard "cumulates" range around 50 wt.%. Wager's compositions inferred from simple mass-balance were found to lie too far from the OlP l-C px boundary to represent a realistic approximation of the low-pressure Skaergaard magma evolution. The main problem of genetic interpretations of the Skaergaard intrusion is a strong misbalance between the parental compositions followed from contact rocks and the results of geochemical thermometry and that of the whole differentiated body. It is most apparent for TiO2 and P2 O5 which are almost twice as high in the average intrusion composition compared to the proposed parents. Moreover, the intrusion composition has of 2­4 wt.% less SiO2 and much more iron. One possible explanation is to assume the Skaergaard magma came to the chamber with an amount of crystals (Ol + P l) equilibrated with the calculated parental liquid. However, even if some amount of "hidden" troctolitic material exists, it is unlikely that crystallization in a closed system could produce large volumes of rocks rich in Fe-Ti oxides without complementary more felsic differentiates.

1. Intro duction

Copyright 2003 by the Russian Journal of Earth Sciences. Paper number TJE03115. ISSN: 1681­1208 (online) The online version of this pap er was published 22 January 2003. URL: http://rjes.wdcb.ru/v05/tje03115/tje03115.htm

Investigations of differentiation of the Skaergaard magma have played a special role in igneous petrology by providing a canonical example of the evolution of a tholeiitic magma that had strong iron enrichment and depletion in SiO2 in the middle stages of differentiation. Excellent, unweathered exposures, detailed sampling, and favorable spatial relations 1


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ariskin: the compositional evolution of differentiated liquids

Figure 1. Magmatic differentiation trends proposed for the Skaergaard intrusion. "Natural" trends are based on mass-balance calculations [Hunter and Sparks, 1987; Wager and Brown, 1967], dike rock compositions [Brooks and Nielsen, 1978, 1990], and results of melting experiments [McBirney and Nakamura, 1973; McBirney and Naslund, 1990]. Zones and subzones are shown next to the relevant symbol: H Z , Hidden zone (after 35% assumed crystallization); LZ , lower zone (76%); M Z , middle zone (88.2%); U Z , upper zone (UZa, 95.7%; UZb, 98.2%; UZx c, 99.3%). Calculated crystallization trends represent phase equilibria modelling with COMAGMAT-3.5 at P = 1 atm and QF M [Ariskin, 1999]: T M is the initial "trapped" melt obtained at 1165 C as a result of geochemical thermometry of MBS rocks (Table 1); E G4507 and K T -39 are the chilled gabbros [Hoover, 1989; Wager and Brown, 1967].

between the main differentiated units have made it an ideal testing ground for a variety of petrologic concepts and techniques. Based on the premise that the initial Skaergaard magma was a pure liquid free of crystals, with the volumes of the differentiates assumed to be proportional to their thickness in the central part of the intrusion, Wager and Deer [1939] proposed a simple graphic technique for the calculation of the composition of the Skaergaard magma at successive stages of differentiation. These early estimates used primarily mass-balance relations between an initial composition inferred from a sample from the chilled margin and average compositions of the main units of the Layered Series. In order to achieve a satisfactory mass balance, Wager [1960] had to postulate a large "Hidden Layered Series". The rocks of the Upper and Marginal border series were not included in the calculations, with the Sandwich Horizon being taken to be the final product of differentiation. The result was a series of compositions that became progressively more iron- rich and silica-poor throughout almost the entire sequence of differentiation [Wager and Brown, 1967]. Silica increased measurably at the very end when it reached a maximum value of about 55 wt.% with 18.5 wt.% FeO in the extreme differenti-

ate (Figure 1). Subsequent investigations revealed two serious flaws in Wager's calculations. First, the E G4507sample of the chilled margin used for the initial magma was shown to be contaminated by metamorphic country rocks and therefore unrepresentative of the original liquid [Hoover, 1989; McBirney, 1975]. Second, a geophysical survey showed that the huge Hidden Zone required by the mass balance does not exist [Blank and Gettings, 1973]. These observations do not necessarily invalidate the main conclusion that the differentiated liquid became increasingly iron-rich throughout essentially the full range of differentiation, but one must take into account two additional circumstances. (1) The principal uncertainty in calculating Skaergaard liquid compositions by mass-balance is the proportions of the individual units of the Layered and Border Series. Although their thickness can be measured, it is difficult to determine their original lateral extent and volume: one can obtain a range of compositional trends simply by using different relative volumes. (2) Any trend calculated by mass-balance must be consistent with phase equilibria constraints if this trend is assumed to have resulted from magmatic fractionation controlled by crystalmelt equilibria. However, attempts to deduce the Skaergaard


ariskin: the compositional evolution of differentiated liquids Table 1. Compositions of chilled gabbros and trapped liquids proposed to be parental to the Skaergaard igneous rocks Sample Chilled gabbro E G4507 SiO2 TiO2 Al2 O3 FeOtot MnO MgO CaO Na2 O K2 O P2 O5 48.58 1.18 17.40 9.73 0.16 8.71 11.50 2.39 0.25 0.10 K T -39 50.46 2.70 13.41 12.96 0.22 6.71 10.34 2.41 0.57 0.22 LA-95 (LZa) 1150 C 51.44 1.61 13.03 13.59 0.26 6.15 11.18 2.38 0.28 0.09


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Trapped liquids E C -22 (MBS) 1160 C 51.30 1.61 12.99 13.55 0.26 6.13 11.15 2.37 0.28 0.09 T M (MBS) 1165 C 49.94 1.68 12.93 13.22 0.19 6.89 12.38 2.37 0.26 0.15

Note: Natural samples: E G-4507 [Wager and Brown, 1967], K T -39 [Hoover, 1989]. Exp eriments at QF M buffer: LA-95 denotes a glass after melting the ro ck from LZa [McBirney and Nakamura, 1973; McBirney and Naslund, 1990]; E C -22 denotes a glass after melting the rock from MBS [Hoover, 1989]. Phase equilibria calculations (geo chemical thermometry): T M represents an average for 6 liquid lines of descent intersected at 1165 C [Ariskin, 1999]. All comp ositions are normalized to 100 wt.%.

liquid compositions from mass-balance calculations based on phase equilibria constraints [Hunter and Sparks, 1987] and magma crystallization models [Ariskin, 1999; Ariskin et al., 1988] drew attention to serious contradictories of the empirically constructed trends which seemed to be inconsistent with experimentally-determined modal mineral proportions. Hunter and Sparks [1987] used a modified version of Wager's calculations. They emphasized that at 48.1 wt.% SiO2 in the sample E G4507 ( initial magma) the monotonous silica depletion trend can not be consistent with the low SiO2 contents (44­46%) of the observed "gabbroic assemblage" Ol-P l-Aug ± M t. Another discrepancy was that the composition at which the iron enrichment changed to silica enrichment did not coincide with the onset of Fe-Ti oxides crystallization in the rocks, which are known to appear quite early in the Skaergaard Layered Series [McBirney, 1989; Wager and Brown, 1967]. Following this reasoning, Hunter and Sparks [1987] argued that the evolution of the Skaergaard magma was similar to that seen in differentiated tholeiitic lavas of Iceland and western Scotland. These speculations were supported by simple mass-balance calculations, including subtraction of specified amounts of the Skaergaard gabbroic components from the composition of chilled gabbro K T -39 assumed to present more accurate the estimate of the Skaergaard parental magma [Hoover, 1989]. The rhyolitic end-product was said to have been erupted without leaving a trace in the presently exposed body of Skaergaard rocks. The geological aspect of the proposed interpretations found little support among petrologists who had worked on the intrusion and were familiar with the field relations [Brooks and Nielsen, 1990; McBirney and Naslund, 1990; Morse, 1990]. They ob jected that Hunter and Sparks had forced the Skaergaard liquids to follow the trend observed in volcanic rocks by assuming unrealistic volumetric proportions and ignoring well-documented geologic evidence. This was the main reason for rejecting the revised direction of the Skaergaard magma fractionation, although the phase equilibria problems were left unresolved [Ariskin, 1998; Hunter and Sparks, 1990].

1.1. Mo delling Fractional Crystallization

An indirect way to define compositional evolution of the Skaergaard magma is to calculate perfect fractionation tra jectories for a liquid thought to be parental to the whole intrusion. Recent calculations with the use of COMAGMAT-3.5 program including more accurate models for the simulation of M t and I lm crystallization [Ariskin and Barmina, 1999] confirme the conclusion [Ariskin et al., 1988; Toplis and Carrol l, 1996] that in systems open with respect to oxygen M t saturation leads to strongly decreasing iron and increasing silica contents of residual liquids, whereas closed systems crystallize lowered amounts of magnetite with a less pronounced iron depletion in the liquid. These simulations were conducted for two initial compositions ­ K T -39 [Hoover, 1989] and T M [Ariskin, 1999]. The latter is an average composition of six modelled melts assumed to be trapped in the most primitive rocks of the Marginal Border Series (Table 1). The phase equilibria calculations give evidence that both compositions are subcotectic (Ol + P l), with K T -39 slightly oversaturated with Ol, whereas T M is closer to the Ol + P l + Aug saturation. Liquid lines of descent representing the fractionation at 1 atm and QF M are plotted in Figure 1. Each of the initial compositions is shown by two evolutionary lines: the upper ones include calculations without any corrections in COMAGMAT-3.5, whereas the lower lines represent a corrected model, providing a maximally possible compositional shift based on the accuracy of the M t- and I lm-models. With these uncertainties [Ariskin, 1999] reached the following conclusions: (i) both parents exhibit only minor decrease in SiO2 as compared to the proposed FeO-SiO2 trend [Brooks and Nielsen, 1978, 1990; Wager and Brown, 1967]; (ii) all of the calculated tra jectories indicate a well-defined inflection point, where the content of SiO2 starts to increase due to the appearance of M t; (iii) the maximum possible iron enrichment of residual melts is approximately 18 wt.% FeO for K T -39 and 20% for T M .


4 ariskin: the compositional evolution of differentiated liquids

Figure 2. Example of geochemical thermometry for the Skaergaard MBS rocks. Lines correspond to the crystallization tra jectories calculated at 1 atm and QF M conditions for six contact gabbros; the temperature of is assumed to indicate the initial temperature for the trapped liquid [Ariskin, 1999]. Experimental data: circles liquids for the Marginal Border Series pristine rocks [Hoover, 1989]; diamonds are those obtained for the Lower Layered Series rocks [McBirney and Nakamura, 1973; McBirney and Naslund, 1990].

equilibrium 1165±10 C are residual and Middle


ariskin: the compositional evolution of differentiated liquids

5

Table 2. Average bulk-rock compositions of the principal units of the Layered Series and equivalent rocks of the Marginal Border Series [McBirney, 1996] Unit Zone SiO2 TiO2 Al2 O3 FeOtot MnO MgO CaO Na2 O K2 O P2 O5 LS LZa 48.12 1.35 16.81 11.13 0.16 9.42 10.11 2.52 0.27 0.11 MBS LZa


LS LZb 48.84 1.44 12.55 12.84 0.21 10.13 11.57 2.13 0.20 0.09

MBS LZb


LS LZc 41.10 6.92 11.02 21.10 0.26 7.61 9.77 1.97 0.20 0.05

MBS LZc


LS MZ 42.79 6.79 11.53 20.00 0.26 6.24 9.87 2.23 0.21 0.08

MBS MZ


LS UZa 43.07 5.67 11.17 22.52 0.31 5.62 8.62 2.55 0.26 0.22

MBS UZa


49.43 1.00 13.71 11.87 0.19 10.77 10.76 1.93 0.25 0.08

50.30 1.08 14.13 11.83 0.18 9.42 10.89 1.88 0.22 0.07

44.86 4.91 11.35 19.02 0.26 6.60 10.59 2.03 0.26 0.13

43.46 5.44 12.48 20.51 0.23 5.64 9.54 2.24 0.35 0.12

44.33 6.24 10.05 21.47 0.29 5.57 9.39 2.08 0.34 0.24

Note: LS is the Layered Series, MBS is the Marginal Border Series. Normalized to 100 wt.%.

1.2. Exp erimental Determination of Trapp ed Liquid Comp ositions Another approach is based on the premise that the Skaergaard gabbros retained a certain amount of trapped or residual liquid within the matrix of cumulus crystals. According to the widely accepted interpretation of layered gabbros as mixtures of the primary minerals and melts, one can estimate the amount of this interstitial component from the bulk-rock concentrations of incompatible elements [e.g., Chalokwu and Grant, 1987; Henderson, 1970]. By melting a sample to a temperature slightly above to its solidus the interstitial liquid can be restored and analyzed. When this method was used to determine the composition of trapped liquid in rocks of the Layered Series [McBirney and Nakamura, 1973], at oxygen fugacities between the QF M and W M buffers and the temperatures lower 1150 C it yielded a series of compositions that follow the "Wager trend" with a steady increase in iron and no pronounced increase of silica [McBirney and Naslund, 1990]. In the FeO-SiO2 diagram this experimental trend corresponds well to that followed from some co-magmatic dike compositions (Figure 1). Hoover [1989] performed several additional melting experiments on gabbros representing the earliest and least differentiated rocks crystallized within the Marginal Border Series. The partial melts obtained at 1160­1180 C and QF M conditions contained markedly more SiO2 and less FeO than the residual liquids observed from "the Layered Series" experiments. In part, the compositional discrepancy may be attributed to the effect of temperature supported by the results of modelling equilibrium crystallization for the chilled gabbro E G4507 (Figure 2). One can see, that at low temperatures the modelled E G4507 line resembles the silica depletion and iron enrichment trend, satisfying both of the above discussed experiments [Ariskin, 1999]. It is noteworthy that at the same temperature the chilled gabbro is characterized by silica contents 1.5 wt.% lower than partial melts of the most of the Marginal Border Series rocks or K T -39 and T M parents. This indicates the possibility that there were differences in the compositions of magmatic liquids parental to marginal rocks and those of presenting the Layered Series cumulates.

1.3. Calculations for the Marginal Border Series Utilizing the COMAGMAT-3.5 crystallization model and a modelling principle known as "Geochemical thermometry" (Appendix 1), the reconstruction of the intersitial liquid composition has been performed for six samples representing the least evolved part of the Marginal Border Series [Ariskin, 1999]. Five of these samples were collected within 1 to 8.5 m of the intrusive contact [Hoover, 1989], whereas the sixth was the "chilled marginal gabbro" E G4507 [Wager and Brown, 1967]. Calculations of equilibrium crystallization at 1 atm pressure, dry conditions and QF M buffer indicated that the mineral crystallization temperatures are correlated with the starting compositions, demonstrating a wide temperature field of Ol for high-magnesia samples and an early crystallization of P l for a plagioclase-rich rock. Two samples (including E G4507) showed sub-cotectic (P l + Ol) relations in the range 1230­1250 C. In accordance with the sequence of crystallization, the calculated lines of descent yielded an obvious intersection in the range of 1175­1155 C (Figure 2). The temperature of 1165±10 C was accepted as representing the initial temperature of the Skaergaard parental magma. The average liquid composition representing this cluster of six evolutionary lines at 1165 C is given in Table 1. It is close to liquid compositions that Hoover [1989] obtained in his melting experiments with marginal rocks. In this paper, estimates of ma jor-element geochemistry for "residual liquids" and primary mineral proportions for the LZ, MZ, and UZ rocks are first presented.

2. Calculations for the Main Skaergaard Units
Before carrying out final geochemical thermometry calculations and a systematic analysis of the modelled liquid lines of descent for real rocks, a set of preliminary simulations of equilibrium crystallization was conducted for the average Layered and Marginal Border Series rocks (Table 2). The Layered Series (LS) of the Skaergaard intrusion represents


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ariskin: the compositional evolution of differentiated liquids

Figure 3. The main units of the Layered Series and the Upper Border Series [McBirney, 1996]. The LS zones are defined by the appearance or disappearance of primary phases, or, in the case of LZb, by the transition from the poikilitic form of the Ca-pyroxene to the granular one. The Upper Border Series is divided into three subzones (, , and ) equivalent to the LZ, MZ and UZ. The compositions given for Ol, P l, and Aug at zone boundaries are approximate and may vary laterally.

a sequence of rocks that formed from the floor upward. It is lithologically and structurally distinct from the Marginal Border Series (MBS) that crystallized inward from the walls and the Upper Border Series (UBS) that crystallized downward from the roof. The upper boundary of the LS is defined by its contact with the UBS along a level known as the Sandwich Horizon ­ SH [Wager and Deer, 1939]. As a rule, this level is recognizable in the field by the contrast between the mafic ferrodiorites of the LS and more felsic UBS [McBirney, 1989]. In accordance with the nomenclature established by Wager and Brown [1967], the Layered Series is divided into Lower, Midd le, and Upper Zones by the disappearance of abundant primary olivine at the base of Middle Zone and by its reappearance at the base of Upper Zone (Figure 3). Olivine is present in Middle Zone only as rare grains and as thin reaction products between pyroxene and Fe-Ti oxides. Lower Zone is further divided into three subzones (LZa, LZb,

and LZc) by the distinctive poikilitic texture of pyroxene in LZa and by the appearance of abundant ilmenite and magnetite at the base of LZc. In a similar way, Upper Zone is divided into three subzones (UZa, UZb, and UZc) by appearance of abundant, coarse apatite at the base of UZb and by the mosaic form of inverted ferrobustamite in UZc. Inverted pigeonite is found in all rocks up to the middle of UZa but is rarely abundant. Interstitial granophyre is common in the upper part of Upper Zone. These relations dictate the main minerals (olivine, plagioclase, augite, pigeonite, ilmenite, magnetite, and apatite) that should be included in any phase equilibria model used for the LS rocks. At present, the thermodynamic basis of COMAGMAT involves mineral-melt equilibria equations well calibrated only for the first six minerals [Ariskin and Barmina, 1999, 2000] and limits its application to LZa through UZa where apatite was not a primary phase.


ariskin: the compositional evolution of differentiated liquids 2.1. Intensive Parameters of Crystallization Other limitations are imposed by the probable range of intensive parameters of the crystallization. The fortuitous relations of the polymorphs of silica and Fe-rich pyroxenes in the Sandwich Horizon rocks allowed Lindsley et al. [1969] to show that that level crystallized at a pressure of 600±100 bars and a liquidus temperature of 970±20 C. Judging from the densities of the overlying gabbros and basalts, this corresponds to a depth of about 2 km for the Sandwich Horizon and places LZa 4.5 km and UZa 3 km below the original surface [McBirney, 1996]. It indicates an initial pressure of 1.3 kbar for the LZa and 0.9 kbar at the boundary between MZ to UZ. For such low pressures COMAGMAT has been shown to produce accurate phase equilibria calculations [Ariskin, 1999; Yang et al., 1996]. Estimates of redox conditions during crystallization are based on mineral equilibria calculations, electrochemical measurements, and results of melting experiments. Thermodynamic calculations for quartz-fayalite-magnetite and magnetite-ilmenite equilibria indicate that the oxygen fugacity at the base of LZa was slightly above QF M [Frost and Lindsley, 1992; Morse et al., 1980; Wil liams, 1971] and declined towards MZ, followed by a sharp decrease at UZ which results in as much as QF M -2 at the Sandwich Horizon [Frost et al., 1988]. Although Fe-Ti oxides were shown not to have preserved their original magmatic compositions, these estimates are in fairly good agreement with indirect evidence from experimental studies of phase relations of the natural rocks [McBirney and Naslund, 1990]. Measurements of intrinsic f O2 indicated values ranging from I W -W M [Sato and Valenza, 1980] to 0.5­1.5 units above the QF M buffer [Kersting et al., 1989]. Using the compositions of Ol-P ig -M t-I lm assemblages from the LZc, MZ, and UZa rocks, Wil liams [1971] calculated "frozen" equilibria temperatures of 1150­1050 C, with the highest values representing Lower Zone. This is in general agreement with both previous estimates for the SH (970 C) and those of melting experiments indicating temperatures of 1150­1002 C for partial melts that seem to represent differentiated (residual) liquids from the UZa to UZc [McBirney and Naslund, 1990]. Melting of the chilled gabbro K T -39 [Hoover, 1989] and results of geochemical thermometry of contact rocks [Ariskin, 1999] produced closed temperatures for the initial Ol-P l cotectics of about 1165 C. Attempts to use temperature-sensitive oxide and two-pyroxene equilibria has resulted in under-estimations of the liquidus temperatures due to significant re-equilibration during subsolidus cooling [Jang and Naslund, 2001].

7

2.2. Differences Between Op en and Closed to Oxygen Calculations Calculations for the average Layered and Marginal Border Series rocks were carried out at total pressure P = 1 kbar, 0.1 wt.% H2 O in the melt, and oxygen buffered conditions (QF M +1 to QF M -2), using a revised COMAGMAT model (version 3.65) which allowed us to calculate low-temperature

relations near 1100 C more accurately than could be done with COMAGMAT-3.5 (Appendix 2). This is critically important because only the lowermost rocks of the LZa an LZb zones fall within the range of experimental glasses used in the calibrations of COMAGMAT, whereas the LZc, MZ, and UZa rocks are too depleted in SiO2 and enriched with FeO because of their abundant Fe-Ti oxides (Figure 4d). I draw attention to this fact because of the validity of genetic conclusions based on the COMAGMAT-3.65 calculations depends strongly on the accuracy of phase equilibria extrapolations to the low-silica field, including estimates of the arrival of I lm and M t [Ariskin and Barmina, 1999]. A general impression of the calculated tra jectories can be obtained from Figure 5, where results of the phase equilibria modelling at QF M are shown. Note that all the calculations were conducted up to 80% crystallized (20% of residual liquid), as marked in the figure by thin solid (LS) and dashed (MBS) lines at the lowest temperatures for each computation. These plots demonstrate both common features and principal differences between the temperature sequences calculated for the main Skaergaard units. There is a general correspondence between the LS rocks and their MBS equivalents in the order of minerals crystallized at each horizon, which is in good agreement with the assumption that two series crystallized at the same time from the same magma (Figure 5). The composition of LZa indicate subcotectic crystallization of Ol + P l, followed by Aug and, at a late stage, little if any I lm. The LZb sequences indicate a large expansion of the field of high-Ca pyroxene, as is consistent with the appearance of abundant granular Aug crystals that characterize that zone, whereas the LZc tra jectories are marked by the early appearance of both I lm and M t. Similar phase relations were obtained for the MZ and UZa rocks (including their MBS equivalents), with the exception that Ol was found to be dissolved completely at a late stage (see below). It is interesting that the P ig field is noticeably wider in the marginal rocks, with the result that Ol disappears earlier owing to its reaction with residual liquids. It is certainly correlated with higher contents of SiO2 in the Marginal Border Series equivalents (Table 2) and may be indicative of the principal compositional differences between the liquids "trapped" in the MBS and the Layered Series rocks [Ariskin, 1999]. The oxide-rich gabbros of the Layered (LZcUZa) and the Marginal Border Series (LZc UZa ) demonstrate almost simultaneous precipitation of ilmenite and magnetite, with I lm crystallized slightly earlier in the LS rocks as 33 compared to MBS. Note, also, that the MBS oxide-rich types seem to have lower-temperature assemblages than those of the Layered Series at the same crystallinity. This latter observation seems to support ideas that the LS has been formed from a liquid that had already crystallized on the wall. The modelled disappearance of Ol at 1100 C in the MZ and MZ equivalents corresponds well to the absence of olivine within these units (Figure 3), but the relatively early disappearance of the mineral obtained for the UZa and UZa average compositions (Figure 5) is inconsistent with the petrographic character of these Ol-bearing units. This must be considered the most significant disagreement of the calculations with the natural observations. In an attempt to resolve the problem we performed an additional set of 20


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ariskin: the compositional evolution of differentiated liquids

Figure 4. Experimental conditions and compositional characteristics of glasses used in the recalibration of the COMAGMAT model (see Appendix 2). The range of temperatures and FeO contents in experimental glasses: A, C as a function of oxygen fugacity; B, D in terms of SiO2 contents. Run duration 48 hours. Extracted from the INFOREX experimental database [Ariskin et al., 1996, 1997]. In the FeO-SiO2 diagram the average bulk-rock compositions of the Layered Series units are shown for comparisons, see Table 2. calculations simulating the course of equilibrium crystallization at closed-to-oxygen conditions, using the average LZa, LZb, LZc, MZ, and UZa bulk-rock compositions. These trajectories were calculated at P = 1 kbar and 0.1 wt.% H2 O, with the initial Fe3+ /(Fe3+ +Fe2+ ) ratios in the melt ranging from 0.10 to 0.25. Some of the results are shown in Figure 6. They include comparisons of the "closed" modal proportions with those obtained in the "open" (QF M ) systems. The basic relations are presented here as a function of the total crystallinity with the amount of the residual liquid given as the ordinate. Such a presentation allows one to visualize the changes in phase proportions between the minerals and melt, as equilibrium crystallization proceeded. 2.3. Changes in Phase Prop ortions of Silicate Minerals and Fe-Ti Oxides The closed stable phase, LZc and UZa crease in the system calculations indicate olivine to be a at least up to a crystallinity of 80%, with the equivalents characterized by a significant inmodal Ol proportion at late stages. This is certainly correlated with the decline of P ig stability relative to those observed in the QF M buffered systems (Figures 5, 6). In fact, the dissolution of Ol is typical of both "open" and "closed" calculations that show an inflection in olivine proportions just after the appearance of magnetite (the later appearance of P ig reinforces the dissolution). It may indicate a peritectic reaction in the iron-enriched basaltic systems, such as: Ol + l Aug (±P ig ) + M t(±I lm) ± P l. The effects of this reaction on phase and chemical relations have commonly been observed in phase equilibria calculations since the end of 1980's [Ariskin and Barmina, 2000; Ariskin et al., 1988]. Petrological signatures of similar peritectic relations are documented in the Skaergaard rocks as thin rims of olivine produced by reaction between pyroxene and Fe-Ti oxides [McBirney, 1996]. As shown in Figure 6, the P ig field is wider in UZa than it is in UZc. It is correlated with higher average contents of SiO2 (43.1 against 41.1 wt.%), even though the Upper Series rocks are richer in FeO (22.5 against 21.1 wt.%, Table 2). This may be explained by: (i) a higher activity of SiO2 in the residual liquid, or (ii) a greater amount of residual liquid in the UZa rocks. Results of geochemical thermometry given below sup-


ariskin: the compositional evolution of differentiated liquids

9

Figure 5. Mineral crystallization sequences modelled for the average bulk-rock compositions of the principal units of the Layered Series and equivalent rocks of the Marginal Border Series (Table 2). Calculations of equilibrium crystallization using the COMAGMAT-3.65 program at P =1 kbar, 0.1 wt.% H2 O in the starting melt, and redox conditions corresponding to the QF M buffer.


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ariskin: the compositional evolution of differentiated liquids

Figure 6. Evolution of phase proportions in the crystallized melts corresponding to the average bulkrock compositions of the LZc and UZa subzones. Calculations of equilibrium crystallization were carried out using COMAGMAT-3.65 at P =1 kbar and 0.1 wt.% H2 O: open systems correspond to the QF M buffer; closed systems are defined by Fe3+ /(Fe3+ +Fe2+ ) = 0.15 in the starting melt.

port the first explanation, but it should be noted that the observed reappearance of Ol in UZa was followed closely by the disappearance of Ca-poor pyroxene at about the same level as the appearance of greater amounts of interstitial granophyre [McBirney, 1996]. The results have an important bearing on the modal proportions of Fe-Ti oxides. The proportions of M t ca