A. A. Ariskin
Vernadsky Institute of Geochemistry and Analytical Chemistry, Russian Academy of Sciences
Hunter and Sparks  used a modified version of Wager's calculations. They emphasized that at 48.1 wt.% SiO2 in the sample EG 4507 ( initial magma) the monotonous silica depletion trend can not be consistent with the low SiO2 contents (44-46%) of the observed "gabbroic assemblage" Ol - Pl - Aug Mt. 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  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 KT -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 objected 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].
An indirect way to define compositional evolution of the Skaergaard magma is to calculate perfect fractionation trajectories 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 Mt and Ilm crystallization [Ariskin and Barmina, 1999] confirme the conclusion [Ariskin et al., 1988; Toplis and Carroll, 1996] that in systems open with respect to oxygen Mt 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 - KT -39 [Hoover, 1989] and TM [Ariskin, 1999]. The latter is an average composition of six modeled 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 sub-cotectic ( Ol+Pl ), with KT -39 slightly oversaturated with Ol , whereas TM is closer to the Ol+Pl+Aug saturation. Liquid lines of descent representing the fractionation at 1 atm and QFM 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 Mt - and Ilm -models. With these uncertainties [Ariskin, 1999] reached the following conclusions: (i) both parents exhibit only minor decrease in SiO 2 as compared to the proposed FeO-SiO 2 trend [Brooks and Nielsen, 1978, 1990; Wager and Brown, 1967]; (ii) all of the calculated trajectories indicate a well-defined inflection point, where the content of SiO 2 starts to increase due to the appearance of Mt; (iii) the maximum possible iron enrichment of residual melts is approximately 18 wt.% FeO for KT -39 and 20% for TM.
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 QFM and WM buffers and the temperatures lower 1150o 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).
Utilizing the COMAGMAT-3.5 crystallization model and a modeling principle known as "Geochemical thermometry" (Appendix A), 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" EG 4507 [Wager and Brown, 1967]. Calculations of equilibrium crystallization at 1 atm pressure, dry conditions and QFM 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 Pl for a plagioclase-rich rock. Two samples (including EG 4507) showed sub-cotectic ( Pl+Ol ) relations in the range 1230-1250o C. In accordance with the sequence of crystallization, the calculated lines of descent yielded an obvious intersection in the range of 1175-1155oC (Figure 2). The temperature of 1165 10oC was accepted as representing the initial temperature of the Skaergaard parental magma. The average liquid composition representing this cluster of six evolutionary lines at 1165oC is given in Table 1. It is close to liquid compositions that Hoover  obtained in his melting experiments with marginal rocks. In this paper, estimates of major-element geochemistry for "residual liquids" and primary mineral proportions for the LZ, MZ, and UZ rocks are first presented.
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.
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.  to show that that level crystallized at a pressure of 600 100 bars and a liquidus temperature of 970 20oC. 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 QFM [Frost and Lindsley, 1992; Morse et al., 1980; Williams, 1971] and declined towards MZ, followed by a sharp decrease at UZ which results in as much as QFM -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 aggreement 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 IW - WM [Sato and Valenza, 1980] to 0.5-1.5 units above the QFM buffer [Kersting et al., 1989].
Using the compositions of Ol - Pig - Mt - Ilm assemblages from the LZc, MZ, and UZa rocks, Williams  calculated "frozen" equilibria temperatures of 1150-1050oC, with the highest values representing Lower Zone. This is in general agreement with both previous estimates for the SH (970oC) and those of melting experiments indicating temperatures of 1150-1002oC for partial melts that seem to represent differentiated (residual) liquids from the UZa to UZc [McBirney and Naslund, 1990]. Melting of the chilled gabbro KT -39 [Hoover, 1989] and results of geochemical thermometry of contact rocks [Ariskin, 1999] produced closed temperatures for the initial Ol - Pl cotectics of about 1165oC. 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].
The closed system calculations indicate olivine to be a stable phase, at least up to a crystallinity of 80%, with the LZc and UZa equivalents characterized by a significant increase in the modal Ol proportion at late stages. This is certainly correlated with the decline of Pig stability relative to those observed in the QFM 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 Pig reinforces the dissolution). It may indicate a peritectic reaction in the iron-enriched basaltic systems, such as: Ol+l Aug ( Pig) + Mt ( Ilm) Pl . 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 Pig 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 support 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 Mt calculated by COMAGMAT-3.65 in the systems closed to oxygen are two-thirds of those calculated at open conditions (Figure 6). This is not surprising, because of the effect has been demonstrated experimentally [Osborn, 1959; Presnall, 1966] and confirmed by phase equilibria calculations [Ariskin et al., 1988; Ghiorso and Carmichael, 1985; Toplis and Carroll, 1996]. A more informative approach is to analyze changes in the modal proportions of ilmenite. At closed conditions an increase in the Ilm proportion is observed, so that the bulk Ilm/Mt ratio shifts in favor of ilmenite (Figure 6). This suggests that closed-system crystallization more realistically reproduces the observed relations, since ilmenite is more abundant than magnetite in the Skaergaard rocks [McBirney, 1989]. The simultaneous crystallization of both oxides from the LZc and UZa melts may be explained from experimental results of Snyder et al. . These authors demonstrated that for a closed iron and titanium enriched system, Ilm precipitation results in the increase of Fe3+/(Fe3+ +Fe2+ ) in the melt, driving the liquid into the Mt stability field. Crystallization of magnetite has the opposite effect depleting the residual melt in Fe2O3 and lowering the relative f O2 value, so that the liquid is quickly driven into the Ilm field. The COMAGMAT calculations are also consistent with this observation (Figure 6).
I also draw attention to the percent melt shown next to the calculated trends (Figure 7). It is clear that the main objective of interest is in the range of 20 to 60% of the remaining liquid, which corresponds to the expected original modal proportions of cumulus grains (primocrysts) of 80 to 40 wt.%. For the case of the initial composition of LZc, one can see small but real differences of the iron contents of the melt at silica contents of 46 to 50 wt.% SiO2: 19-20 wt.% FeO at closed conditions and 17 to 18 wt.% at QFM. The simulations of UZa at QFM have a much wider range of liquids evolving towards somewhat unusual ferroandesitic compositions containing as much as 13 to 14 wt.% FeO at 60 wt.% SiO 2. Nevetheless, the UZa trend for closed conditions falls in a more realistic range of 47.5 to 52.5% SiO2 with 18 to 20% FeO. These results may indicate that our estimates of the differentiated liquid compositions could be best for the LZa and LZb rocks (where primocrysts of oxides are absent). If so, they could provide a good approximation for LZc (and possibly MZ) but yield an approximate composition for the UZa rocks, for which accurate estimates of the initial f O2 values and bulk Fe3+/(Fe3+ +Fe2+ ) ratios are necessary.
Thus, the preliminary calculations indicate complex relations between modal proportions of silicate and oxide minerals, residual liquid compositions, and redox conditions. In general, at relatively oxidizing conditions ( > QFM ) one would expect early precipitation and larger proportions of Mt resulting in the pronounced increase in the silica with a complementary depletion in FeO in the melt. Sharply increasing SiO 2 causes Ol dissolution to occur after the appearance of Ca-rich pyroxene and magnetite, owing to the peritectic reaction discussed above. This process is enhanced by the expansion of the Pig stability field resulting from the enrichment with SiO2 and moderation of the Mg/(Mg+Fe) ratio in the melt [Grove and Baker, 1984], leading to the disappearance of olivine at late stages. Under more reduced conditions Ilm may crystallize earlier than Mt, so that the increase in SiO2 comes relatively late. In this case the field of Pig decreases, whereas the olivine stability field increases (as the result of uncompleted Ol dissolution at the same amounts of residual liquid).
Under closed conditions, the Ol - Pig relations depend strongly on iron oxidation in the initial melt. At higher Fe3+/(Fe3+ +Fe2+ ) values, the Pig field is slightly increased because of the enrichment with SiO2 caused by earlier precipitation of magnetite. The decline of ferric iron is equivalent to melt reduction accompanied by the decrease in the Mt crystallization temperature and an expansion of the Ilm stability field (Figure 6). The general effect of closed conditions is shown to decrease the total amount of Fe-Ti oxides crystallized, so that the enrichment with SiO2 and depletion of iron becomes less pronounced (Figure 7). From the results of phase equilibria calculations one can conclude that only perfectly closed crystallization produces the predominance of Ilm over Mt observed in the LZc rocks, as well as the preservation of a part of primary Ol crystals within the Upper Zone. Note, however, the latter inference is based on the results obtained for relatively oxidized systems with a bulk Fe3+/(Fe3+ +Fe2+ )=0.15 (compare Figures 6 and 7). Judging from the liquidus temperatures (Figure 5), it corresponds to an oxygen partial pressure slightly below NNO. Thus, one should not exclude the possibility of more reduced conditions at which the liquid was depleted in SiO 2 so that Pig was not crystallized at the late stage, whereas Ol continued to crystallize in the Upper Zone.
Geochemical thermometry is a computer-based technique designed to extract initial temperature, interstitial liquid composition, primocryst mineral compositions and modal proportions, as "recorded" in the whole-rock chemistry [Frenkel et al., 1988a, 1988b]. This approach is based on interpretations of the results of phase equilibria calculations for igneous rocks and their "parents" using the COMAGMAT magma crystallization model [Ariskin, 1999], see Appendix A. Several varieties of geochemical thermometry have been proposed with the purpose to evaluate the original temperature, crystallinity and liquid composition for intrusive magmas at an initial stage corresponding to their emplacement into a magma chamber [Ariskin and Barmina, 2000]. Here the technique is applied to the Layered Series cumulates from of the Skaergaard intrusion.
A dataset of 65 bulk-rock compositions from the LZa ( n = 8), LZb (16), LZc (11), MZ (20), UZa (10) units have been selected to use in geochemical thermometry. The structural location and chemical characteristics of the samples are listed in Table 1 of McBirney . The only change is in the sample LB-238 moved from LZb to LZa. In addition, five average bulk-rock compositions of the principal units of the Layered Series and five average compositions of equivalent rocks of the Marginal Border Series were used (Table 2).
As was shown in the previous paragraph, the calculations for a system closed to O 2 seem to be more suitable if one tries to reproduce both the naturally observed modal mineral proportions and corresponding differentiated melt compositions. To conduct the closed system calculations for the Skaergaard rocks, I decided to use the bulk-rock FeO and Fe2O3 contents as a starting point for each of the above selected samples. This decision is based on the premise that most of the Skaergaard rocks are remarkably fresh and show little if any secondary oxidation. It allows one to assume that the phase equilibria simulations with the initial FeO and Fe2O3 contents may result in the intrinsic magmatic Fe3+/(Fe3+ +Fe2+ ) ratios in the residual liquids, providing a means of estimating redox conditions during differentiation.
To make sure that such calculations realistically reproduce both the original melt-mineral proportions and the residual liquid compositions we first calculated phase relations for 11 of the LZc rocks for which modal mineral proportions were available, see Table 3 of McBirney . Part of these data including those for Ilm and Mt is listed in Table 4. This tabulation also includes initial TiO2, FeO and Fe2O3 contents in the LZc rocks with the normative proportions of the oxides calculated on the assumption that Ilm and Mt are pure end-members (FeTiO3 and Fe3O4 ).
Results of phase equilibria calculations for the same LZc rocks are given in Table 4 in the form of modeled modal proportions at given amounts of crystallinity and chemical compositions (TiO2, FeO, Fe2O3 ) of the complementary residual liquids for each sample. The calculations for conditions of closed O2 were carried out at total pressure P=1 kbar and an assumed H2O content of 0.1 wt.% in all rocks. The results permit one to calculate total proportions of Fe-Ti oxides in the "computer simulated rocks" by summation of the liquidus mineral proportions at a given system crystallinity with the normative oxide proportion in the residual liquid weighted in accordance with the modal amount of this melt. These estimates are given in two lower strings of Table 4. Comparisons with the observed modal proportions indicate that COMAGMAT-3.65 closely reproduces the final amounts of ilmenite and Ti-magnetite for all but two samples, LC-417 and LC-465.
Using compositions of eight LZa rocks to estimate starting liquids, we have calculated the course of equilibrium crystallization up to a total crystallinity of 75 to 85%. These calculations were conducted at the same conditions as for the LZc rocks, with crystal increments Ds = 1 mole %. For 5 of the initial compositions with MgO > 9.5 wt.% (LA-36, -95, -237, -401, -770) a wider temperature range was obtained for Ol, and Pl was the second mineral to crystallize, followed by Aug and in some cases Mt, and finally by Pig: Ol (1322 to 1258oC) Pl (1237 to 1193oC) Aug Mt (1140 to 1110oC) Pig (1125 to 1110oC). Relatively early crystallization of plagioclase and pyroxene was observed for the most calcium-rich sample, LA-238: Pl (1230oC) Ol (1196oC) Aug (1181oC) Mt+Pig (1120-1110oC), whereas the most silica-rich sample, LA-753 (51 wt.% SiO 2 ) indicated sub-cotectic relations: Ol+Pl (1210 to 1200oC) Aug (1153oC) Mt+Pig (1125 to 1115o C). Ilm was found in minor amounts in three cases at the temperatures < 1110oC. Only in the most titanium- and iron-rich sample, LA-347, (5.8% TiO2 and 18.3% FeO ) was Mt the first mineral to crystallize at 1190oC followed with Pl (1140oC), pyroxenes and ilmenite at 1120-1110o C.
Initial mineral compositions. Calculated average compositions of Ol and Pl for the six samples at 1145oC are Fo74.91.5 and An66.31.9. The estimated An content in plagioclase is in excellent agreement with observed compositions (Figure 3), whereas the calculated Ol composition is 7 mole % more magnesian. This difference cannot be attributed to uncertainties of the COMAGMAT model, which generally has an accuracy of about 1-2% Fo [Ariskin and Barmina, 2000]. It is more likely to indicate re-equilibration of the initial olivine crystals with an iron-enriched residual liquid constituting on average up to 50 wt.% of the mineral-melt assemblages (see below).
Oxygen fugacity. Also shown in Figure 8 are the T - log f O2 dependencies calculated for each liquid line of descent as a function of the evolving Fe3+/Fe2+ ratio in the melt. These calculations use the equation by Sack et al.  rearranged to compute log f O2 for a given temperature and melt composition. As was shown by Nikolaev et al. , such calculations have an accuracy of 1 log unit. In ideally closed systems, the modeled T - log f O2 trajectories should intersect at the same temperature as the liquid compositions. In such a case, an average f O2 value calculated for different evolutionary lines could be considered as an estimate of the oxygen fugacity intrinsic to the initial mineral-melt equilibrium assemblage. For the LZa molten rocks these lines form a wide cluster of log f O2 values (Figure 8) with an average of -7.50.6 at 1145oC. This is about 1.5 log units above the QFM buffer and is consistent with the measurements of Kersting et al. . Note that the standard deviations of 0.6 log units can be attributed to both the accuracy of Sack's equation and to the analytical uncertainties in the bulk-rock FeO and Fe2O3 contents. There is no reason to suspect that the COMAGMAT mass-balance calculations for the Ol+Pl field are incorrect [Yang et al., 1996].
Three main crystallization sequences were found to be peculiar to the LZb rocks. Four trajectories (LB-104, 234, 600, and 724) indicate the order: Ol (1370 to 1240oC) Pl (1220 to 1170oC) Aug Mt (1160 to 1110oC) Pig (below 1130oC); in five cases (LB-68, 235, 269, 476, 655) augite appeared on the liquidus at temperatures of 1200 to 1170oC simultaneously with or just after olivine, and in all seven samples plagioclase was the first crystallizing mineral (1250 to 1190oC) followed by Ol and Aug. These phase relations indicate that Ol and Pl continued to be the main cotectic phases with Ca-rich pyroxene crystallizing as the third mineral. Ti-magnetite usually is the fourth, whereas Pig and Ilm were found only at late stages at temperatures below ~1120oC and 1100oC, respectively. These calculations are in general agreement with petrological data indicating the change in the structural position and abundance of augite within the LZb horizon and do not contradict with a standard interpretation that Aug crystallized on a cotectic line with Ol and Pl [McBirney, 1989; Wager and Brown, 1967].
Compared to LZa and LZb, simulations for eleven LZc compositions show much more diversity in the order of appearance of minerals on the liquidus. This is due to the abundant crystallization of oxides, so that for 6 samples (LC-306, 342, 465, 727, 746, and 928) Ilm and Mt are the first minerals to crystallize in the range of 1270 to 1170o C. The order of crystallization of silicate minerals for these samples is variable ( Ol Pl Aug + Pig or Ol Aug + Pig Pl ), however pyroxenes never crystallized at temperatures above 1110oC. This can be taken as the upper limit for the temperature controlling original mineral-melt equilibria. In three cases Ol crystallized first (LC-67, 534, and 605) followed by Pl, pyroxenes and oxides. Sample LC-406 indicated over-saturation with plagioclase (1203oC), whereas LC-417 started with the cotectic assemblage Ol+Pl (1157oC). These results support the notion that Ilm and Mt should be considered among the major primary phases of LZc [Wager and Brown, 1967].
Mineral compositions of Ol, Pl, and pyroxenes equilibrated with the LZc residual liquid have the same tendency to have low amounts of the refractory components compared to the underlying mineral-melt assemblages (Table 3). Appearance of Pig crystallizing near 1100oC was observed only five times of the 11 samples, so that the question of its original primary nature in the LZc rocks is still open to question. The initial Ilm and Mt compositions are given in Table 3 in the form of the FeTiO3 and Fe2 TiO4 activities calculated using the equations of Stormer . These data indicate almost stoichiometric ilmenite and moderately titaniferous magnetite. This is consistent with a shift of the calculated f O2 values towards more reduced conditions close to the QFM buffer (Figure 11).
Twenty of the MZ samples analysed by McBirney  were used in the phase equilibria calculations. Differences in the starting compositions and the extent of iron oxidation result in various crystallization sequences in which six minerals ( Ol, Pl, Aug, Pig, Ilm, Mt ) change or replace each other over small temperature intervals, often within the accuracy of the COMAGMAT model (10-15oC). Only their general features are described here. The most surprising result is that all Ol -free samples produce some amount of olivine at early and middle stages of crystallization. On average, the amount of olivine was 3 to 5 wt.% at 30% total crystallinity, but in some cases it was as much as 8 to 10 wt.% Ol. In eight samples Ol was the first mineral to crystallize at temperatures of 1190 to 1160oC (MZ-53, 55, 58, 61, 747, 824, 826, 831). However, at lower temperatures corresponding 50 to 70% crystals this olivine was completely dissolved by reaction with pyroxenes, oxides and melt. For the eight samples that contain the most TiO2 and iron (MZ-28, 172, 438, 603, 613, 729, 830, 901), Fe-Ti oxides crystallized first, followed by Ol, Pl, and pyroxenes. In four cases, the starting melts were slightly oversaturated with plagioclase at 1210 to 1160oC (MZ-99, 343, 530, 864). Note that in only two cases was Aug found at temperatures above 1105oC, whereas Pig always crystallized below 1100oC. It is important to note that in half of the samples pigeonite appeared at a total crystallinity of 20 to 40%. This may indicate that the mineral may have been among the primary crystallizing phases. In any case, a substantial amount of Pig was certainly formed at later stages by the peritectic replacement of olivine.
Mineral compositions in Table 3 are weighted average values for Ol, Pl, pyroxenes and oxides at 1090oC. The modelled Pl composition (An54.22.9 ) is close to the observed plagioclase composition of An51, while calculated pyroxenes and olivine relicts are again more magnesian. Ilmenite became slightly richer in FeTiO3, whereas the activity of the ulvospinel component in Ti-magnetite increased considerably ( Ulv72 in MZ versus Ulv54 in LZc). At first view this seems strange, because the calculated TiO2 in the melt was found to decrease somewhat. However, if one considers the T - log f O2 diagram (Figure 12), it is clear that within the MZ zone the system was evolving to more reduced conditions, slightly below QFM (Table 3). This conclusion is in excellent agreement with the previous thermodynamic and experimental estimates indicating a smooth f O2 decline towards MZ, where redox conditions corresponded to about 0.5 log units below QFM [Frost and Lindsley, 1992; McBirney, 1996; Morse et al., 1980; Williams, 1971]. On the other hand, this somewhat surprising behavior may indicate that the initial Fe2O3 and FeO contents used in the closed system calculations really reflect both the total amount of oxides and Fe3+/Fe2+ ratios intrinsic to the MZ residual liquids. It supports the premise that most of the Skaergaard rocks have undergone insignificant secondary oxidation.
Phase equilibria calculations for ten LZa rocks produced the lowest temperatures for those parts of the Layered Series studied thus far. None of the assemblages had a silicate liquidus temperature above 1180oC. On average, the modeled sequences are similar to those of MZ (Figure 5), starting from sub-cotectic ( Ol + Pl oxide) or compositions slightly oversaturated with olivine (UA-102, 136, 199, 200, 300, 315, 614, 786). The two samples with the highest FeO contents (UA-131 with 35 wt.% and UA-425 with 46 wt.%) have a large magnetite field and were excluded from further considerations because of strong deviations of their trends from those calculated for other samples. Both pyroxenes crystallized almost simultaneously at temperatures below 1100oC, with the Pig crystallization interval slightly decreased with respect to the MZ simulations. This resulted in an increase in the modal proportion of Ol (Figure 10), but Pig did not disappear completely (as can be seen in Figure 3) but continued to be a primary phase. This corresponds observations of nature that Pig continued continued to present in many rocks in the lower half of UZa [McBirney, pers. comm.].
Mineral compositions of the six-phase liquidus assemblage calculated at 1085oC follow the same trend of the depletion in the refractory components as in lower horizons (Table 3). The modeled plagioclase (An 48.03.1 ) is similar to the composition of An 46 observed at the base of UZa (Figure 3), whereas the calculated Ol (Fo52.52.5 ) is much more magnesian than the observed composition (Fo40 ). Both pyroxenes have higher Fs/En ratios, although the modeled augite is of 4 to 5 mole % richer in Wo than the natural one. Although the ilmenite composition remains unchanged, magnetite demonstrates a further increase in Fe2TiO4 (81.4 6.7 mole % Ulv ).
Geochemical thermometry for the Layered Series rocks produced a probable range of the "formation temperatures" that represent original intersitial material (Table 3). It spans a range of only 60oC between 1145 10oC at LZa to 1085 10oC at UZa. The observed values correspond to other estimates based on other approaches [Hoover, 1989; Jang and Naslund, 2001; McBirney and Naslund, 1990; Williams, 1971]. The phase equilibria simulations also define the range of oxygen fugacities intrinsic to the initial liquids. Our results indicate a sharp decline in log f O2 values from 1 to 1.5 units above QFM in LZa and LZb to QFM and lower oxygen fugacities starting from LZc. This is also in fairly good agreement with previous estimates [Frost and Lindsley, 1992; Frost et al., 1988; Kersting et al., 1989; Lindsley et al., 1969; Morse et al., 1980; Williams, 1971].
The main result of the geochemical thermometry is an estimate of the initial compositions of differentiated liquids (Table 3). The calculated liquids demonstrate strong enrichment in FeO (up to ~18 wt.%) and TiO2 (up to ~5.5 wt.%), consistent with the notion that the Skaergaard magma evolved in a manner typical of iron- and titanium-rich tholeiitic magmas [Brooks et al., 1991]. As Wager and Brown  inferred from simple mass-balance considerations, there was essentially no silica depletion in the liquids from LZa through UZa. Phase equilibria calculations do, however, indicate a small increase in SiO 2 during formation of LZa. This agrees with the results of melting experiments for Ol - Pl assemblages in low-pressure tholeiitic systems [Grove and Baker, 1984]. The transition from the LZb ( Ol+Pl+Aug ) to LZc ( Ol+Pl+Aug + oxides) rocks is accompanied by a decrease in SiO2 as a result of continued crystallization of Ca-rich pyroxene, whereas the appearance of Mt and Ilm brings renewed silica enrichment.
It is interesting that this increase in SiO 2 is not as significant as it commonly is during crystallization of Fe-Ti oxides from tholeiitic magmas buffered for f O2 [Ariskin and Barmina, 1999, 2000]. It is certainly consistent with the absence or very small decrease in FeO (Table 3) that could be interpreted as evidence that the Skaergaard magma evolved as a system closed to oxygen [Osborn, 1959; Presnall, 1966; Toplis and Carroll, 1996]. Nevertheless, we draw attention to the fact that the decrease of f O2 within this temperature range is too sharp to be generated during simple closed-system crystallization. It moves us to conclude that the Skaergaard magma chamber has not been absolutely closed and could have been open to a reducing agent that was responsible for the change of oxygen fugacity during a middle-to-late stage of differentiation.
Another important factor bearing on the trend of differentiation is the modal proportions of primary mineral assemblages and residual (interstitial) liquids. Figure 10 summarizes the results of the geochemical thermometry calculations related to the original characteristics. The diagram to the right of the figure was constructed from average modal proportions calculated for each unit at the corresponding temperature (Table 3). The observed modal assemblages closely resemble the major calculated proportions of minerals for the main Skaergaard units. The only exception is probably over-estimated amounts of pigeonite in the UZa zone, although the mineral is known to present in small amounts there [McBirney, 1989; 1996]. Similarly, the modeled proportions of the evolving residual liquid through the LS section considered here was found to be about 50 wt.% (Figure 10). This may indicate that the compaction process was not efficient, at least at the early and middle stages of the Skaergaard magma differentiation. This conclusion is consistent with petrographic observations that adcumulus textures (expected to indicate a higher crystallinity) were widely observed at higher levels corresponding to UZb and UZc horizons [Wager and Brown, 1967]. Note, however, that the modal-proportion diagram in Figure 10 represents average values for the main units: in fact, the calculated amounts of intercumulus liquid differ for each particular rock, ranging from about 30% to 60% (as a residual porosity - Morse ).
The most conspicuous contradiction of the results of geochemical thermometry is the obvious gap between LZa and LZb liquids and those of LZc through UZa (Table 3). It is seen in a small (~1 wt.%) return of CaO contents to higher and SiO 2 contents to lower values starting from the LZc liquid (Figures 11-13). The simplest explanation would be that the differences are due to uncertanties of the COMAGMAT program, because the gap is within the accuracy of the technique (compare the 1s deviations in Table 3). One could say that the COMAGMAT-3.65 model slightly under-estimates the crystallization temperature for Aug (Figure 5), so that later precipitation of the high-Ca pyroxene was modeled from more calcium rich liquids. Nevertheless, a nearly two-fold increase in TiO2 from LZb to LZc must also be explained because it would require at least 40 to 50% crystallization of the LZb melt. Thus, the observed gap may be an important genetic clue that should be examined more closely.
First of all, note that Wager's classical trend is totally inconsistent with the standard low-pressure evolution of crystallizing tholeiitic magmas. With one exception, these points lie in the Ol - Pl field far from the experimentally defined Ol - Aug - Pl boundary, even though most of points should correspond to co-saturation with olivine and Ca-rich pyroxene. Only the last point is found to fall in the Ol - Pl - Aug - Pig field, but this is not surprising because the composition is similar to that of the Sandwich Horizon which is thought to represent an almost pure liquid that was the final product of differentiation [Wager and Brown, 1967]. It allows us to conclude that the trend of FeO versus SiO 2 calculated using inferred proportions of the main differentiated units of the Skaergaard intrusion (Figure 1) is actually artificial and does not provide a valid interpretation of all the major components of tholeiitic magmas. The only way it can be made to correspond to experimental data is to assume crystallization at elevated pressures where the CPX field is increased. In such a case, one should re-examine the previous pressure estimates of Lindsley et al.  and assume that Aug crystallized as a primary mineral beginning from the LZa zone.
The experimentally determined compositions of McBirney and Naslund  and the calculated compositions of residual liquids of Ariskin ; this study] have a more realistic distribution in Figure 14. Most of the points are in the vicinity of the Ol - Aug boundary which spans a range of liquids in which the upper margin is shifted slightly towards CPX. This is consistent with the presence of augite as a primary liquidus mineral dominating the course of crystallization for the most of the Layered Series rocks. Note that the calculated TM, LZa, and LZb liquids are located very close to the curve with a trend directed straight towards the Pig field, whereas the LZc to UZa melts with more normative diopside are located above the Ol - Aug - Pig peritectic point. It is difficult to find a reasonable petrologic interpretation that can link the two slightly different series of residual liquids by simple crystal fractionation. The first one is most likely to reflect the early stages of differentiation along the normal trend of tholeiitic magmas as seen in volcanic suites [Ariskin et al., 1988; Hunter and Sparks, 1987; Toplis and Carroll, 1996]. The second trend may represent another type of chemical evolution that has only recently been suspected. Possible explanations include the possibility that the diopside normative and Ti-rich component was (i) a late stage residual liquid percolating upward through compacting crystals, (ii) an influx of the titanium and chlorine rich fluid at a middle-to-late stage [McBirney, in press], or (iii) some sort of compositional stratification in the main mass of magma. I think that any of the proposed mechanisms is linked with a more general Skaergaard problem concerning the absence of the mass-balance between probable parental magmas and the average Skaergaard intrusion composition.
If we assume the Skaergaard Layered Series was formed during a closed system process of crystal fractionation from a parental magma close in composition to the marginal gabbro or the calculated initial "trapped" melt (Table 1), one would expect the weighted average composition of the Layered Series and two Border Series to resemble that of the proposed parental magma. Because the original shape of the intrusion is unknown, it is very difficult to estimate reliable proportions between the LS, UBS, and MBS rocks. Nevertheless, there are reasons to consider the LS rock as constituting at least 40-50% of the intrusion rocks providing the main contribution to the whole-body composition [McBirney, pers. comm.].
The main feature of the Layered Series is the predominance of magnetite- and ilmenite-rich rocks that compose of about 70% of the observed section (Figure 10). These rocks contain of 43 to 44 wt.% SiO 2, 5 to 6 wt.% TiO 2, and 22 to 23 wt.% FeO , so that the average Layered Series composition must also be depleted in silica. The LS characteristics used to calculate a range of possible Skaergaard intrusion average compositions ( SIAC ), using varied proportions of the UBS and MBS rocks. Some of the results for three different proportions of the LS, UBS, and MBS are listed in Table 5. For comparisons, two compositions representing the chilled gabbro KT -39 and our calculated initial "trapped" liquid (Table 1) are also included into the table. Even a superficial view of the compositions reveals a strong misbalance between the parental compositions and that of the differentiated body. It is most apparent for TiO2 and P2O5 which are almost two fold higher in SIAC as compared to the proposed parents. Moreover, the intrusion composition has of 2-4 wt.% less SiO 2 and much more iron. The observed discrepancy may be explained by the effects of closed- or open-system crystallization. The "closed" interpretation includes two possibilities, one in which the bulk composition of the Layered Series has more SiO2 and less TiO2, FeO, and phosphorus than the exposed section. This would imply a "Hidden Layered Series" with relatively silica-rich Pl - Opx rocks at depth [Wager and Brown, 1967]. The second possible explanation postulates substantial amounts of felsic differentiates almost completely removed by erosion [Hunter and Sparks, 1987].
Results of our calculations indicate that the compositional difference between liquids "trapped" in marginal rocks ( TM ) and that of LZa is small (Tables 1 and 4, Figure 14). They are consistent with the observation that the composition of Pl at the base of the exposed Layered Series is similar to that in the uppermost part of the Upper Border Series (Figure 3). Thus, there are no reasons to assume a large volume of a "Hidden Layered Series", assuming, of course, that the initial Skaergaard magma was free of suspended crystals. If one suggests the Skaergaard magma came to the chamber with an amount of crystalline material ( Ol+Pl ) equilibrated with a liquid close to the TM composition [Ariskin, 1999], the existence of complementary troctolitic rocks is inevitable, despite of some geophysical data arguing against the huge Hidden Zone [Blank and Gettings, 1973; McBirney, 1975]. Even if some amount of hidden troctolitic material exists, however, it is unlikely that crystallization in a closed system could produce large volumes of rocks rich in Fe-Ti oxides without more felsic differentiates (see Figure 7 and Figures 11-13).
In recent years more attention has been given to open-system crystal fractionation in the Skaergaard magma chamber with a later influx of a fluid or melt that was very rich in iron and titanium. There is some evidence of this in the geological relations, trace-element distribution, and petrographic features [McBirney, 1995]. Thus, this alternative interpretation may be also attractive in the light of the calculated shift of the LZc to UZa residual liquids plotted in the OLIV - CPX - QTZ diagram (Figure 14). Note, however, that a postulated influx of iron-rich material is difficult to reconcile with the vast amounts of the fluid needed to alter the primary igneous rocks metasomatically, or with the recent geochemical evidences for differentiation under closed conditions [Jang and Naslund, 2001].
A specially calibrated version of the COMAGMAT program allowed the author to define the range of initial temperatures (1145 to 1085oC) and oxygen fugacities (1-1.5 log units above QFM to slightly below QFM ) intrinsic to the original crystal mush from which the rocks from LZa to UZa have been crystallized. In parallel, average major-element compositions of residual (interstitial) liquid were calculated for the same units (Table 3). The result demonstrates a trend of continual enrichment of FeO (up to ~18 wt.%) and TiO 2 (up to ~5.5 wt.%) with only minor variations in the SiO 2 contents (48 to 50 wt.%).
Projection of the compositions onto the low-pressure OLIV - CPX - QTZ diagram (Figure 14) supports several lines of petrologic evidence that most of the Layered Series crystallized on an olivine-clinopyroxene cotectic (along with plagioclase and oxides). Wager's compositions, however, lie too far from the Ol - Aug boundary to represent a realistic approximation of the low-pressure Skaergaard magma evolution. The simple mass-balance approach that Wager  proposed is not consistent with phase equilibria constraints.
Systematic differences observed 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 phase equilibria used in the model. If one examines the T - X diagrams in Figures 11-13, one is easily convinced that the liquid lines of descent at temperatures below 1100o C may be changed dramatically. For example, CaO contents in the calculated liquids may differ of 2 to 2.5 wt.% in a temperature range of 10oC. Nevertheless, we cannot exclude the possibility that the compositional differences discussed above could be the geochemical effects of a late-stage process involving migration and re-equilibration of interstitial liquids.
A remarkable feature of geochemical thermometry is its ability to estimate original modal proportions of the primary liquidus minerals and residual liquid. The results summarized in Figure 10 are in fairly good agreement with the observed mineralogical assemblages that define the main units of the Layered Series. Estimated amounts of interstitial melt range around 50 wt.%. I believe that estimate is correct for the rocks of LZa and LZb, but it may be underestimated by 10 to 20 wt.% for gabbros with abundant magnetite and ilmenite owing to a slightly over-estimated (up to 10oC) initial rock formation temperatures.
Finally, our interpretations of the results of phase equilibria modeling depend upon the accuracy of the COMAGMAT calculations. The problems we have outlined indicate that even careful calibrations giving a precision of 10oC may be insufficient to give unambiguous conclusions, especially when dealing with multi-phase assemblages including 5-6 minerals plus melt. This does not mean we are not confident of the utility of magma crystallization models. Despite the difficulties we have oulined, systematic phase equilibria modeling of intrusive rocks can be a controlling factor allowing petrologists to surmount the inadequacies of other techniques, such as mass-balance calculations. In the context of modeling the Skaergaard rocks, the most serious need are experimental data on the melting relations of magnetite and ilmenite in assemblages spanning the unexplored range of Fe-Ti rich and silica depleted compositions (Figure 4d). These data could be used in the future for more accurate calibrations of COMAGMAT and other magma crystallization models.
This method is based on the premise that any crystal-liquid mush that produces a completely crystalline rock should pass through a stage when relative motions of solids and interstitial liquid have ceased, so that the system may be considered as closed. In this case, it is theoretically feasible to calculate all thermodynamic parameters for the initial system, using phase equilibria models designed for the calculation of equilibrium melting/crystallization relations in igneous rocks. Like the other methods, this approach entails certain assumptions as to the mechanisms of differentiation of intrusive rocks.
If crystals grew to sufficiently large sizes before settling to the floor, then the interstitial liquid would have approximately the same major- and trace-element concentrations as the liquid constituting the main part of convecting magma in which they were accumulated. This is also true if the crystals nucleated and grew in situ, with the residual liquid being exchanged by chemical diffusion with the main magma body. In both cases the average phase composition for the rock at any particular horizon may be given by the expression:
where Fjcr and Flin are initial modal proportions of a particular mineral j (1 j M ) and interstitial liquid ( l ). The bulk chemical composition of such a crystal mush is expressed as:
where Cil(0) is the interstitial liquid composition, whereas Cij(0) should be attributed to average weighted contents of element i in mineral j accounting for a range of original compositions for each particular mineral.
The physical history of the formation of a rock from the crystal mush may be consistent with "closed" and "open" system evolution. The first situation means that cooling of the heterogeneous layer results in simple chilling (contact rocks), or interstitial crystallization with new minerals precipitated from the liquid matrix (orthocumulates). It is also possible that primary crystals undergo some amount of additional growth at the expense of the interstitial melt, followed by crystallization of the residual matrix. In all cases closed systems conserve the whole composition given by Equation (2), even though the compositions of individual primocrysts and interstitial material may change. The primary "trapped melt" temperature may be called by the formation temperature of the magmatic rock in the geochemical sense that the bulk rock composition did not change during further cooling. If so, Fjcr and Flin should be treated as initial or original modal proportions, with Cij(0) representing original mineral compositions [Ariskin and Barmina, 2000; Barmina et al., 1989a].
In "open" systems the original modal proportions may change by later growth or compaction of the crystals (increasing Fjcr ) and the accompanied exclusion of part of the initial interstitial liquid (decreasing Flin ). In addition, liquid percolation may give rise to a compositional re-equilibration of the initial primocrysts and migrating melts and significant modification of the initial bulk compositions, Cir. Nevertheless, in all of the open systems one can conceive of a stage when transport events cease. Beginning from this point, the percolating system is closed [Hunter, 1996], so that its whole chemistry defines the composition of the rock to be formed further in situ. It means that bulk modal and chemical compositions of the rocks produced during open system evolution can be also expressed by Equations (1, 2), but genetic interpretations of the values Fjcr, Flin, Cij(0), and Cil(0) must attribute them to one or more dynamic, probably late-stage, processes of reorganization.
If we assume the whole rock chemistry contains a record of the original rock formation parameters, there is the possibility of extracting such genetic information. Two main assumptions should be done to resolve the problem. The most important is that the liquid and crystals were in chemical equilibrium, with the formation temperature as an intensive parameter of the equilibrium. The second assumption is that there are samples having bulk compositions different from each other only due to variations of the modal proportions Fjcr and Flin, with the "trapped" liquid and temperature being the same.
If both assumptions are true, one can state that during equilibrium melting of samples having the same "formation temperature" the modeled liquid lines of descent will converge with falling temperature and intersect at a point in temperature-compositional space that represents the conditions under which the assemblage was formed [Frenkel et al., 1988a, 1988b]. Failure to find such an intersection would indicate the absence of the local thermodynamic equilibrium and a link between the values for Cij(0) and Cil(0). Thus, to decipher the genetic record of the rock it is necessary: (1) to select a group of samples for which one would expect the initial formation conditions to have been the same, (2) to conduct a series of physical or numeric equilibrium melting experiments for these samples, (3) to determine or calculate the chemical evolution of modeled liquids, (4) to construct T - X lines of this evolution for each sample and to identify the intersection point. In some aspects the approach is similar to the method of isotope geochronology in which an isochron is constructed by searching for an intersection point corresponding to the ages of the last isotopic equilibrium for both rocks and minerals [Ariskin and Barmina, 2000].
In practice, the method of geochemical thermometry is accomplished by means of the COMAGMAT phase equilibria model [Ariskin, 1999; Ariskin et al., 1993]. Although the program was primarily designed to simulate fractionation of basaltic magmas, it can also be used to calculate the course of equilibrium crystallization both for melts and cumulus rock compositions. Because equilibrium crystallization is reversible with respect to equilibrium melting, the construction of "melting" lines can be replaced with the crystallization trajectories for starting compositions identical to those of the rocks under consideration. Applying the method to natural assemblages, one should account for both analytical and computational uncertainties that yield formation temperatures with a determined error. As a rule, instead of an ideal intersection point (for a small number of modelled lines their convergence for each petrogenic oxide can be determined visually) one can see on the T - X diagrams a compact cluster of liquids lines of descent closing on one another in a range of 10-20o C that corresponds to the accuracy of the COMAGMAT model [Ariskin, 1999]. Our experience in the phase equilibria calculations for dolerites and diabases from differentiated sills [Barmina et al., 1989a, 1989b; Frenkel et al., 1989; Krivolutskaya et al., 2001], gabbros and troctolites from the Partridge River intrusion [Chalokwu et al., 1993, 1996] and large layered complexes [Ariskin et al., 2003; Barmina and Ariskin, 2002] has shown that the greater the contrast in compositions that are used, the more accurate the estimates of the trapped (interstitial) liquid temperature and composition that are obtained. This is confirmed by the results obtained for the rocks of the Skaergaard Marginal Border Series (Figure 2).
The results presented in this paper indicate that the determination of formation temperatures graphically, based on intersections of the modeled T - X trajectories (Figures 8-13), is somewhat subjective, because the approximate uncertainties of 10-15o C for this approach are close to the cited internal calibration accuracy of the COMAGMAT model. However, one must take into account the fact that the Layered Series rocks are differentiated. This imposes a set of important limitations on how one determines the original temperature and liquid composition for each sample: (1) the calcuated and observed "cumulus" mineral assemblage must correspond, (2) the modeled mineral compositions should be close (as much as possible) to those observed in selected rocks; and (3) melts entrapped in cumulates from higher levels of the Layered Series should represent lower temperature melts compared to those from underlying zones. Thus, the average values given in Table 4 were calculated for clusters of modeled T - X lines determined according to the petrologic restrictions given above. This does not allow one to specify strictly the accuracy of each calculated thermodynamic parameter, but provides an internally consistent approximation of T, log f O2, liquid component concentrations, mineral compositions, and original modal proportions.
The COMAGMAT programs rely upon mineral-melt equilibria geothermometers for olivine, plagioclase, augite, pigeonite, orthopyroxene, titaniferous magnetite, and ilmenite appropriate for a wide range of f O2 (from IW to NNO+1 buffers) and pressures from 1 atm to 10-15 kbar [Ariskin, 1999]. These semi-empirical dependencies were derived from a large dataset of experiments carried out mostly in tholeiitic systems covering compositions ranging from basalt to dacite. The general precision of phase equilibria calculations is about 10 to 15oC and 1 to 3 mol. % for each individual mineral. With such uncertainties, the use of mass-balance constraints in the COMAGMAT-3.5 model allows one to calculate equilibrium liquid compositions with an accuracy of 0.1 to 1 wt.%, which varies for different components and depends on the modal proportions between the melt and solids [Ariskin and Barmina, 1999, 2000].
A problem of application of COMAGMAT-3.5 to the Layered Series is the relatively low temperature mineral assemblages in LZc, MZ, and UZa rocks, ~1115-1080o C [McBirney, 1996]. In fact, the "silicate basis" of the COMAGMAT model was calibrated at higher temperatures of 1120-1350oC, with only a few points falling in the lower range. As a result, some calculations near 1100oC and below resulted in a shift of the modeled phase boundaries, e.g. producing an unrealistic enlargement of the Aug field at the expense of Pl and Pig. Moreover, calculated plagioclase compositions were found to be systematically more sodic and potassic than those observed in the natural assemblages. This forced me to reconsider and to modify the initial experimental dataset, drawing more attention to ferrobasaltic systems equilibrated in the vicinity of 1100oC.
To achieve this goal, we used the INFOREX experimental database [Ariskin et al., 1996, 1997], which now contains information from 290 melting experiment studies including over 10500 subliquidus runs. Searching the database for 1 atm experiments that resulted in saturation with Pl, Ol, Aug, and Pig oxides in the range 1050-1250o C (run duration 48 hours) yielded 270 experimental glass compositions containing 7 FeO 18 wt.%, 45 SiO 2 60 wt.%, and 2 Na2 O+K 2 O 5 wt.%. For most of the glasses (which are classified as ferrobasalts to ferrodacites - Figure 4), equilibrium mineral compositions have been determined, so that four subsets of 153 points for Ol, 187 for Pl, 125 for Aug, and 43 for Pig have been formed. That mineral-melt information was processed in the manner of Ariskin et al.  to develop a system of new mineral-melt geothermometers. The semi-empirical dependencies have been incorporated into COMAGMAT and tested by solving the inverse temperature-compositional problem, based on the initial dataset for each mineral. The new COMAGMAT program (version 3.65) produced a small improvement in the temperature calculations for Pl, Aug, and Pig (as compared to COMAGMAT-3.5), although modeled plagioclase compositions and phase relations between pyroxenes became more realistic. New olivine geothermometers resulted in the temperatures and compositions similar to those obtained in the older version and were left unchanged. Thus, the new "ferrobasaltic" COMAGMAT-3.65 model was used for phase equilibria calculations simulating equilibrium crystallization of the Skaergaard rocks.
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