Fluid inclusions in magnetite-apatite ore from a cooling magmatic system at El Laco, Chile

GFF volume 130 (2008), pp. 177–188. Article

Oxygen isotope composition of magnetite in iron ores of the Kiruna type in Chile and SwedenJAN OLOV NYSTRÖM1, KJELL BILLSTRÖM1, FERNANDO HENRÍQUEZ2, ANTHONY E. FALLICK3 andH. RICHARD NASLUND4

Nyström, J.O., Billström, K., Henríquez, F., Fallick, A.E. & Naslund, R.H., 2008: Oxygen isotope composition of magnetite in iron ores of the Kiruna type from Chile and Sweden. GFF, Vol. 130 (Pt. 4, December), pp. 177–188. Stockholm. ISSN 1103-5897.

Abstract: Magnetite-apatite iron ores of the Kiruna type, unaffected by deformation, have structures and textures similar to those of igneous rocks. The best examples are the El Laco deposits in northern Chile which resemble lava flows, pyroclastic deposits and dikes. El Laco magnetites have δ18O values between 2.3 and 4.2‰ (V-SMOW). Magnetite from ore with a magmatic texture has a mean of 3.7‰, and the mean for magnetite intergrown with pyroxene in veins is 2.4‰. Oxygen isotope data given here, fluid inclusion results and geological evidence indicate that ore formation took place in a cooling magmatic system. Major orebodies resembling lava flows and near-vent pyroclastic deposits crystallized from magma at ca. 1000°C. Fluids from cooling magma deposited magnetite and pyroxene (± apatite) at ca. 800°C in fissures and open spaces, now present as veins cutting major orebodies. There is little evidence for significant magnetite precipitation during hydrothermal conditions. A large province of magnetite-apatite iron ore in central Chile (the Cretaceous iron belt) and the Kiruna district in northern Sweden also contain primary ore of magmatic appearance. Major deposits in the Chilean iron belt and Kiruna contain magmatic-textured magnetites with the following δ18O means: Algarrobo = 2.2‰, Romeral = 1.2‰, Cerro Imán = 1.6‰, and Kiirunavaara = 1.5‰. We consider all oxygen isotope data for unoxidized, mag-matic-textured magnetite as representative of the Fe-rich magmas. Magnetites affected by hydrothermal alteration, recrystallization and subaerial oxidation have modified isotope signatures.

Keywords: Oxygen isotope, magnetite, iron ore, El Laco, Kiruna, Chile, Sweden.

1 Swedish Museum of Natural History, Box 50007, SE-104 05 Stockholm, Sweden; [email protected], [email protected] Departamento de Ingeniería en Minas, Universidad de Santiago de Chile, Casilla 10233, Santiago, Chile; [email protected] Isotope Geosciences Unit, SUERC, Scottish Enterprise Technology Park, East Kilbride, Glasgow G75 OQF, UK; [email protected] Department of Geological Sciences, SUNY, Binghamton, NY 13902, USA; [email protected] received 20 December 2007. Revised manuscript accepted 10 December 2008.

IntroductionSeveral studies of well-preserved magnetite-apatite ores of the Kiruna type suggest that they originated from iron-oxide magmas that erupted onto the surface or were emplaced at subvolcanic levels. Park (1961), Henríquez & Martin (1978), Nyström & Henríquez (1994) and Naslund et al. (2002) described Pliocene iron ores with volcanic structures and textures at El Laco, a vol-cano in northern Chile. The Oligocene iron deposit at Cerro de Mercado and related iron ores in Durango, Mexico, were inter-preted as volcanic deposits formed in a caldera by Lyons (1988). Magnetite lavas and tuffs occur in several places in the Chilean iron belt, an extensive province with numerous iron deposits of Cretaceous age along the Pacific coast, according to Travisany et al. (1995) and Nyström et al. (1996). Förster & Jafarzadeh (1994) reported Infracambrian [Neoproterozoic to Early Cam-

brian] diatremes with pyroclastic, extrusive and intrusive iron ore in the Bafq mining district of Iran. Moreover, Nyström & Henríquez (1994) gave evidence for a volcanic origin of the Paleoproterozoic Kiirunavaara deposit in Sweden, the type lo-cality of the apatite iron ores (Geijer 1931).

Alternative genetic models for the magnetite-apatite ores have been proposed. For a long time the discussion was focused on the iron deposits in the Kiruna district. The magmatic-intrusive model of Geijer (1931, 1967) and Frietsch (1978, 1984) was rejected by Parák (1975, 1984) who argued that the ores were deposited as exhalative-sedimentary bodies in a marine envi-ronment. However, such an origin is inconsistent with many observations (see Frietsch 1978; Nyström & Henríquez 1994). The presence of sedimentary structures and other features of the

ores offered by Parák (1975, 1984) as evidence for his model are easily explained by formation in a volcanic setting.

Some years ago the genetic discussion shifted to the iron de-posits at El Laco which generally are considered to be modern examples of the Kiruna ore type. Most geologists visiting the El Laco district have regarded the major orebodies as magnetite lava flows, but this interpretation has been disputed. Hitzman et al. (1992), developing an idea of Hildebrand (1986), suggested that the El Laco deposits were generated by subaerial precipita-tion from hydrothermal fluids or by hydrothermal replacement of viscous flows, such as rhyolites. Rhodes & Oreskes (1994, 1999), Sheets (1997) and Rhodes et al. (1999) argued that stable isotope, fluid inclusion and REE data, and field observations show that the ores are metasomatic-hydrothermal. A replacement origin has also been suggested by Sillitoe & Burrows (2002). However, fluid inclusion work of Broman et al. (1999) indicated that the main orebodies at El Laco formed in a volcanic system at temperatures above 800°C, with hydrothermal overprinting at lower temperature, consistent with the model of Nyström & Henríquez (1994, 1995) and Naslund et al. (2002).

Nyström & Henríquez (1994) presented major and trace ele-ment data of magnetite separated from many deposits in three apatite iron ore provinces, namely the El Laco district, the Chilean iron belt, and the Kiruna district. They found that the magnetites were similar chemically, with minor and trace element patterns differing from those of magnetites in stratabound sedimentary iron ores. The purpose of the present paper is to characterize the oxygen isotope composition of magnetite in ores of the Kiruna type based on samples from the three ore provinces, and to show that the δ18O values are consistent with deposition from mag-matic melts and magmatic fluids that evolved to hydrothermal solutions during cooling. The emphasis of the discussion will be on the El Laco deposits, since they are unaffected by metamor-phism and the best preserved examples of this ore type. The El Laco deposits have been studied by three of us (FH, HRN and JON) during several field seasons, before, in-between, and after the mining operations at Laco Sur during the 1990ʼs.

Investigated depositsThis study is based on the same magnetite separates that were analyzed for minor and trace elements by Nyström & Henríquez (1994). Magnetite separates from nine additional samples, not analyzed by these authors, are also included. In total, we have determined the δ18O ratios for 60 samples: 16 from El Laco and two other deposits in the High Andes, 25 from the Chilean iron belt, and 19 from Kiruna.

El Laco and other Andean depositsThe El Laco iron deposits in northern Chile (Park 1961; Rogers 1969; Haggerty 1970; Frutos & Oyarzún 1975; Henríquez & Martin 1978; Wegner 1982; Nyström & Henríquez 1994, 1995; Broman et al. 1999; Rhodes et al. 1999; Naslund et al. 2002; Sil-litoe & Burrows 2002; Henríquez et al. 2003) are situated on the flanks of a Pliocene volcanic complex at altitudes of 4600–5200 m. Within an area of 30 km2 there are seven deposits (Fig. 1A) with total resources of ca. 500 million tons of high-grade iron ore hosted by andesitic rocks. Geophysical studies indicate that a thin, shallow, magnetic layer extends for tens of kilometers to the north (Alva-Valdivia et al. 2003). The ores are largely composed of magnetite except in the upper part of the major deposits that

resemble lava flows where hematite predominates. Martitiza-tion along octahedral planes of magnetite and relict patches of magnetite in hematite demonstrate that the latter is an oxidation product. Fluor-apatite and pyroxene are usually subordinate min-erals. Fission track dating of apatite from the ore yielded an age of 2.1±0.1 Ma (Maksaev et al. 1988).

Although there are different opinions on the origin of the El Laco deposits, their appearance is not in dispute (cf. Rhodes et al. 1999; Naslund et al. 2002). Morphologically and structurally, the largest deposits (Laco Norte, Laco Sur and San Vicente Alto) look like short and relatively thick lava flows with associated near-vent pyroclastic material and dikes (Fig. 1 in Henríquez et al. 2003), and these terms will be used here in a descriptive sense for the sake of simplicity (Table 1) – no genetic meaning is intended. Another large orebody (San Vicente Bajo) has a dome-like form. These deposits have thicknesses of the order of 100 m or more and consist of >98 wt percent iron oxide.

The orebodies resembling lava flows have rubbly top surfaces like aa lava, but locally there are contorted flow structures (Fig. 5 in Nyström & Henríquez 1994), and even pahoehoe features. Vesicle-like cavities are common in the ore, which shows razor-sharp, seemingly chilled contacts towards andesite (Fig. 1 in Henríquez et al. 2003). Apatite is an accessory mineral occurring as interstitial grains and as needles in cavities and pores. In explo-ration pits at Laco Norte scoriaceous ore is intercalated between fresh andesite flows. Open pit mining at Laco Sur and drilling reveal that large parts of the orebodies consist of fragmental ore, described as pyroclastic by Nyström & Henríquez (1994) and Naslund et al. (2002). This ore is a friable, porous aggregate of fine-grained magnetite octahedra. A massive appearance with-out visible bedding is common, though at places the ore shows stratification, and even crossbedding, brought out by layers rich in apatite needles (Figs. 4A, B in Nyström & Henríquez 1994). One of the additional samples was separated from the depicted stratified ore. Interbedded fragments of magnetite with the size range of ash, lapilli and bombs (cf. Henríquez & Nyström 1998; Naslund et al. 2002; Henríquez et al. 2003) are exposed in road cuts beneath the main Laco Sur orebody.

The remaining three deposits (Rodados Negros, Laquito and Cristales Grandes) are small compared to the other deposits at El Laco. Rodados Negros and Laquito have dikelike form. The ore is massive and indistinguishable in texture from the ore in small magnetite dikes and veins that crosscut major deposits. Comb layering, defined by repeated layers of columnar magnet-ite dendrites with or without intergrown pyroxene and apatite (Nyström & Henríquez 1994), is locally conspicuous at Laquito and Rodados Negros. The pyroxene occurs as thin dendrites that commonly are altered to talc, silica, goethite and clay minerals, and the apatite forms euhedral, conical prisms. Ore of pyroclastic appearance has been found in a trench at Rodados Negros.

Cristales Grandes is a dike-vein system with prominent ore breccias and aggregates of magnetite, pyroxene and apatite, rich in quartz-lined open spaces. Pyroxene crystals are enclosed in and overgrown by apatite whereas the reverse has not been observed, indicating that apatite formed after pyroxene (Fig. 2 in Broman et al. 1999). Ore breccias also envelope the dike-like orebodies and San Vicente Bajo. The relationship between magnetite and pyroxene is discussed by Naslund et al. (2002) who concluded that in most places the two minerals crystallized simultaneously, after the main ore-forming process, as indicated by veins of magnetite intergrown with pyroxene cutting the ore at Laco Sur. Cristales Grandes appears to have formed at lower

178 Nyström et al.: Oxygen isotope composition of magnetite in iron ores of the Kiruna type GFF 130 (2008)

temperature than any of the other El Laco deposits, given the low homogenization temperatures and salinities of fluid inclusions in apatite here, and abundance of hydrothermal features (Broman et al. 1999; Rhodes et al. 1999).

Two alteration patterns are recognized at El Laco: minerals spatially linked to the ies, and a later hydrothermal alteration affecting large areas in the ore district spatially unrelated to ore-bodies. The first type is expressed principally as metasomatic ha-los of pyroxene in altered andesitic rock hosting ies and enclosed in ore. Nyström & Henríquez (1994), Broman et al. (1999), and Naslund et al. (2002) considered the pyroxene to be a product of the ore-forming process, whereas Rhodes et al. (1999) and Sillitoe & Burrows (2002) regarded the formation of pyroxene as an intermediate step in a multistage replacement process leading to magnetite mineralization. The second alteration pattern is hy-drothermal and of the same type found in other volcanic centers in the Andes. It is characterized by argillization, silicification,

propylitization and fumarolic activity represented by deposition of sulfates and sulfur (Gardeweg & Ramírez 1985).

The El Laco deposits are the best known members of a larger province of magnetite-apatite iron ore in the High Andes of northern Chile. Two other deposits belonging to this province are Incahuasi and Magnetita Pedernales, represented by one sample each in the study (Fig. 1B; Table 1). Incahuasi is situated at the southern foot of an eroded Late Miocene volcanic complex of the same name, 26 km south of El Laco. Fission track dating of the ore gave an age of 10.3±0.8 Ma (Maksaev et al. 1988). The ore occurs as apatite-bearing magnetite dikes and veins emplaced in a monzonitic porphyry. The two largest exposed dikes are up to 3 m thick and can be followed for 50 m. A few loose pieces of a highly vesicular ore type has been found which indicate that the dikes were emplaced close to the surface.

Magnetita Pedernales, located ca. 300 km south–southwest of El Laco, is a pluglike hill of magnetite ore hosted by dacitic

GFF 130 (2008) Nyström et al.: Oxygen isotope composition of magnetite in iron ores of the Kiruna type 179

Fig. 1. Sampled magnetite-apatite deposits in A. the El Laco district, and two other deposits situated in the High Andes of northern Chile. B The Cretaceous iron belt along the Pacific coast of central Chile. C. The Kiruna district in northernmost Sweden.

volcanic rocks of probable Early to mid-Tertiary age (Grez et al. 1991). The size of the orebody is uncertain due to scarcity of good outcrops and unexposed contacts with the host rock but it appears to be at least 200 m across at the surface. Magnetic measurements suggest that the form of the orebody is dikelike at depth.

Chilean iron beltThe ca. 600 km long Chilean iron belt (Geijer 1931; Park 1972; Bookstrom 1977; Oyarzún & Frutos 1984; Ruiz & Peebles 1988; Espinoza 1990; Nyström & Henríquez 1994; Travisany et al. 1995; Ménard 1995; Nyström et al. 1996; Oyarzún et al. 2003) is composed of several large and about 40 smaller magnetite-apa-tite deposits of Cretaceous age (Fig. 1B). Most of the deposits, including the large ones, occur associated with coeval andesites and basalts in tectonic contact with granitoids within the north-trending Atacama fault zone. Some deposits (e.g. Carmen) are hosted by volcanic rocks without presence of granitoids in the vicinity. The volcanic host rocks are affected by burial metamor-phism, usually at prehnite-pumpellyite facies (Levi et al. 1988). The intensity of the alteration, reflected by recrystallization and extensive development of amphibole, is higher near the orebod-ies, probably due to processes related to the formation of the ores. The original textures of the volcanic rocks are preserved outside shear belts in the fault zone. Limestone, recrystallized to garnet-bearing marble, occur close to and stratigraphically below one of the deposits (El Tofo). No carbonate rock have been reported to underlie any of the other sampled deposits.

The iron belt is represented here by four major deposits (Algar-robo, Romeral, El Tofo, Cerro Imán; Table 2) and three smaller deposits (Carmen, El Dorado, Ojos de Agua; Table 3). The ore-bodies tend to be elongated parallel to the Atacama fault zone. Most of them have tabular shapes; some are domelike. The ores show discordant as well as concordant relations with the host rock and the contacts are sharp. An upwards increasing extent and intensity of hydrothermal alteration (Naslund et al. 2002), and presence of fumarolic deposits of silica in topographically high parts of some ore districts (e.g. at Algarrobo) are consistent with a subvolcanic character of the ores. Travisany et al. (1995) reported the occurrence of ore lavas and pyroclastic ore in some districts.

The major deposits consist predominantly of massive mag-netite ore with low contents of amphibole and apatite gangue. However, actinolite altered to chlorite, talc and other phases, and at places coarse prisms of apatite, are common in marginal parts of the smaller deposits, locally giving rise to comb lay-ering. The grain size of magnetite varies; coarse crystals tend to be more common in the upper part of the deposits. Primary ore textures such as magnetite of columnar habit and dendritic pseudomorphs after pyroxene or amphibole have been found in almost all deposits (Figs. 8E and 9B, C in Nyström & Henríquez 1994). A search for dendritic minerals at El Tofo failed to reveal primary ore textures. The ore typically forms granular aggregates of polyhedral magnetite crystals suggestive of recrystallization (Nyström & Henríquez 1994). Hematite and small amounts of maghemite are widespread in samples from what appears to have been the original top of the ore. These minerals are probably oxidation products after magnetite as suggested by ubiquitous relicts of this phase. Pyrite occurs locally as dissemination in the deposits (e.g. in samples ALG-3, ROM-59 and ROM-58). A variety of ore types ranging from massive high-grade ore to ore

breccia and mineralized rock were sampled in one of the deposits (Romeral; Table 2).

The smaller and often apatite-rich deposits can be regarded as intermediate in character between major deposits and steep apatite-actinolite-magnetite veins. Such veins occur in great numbers and constitute swarms in the iron districts. Large veins, which have been exploited for apatite down to a depth of 100 m or more, are 30–800 m long and 1–10 m wide. Texturally, the veins are pegmatites, and a pegmatitic appearance also char-acterizes marginal parts of the ore in Carmen, El Dorado, and Ojos de Agua (Geijer 1931; Dobbs & Henríquez 1988). Two large apatite veins have been sampled (La Escoba and Yayita in the Romeral district; Table 3). A limestone-hosted stratiform iron deposit that is not of Kiruna type, situated at the eastern margin of the Chilean iron belt (Bandurrias, Fig. 1B; Cisternas 1986; Espinoza 1986), is included in the study for comparison (Table 3). One of the Bandurrias samples is carbonate-banded, whereas the other sample lacks a primary texture, possibly due to recrystallization.

Kiruna districtThe magnetite-apatite ores of the Kiruna district in northernmost Sweden (Geijer 1910, 1931, 1967; Parák 1975; Frietsch 1978; Nyström & Henríquez 1994; Romer et al. 1994; Harlov et al. 2002) are steeply dipping tabular bodies (Fig. 1C) occurring con-cordantly in volcanic rocks. The giant Kiirunavaara deposit, with pre-mining reserves of ca. 2 billion tons of high-grade iron ore, formed in the age interval from ca. 1900 to ca. 1880 Ma (Cliff et al. 1990). It is the most well-preserved deposit in the district. The footwall consists of trachyandesitic lava that traditionally is called syenite porphyry and the rocks of the hanging wall are rhyolitic ignimbrites and tuffs referred to as quartz porphyry. Presence of primary textures in many parts of the Kiirunavaara deposit (see Nyström & Henríquez 1994, and references therein) are consistent with a lack of penetrative deformation in sur-rounding volcanic rocks which are burial metamorphosed. The secondary mineral assemblages in the volcanic rocks correspond to greenschist facies.

Fifteen of the Kiruna samples come from Kiirunavaara (Table 4). In addition to the magnetites analyzed chemically by Nys-tröm & Henríquez (1994), the present study includes a sample of what these authors interpreted as vesicular magnetite lava and four samples of magnetite from amygdules in the footwall por-phyry. These amygdules are locally abundant in discontinuous zones up to many meters away from the ore contact; they are absent in the hanging wall. The amygdules consist of magnetite and/or actinolite; other minerals observed are titanite, and minor apatite, quartz, calcite, biotite (±chlorite) and pyrite. Some of the large amygdules are only partly filled. Similar amygdules have been reported from two other iron deposits of Kiruna type in the region by Lundberg & Smellie (1979).

The Per Geijer deposits (Rektorn, Henry and Nukutusvaara; Fig. 1C), described by Parák (1975), are located higher up in the volcanic pile. They are relatively rich in apatite. Relict primary textures occur (samples PF-HE and NUK-9; Table 4) but large portions of the ores are deformed with a foliated to gneissic appearance and apatite-banding of tectonic nature. The magnetite is partly to strongly oxidized to hematite. Much silica and carbonate, now present as quartz and ankerite in the ore, was seemingly introduced during an episode of alteration. These changes are pervasive which suggests that the deformation and


alteration took place together. Four samples from the Per Geijer deposits have been analyzed.

Analytical procedureThe magnetites were concentrated by crushing, washing in distilled water, and magnetic separation. Three of the magnetite separates from the Kiruna district contain considerable amounts of apatite (PG-235, PG-K5 and PG-531; Table 4), but generally the content of this mineral is less than 2 wt percent in separates from apatite-rich ore types and below detection level in other ore types. Given that apatite fractionates in a manner similar to oliv-ine (Taylor 1967), the fractionation equation 1000 lnαol-mt = 2.62 × 106/T2 can be used to estimate the effect of apatite impurities on the obtained magnetite data. For example, at 600°C, which we consider to be at the lower end of the depositional temperature range for the El Laco deposits, a 2 wt percent apatite contamina-tion will raise the δ18O value with 0.1 per mil. Any apatite im-purities in magnetite will thus have very limited influence on the measured oxygen isotope signatures. Duplicate analysis of untreated, apatite-bearing portions and HCl-leached, apatite-free portions of two samples gave similar δ18O values: 1.6‰ (un-treated) compared with 1.8‰ (leached) for PG-235, and 2.2‰ (untreated) compared with 2.3‰ (leached) for PG-531.

Small amounts of actinolite or pyroxene, and quartz are present in a few samples (in total up to 4 wt percent in ROM-61, PG-K14 and EL-C:24); the other samples contain <1 wt percent of sili-cates as a rule. The percentage of hematite in the samples (% hm) was determined with X-ray diffraction, using a calibration curve based on the ratio of the 2.69Å (hematite) and 2.10Å (magnetite) peak areas for mixtures of pure magnetite and hematite in differ-ent proportions from the investigated deposits.

The oxygen isotope work was carried out at SUERC in Scot-land. Most of the samples (i.e., the samples included in Nyström & Henríquez 1994) were analyzed according to the conventional fluorination procedure of Clayton & Mayeda (1963), as modi-fied by Borthwick & Harmon (1982) for use with ClF3. The 1σ precision for isotopically homogeneous material is better than ±0.2‰, with the NBS 28 quartz standard giving 9.6‰. The laser technique (cf. Fortier et al. 1995) was used for the nine additional samples which were all run in duplicate with a reproducibility that slightly exceeded 0.1 per mil. The laser data for unmetamor-phosed Andean deposits fall at the high end of the conventional data range for El Laco (Table 1). However, this potential meth-odology effect is likely to be at the most a few tenths of per mil and the order of magnitude of data displacement will not affect the conclusions discussed below. All oxygen isotope data deter-mined in this study (Tables 1–4) and quoted from other sources are reported as delta values in per mil relative to V-SMOW.

Results: δ18O values of the iron oxideThe δ18O values obtained for magnetite from the El Laco depos-its are given in Table 1. Samples from deposits that according to Rhodes et al. (1999, p. 323) show “ostensibly magmatic tex-tures (bladed [columnar], fine-grained, massive, and vesicular magnetite)” will be referred to here as magmatic-textured ore samples. They are listed in the upper part of Table 1 (samples FHL-121 to FHL-18), followed by three samples from orebodies of dikelike form (Rodados Negros and Laquito), and ending with three samples from the Cristales Grandes dike-vein system. The analyzed magnetites have δ18O values that range from 2.3 to 4.2


per mil (Fig. 2), with a mean of 3.0±0.6‰ (1σ; n = 14). Within this limited range small differences can be discerned between magnetites that on textural grounds and fluid inclusion work (Broman et al. 1999) may be assumed to represent the high and low-temperature ends of crystallization, respectively.

The average for magmatic-textured magnetites with a maxi-mum of 5 wt percent hematite is 3.7±0.4‰ (Table 1: samples ELS-2, EL-2:6 and FHL-126). Samples containing more hema-tite are excluded, since oxidation lowers the δ18O value of the iron oxide, as shown in the next section. The magnetites from Cristales Grandes, the lowest-temperature deposit at El Laco (Sheets 1997; Broman et al. 1999), have more uniform and lower δ18O ratios (mean = 2.4±0.1‰). Since hydrothermal features are abundant at Cristales Grandes the magnetites from this dike-vein system are regarded as ʻhydrothermalʼ. The difference between the ̒ magmatic ̓and ̒ hydrothermal ̓means is 1.3 per mil. Samples from the dikelike Rodados Negros and Laquito orebodies that show evidence of deposition from both melts and hot fluids are avoided in this calculation in order to enhance the contrast. Their δ18O ratios fall between the ʻmagmatic ̓and ʻhydrothermal ̓val-ues as should be expected from the mixed geological signature of the two deposits. The two other Andean deposits, Incahuasi and Magnetita Pedernales, have δ18O ratios close to the average for magmatic-textured El Laco magnetite (Table 1).

The δ18O values for the major deposits in the Chilean iron belt are listed in Table 2. They are significantly lower than the results for El Laco, with only a small overlap between the two data sets (Fig. 2). The magmatic-textured magnetites from the iron belt have the following means: Algarrobo = 2.2±0.3‰ (all samples except pyrite-bearing ALG-3), Romeral = 1.2±0.5‰ (samples ROM-60 and ROM-66), and Cerro Imán = 1.6±0.1‰ (both samples). Averaging the three individual means gives a value of 1.7±0.5‰ for the iron belt. Magnetite samples with disseminated pyrite (ALG-3, ROM-59 and ROM-58) have lower δ18O ratios than magmatic-textured magnetites from the same district (Ta-ble 2). Four Romeral samples of uncertain type are disregarded. The mineralized rock samples (ROM-63 and ROM-64) and the very low-grade disseminated ore (ROM-61) contain much more Cr than the other samples (Nyström & Henríquez 1994), and it cannot be excluded that part of the magnetite is inherited from the rock. Sample ROM-62 is not magmatic-textured. The recrystallized El Tofo magnetite has an anomalously high δ18O value (4.5‰).

The variation in δ18O values for the smaller apatite-rich ores and the veins (from 0.1 to 3.0 per mil) is similar but wider than the range for major deposits, disregarding the anomalous El Tofo sample (from 0.8 to 2.5 per mil; Fig. 2, Table 3). The magnetites of the stratiform Bandurrias deposit have 3 to 4.5 per mil higher δ18O values than the mean for the deposits of the iron belt.

The δ18O values for the Kiruna magnetites (Table 4) are slightly lower than those for the Chilean iron belt, and there is no overlap with the El Laco data (Fig. 2). The first ten magnetites listed in Table 4 may be regarded as magmatic-textured. Exclusion of partly oxidized samples gives a mean of 1.5±0.7‰ (n = 5). The magnetite occurring in amygdules in the footwall, deposited from fluids, has a mean that is ca. 0.6 per mil lower (0.9±0.4‰; n = 4). All the samples from the deformed Per Geijer deposits, including vein sample PG-618 from Kiirunavaara, are moder-ately to strongly oxidized and have negative or near-zero δ18O values. Sample KRE-1 from Rektorn, separated from foliated ore rich in secondary ankerite and quartz, has the lowest δ18O value measured in this study (–3.3‰; Fig. 2).

Table 2. δ18O values (‰ V-SMOW) for magnetites from four major deposits in the Chilean iron belt.Sample δ18O Ore type according to texture % hmAlgarroboALG-21 1.9 Columnar magnetite in massive ore 0ALG-25 2.3 Massive ore with dendritic ʻpyroxeneʼa 0ALG-20 2.5 Brecciated ore 0ALG-3 0.6 Ore with dendritic ʻpyroxeneʼa and some pyrite 0RomeralROM-60 1.6 Massive ore, Cuerpo Principal (=CP) 0ROM-66 0.8 Columnar magnetite with dendritic ʻpyroxeneʼa, CP 0ROM-62 1.3 Ore breccia with apatite, CP 0ROM-61 1.7 Low-grade ore (magnetite dissemination), CP 0ROM-59 –1.3 Pyrite-rich ore vein in ore breccia, CP 0ROM-58 0.4 Recrystallized ore with pyrite, Romeral Norte 75ROM-63 2.0 Ore patches in mineralized, altered diorite, Siciliano 70ROM-64 0.3 Ore patches in mineralized, altered diorite, Siciliano 45El TofoTOF-40 4.5 Recrystallized ore with saccharoidal texture 0Cerro ImánCIM-5 1.6 Columnar magnetite in massive ore 4CIM-9 1.7 Columnar magnetite in massive ore 0a ʻpyroxene ̓= pseudomorphs composed mainly of actinolite and talc after pyroxene or amphibole.

Table 1. δ18O values (‰ V-SMOW) for magnetites from El Laco and two Tertiary iron deposits in the High Andes of northern Chile.Sample δ18O Ore type according to texture % hmEl LacoFHL-121 3.2 Volcanic ore bomb, east of San Vicente Altoa 15ELS-2 4.2 Apatite-banded fallout ore tuff, Laco Sur (from mine) a, b 1FHL-82 2.7 Pyroclastic ore, San Vicente Altoa 30FHL-101 3.4 Pyroclastic ore, Rodados Negrosa 25EL-2:6 3.4 Vesicular ore, Laco Sur (from drill core; depth 50 m)a 0FHL-67 2.3 Ore with ropy surface, Laco Nortea 12FHL-126 3.7 Ore at vug with apatite, San Vicente Bajoa 5FHL-18 3.0 Ore with pyroxene layers, Laco Sur 75FHL-105 3.1 Columnar magnetite, Rodados Negrosc 1FHL-106 3.5 Columnar magnetite with pyroxene, Rodados Negrosc 7FHL-76 2.3 Columnar magnetite, Laquitoc 13FHL-109 2.6 Spherulitic ore, Cristales Grandes 3FHL-113 2.4 Pyroxene-rich ore, Cristales Grandes 0EL-C:24 2.4 Pyroxene-apatite-rich ore breccia, Cristales Grandes 2Other Andean depositsINC-2 3.4 Radiating platy dendrites, Incahuasi b 15CMP-13 4.0 Columnar magnetite, Magnetita Pedernales b 20a Magmatic-textured ore.b Analyzed with laser technique.c From ore dike with comb layering.


Factors influencing the stable oxygen isotope ratios of magnetiteThe δ18O value of magnetite is principally controlled by the depositional temperature and the isotope signature of the fluid or magma from which magnetite crystallizes. Issues with respect to the magnetite-water oxygen isotope fractionation factor (α(mt-w)) have been discussed by Fortier et al. (1995) and Zheng (1995). Whereas the former offer a well-constrained value of around –8.5 for 103lnα(mt-w) at 350°C, the latter suggests that different mineral structures during magnetite crystallization result in dif-ferent fractionation factors. The value determined by Fortier et al. (1995) is consistent with that expected (–8.8) for the spinel structure in Zhengʼs (1995) scheme. For the magnetites dis-cussed in this paper, we propose that the crystallization structure was rather the inverse-spinel one, and so adopt Zhengʼs (1995) calculated magnetite-water fractionation factor via:

103lnα(mt-w) = 2.88 × 106T–2 – 11.36 × 103T–1 + 2.89

where T is in kelvins.

Magmatic fluids have δ18O values in the 5.5–9.5 per mil range (Shephard et al. 1969) with most values in the 6–8 per mil in-terval. Assuming isotopic equilibrium, magnetite that grew from an aqueous fluid at temperatures of 800–850°C will have a δ18O value about 5 per mil lower than that of the fluid. At temperatures of 400–500°C the slow rates of exchange of oxygen between mineral and fluid means that isotopic equilibrium may not be attained (Freer & Dennis 1982). A hydrothermal fluid can have a very variable oxygen isotope composition dependent on the ultimate source of its components. In addition, interaction with the rock pile through which the fluid passes before deposition can modify its isotope signature.

The δ18O values given here for magnetites from magmatic-tex-tured ore (Fig. 2) coincide with the range for magnetite from ig-neous rocks reported in the literature (e.g. Taylor 1968). Accord-ing to Taylor (1967, p. 134) ”igneous magnetite is very uniform in O18/O16 ratio (δ18O = +1 to +4‰), and any magnetite deposit with δ values appreciably outside this range must have a differ-

Table 4. δ18O values (‰ V-SMOW) for magnetites from Kiruna, Sweden.Sample δ18O Ore type according to texture % hmKiirunavaaraPG-531 2.2 Apatite-bearing ore with apparent cross-bedding 0PG-235 1.7 Stratified ore rich in apatite 25KIR-51 1.1 Vesicular orea 15PG-K9 1.8 Ore with dendritic actinolite pseudomorphs after pyroxene 12PG-K14 0.9 Ore with dendritic actinolite pseudomorphs after pyroxene 14PG-K8 2.1 Skeleton ore 0PG-37:6 0.7 Skeleton ore 0HjL-1 1.3 Columnar magnetite 0KUJ-3 1.1 Columnar magnetite 0PG-530 –0.2 Columnar magnetite 30PG-618 –0.7 Magnetite-banded apatite rock (foliated vein) 30KIR-86 1.0 Ore amygdules from the footwall porphyrya 0KIR-87 1.1 Ore amygdules from the footwall porphyrya 04564:240 0.3 Ore amygdules from the footwall porphyrya 04564:297 1.2 Ore amygdules from the footwall porphyrya 0Per Geijer depositsNUK-9 0.2 Pyroclastic (?) ore, Nukutusvaaraa 15PF-HE –0.8 Skeleton ore, Henry 35PG-K5 –0.4 Apatite-banded ore with folding, Rektorn 75KRE-1 –3.3 Apatite-ankerite-quartz-banded ore, Rektorn 80a Analyzed with laser technique.

Table 3. δ18O values (‰ V-SMOW) for magnetites from three apatite-rich iron deposits, two apatite veins, and a stratiform exhalative-sedimentary deposit (Bandurrias) in the Chilean iron belt.Sample δ18O Ore type according to texture % hmCarmenCMN-12 2.6 Columnar magnetite with apatite and ʻpyroxeneʼa 10CMN-13 2.4 Ore with apatite 20CMN-14 0.9 Columnar magnetite with apatite 50El DoradoDOR-2 0.1 Columnar magnetite in apatite rock 30DOR-3 3.0 Apatite-rich ore 0Ojos de AguaODA-31 0.5 Columnar magnetite with actinolite 40Apatite veinsYAY-2 1.5 Apatite-magnetite rock 30ESC-14 1.1 Apatite-actinolite-magnetite rock 6BandurriasBAN-1 6.3 Massive recrystallized ore 15BAN-3 4.8 Carbonate-banded ore 50a ʻpyroxene ̓= pseudomorphs composed mainly of actinolite and talc after pyroxene or amphibole.


ent origin or have been recrystallized and altered subsequent to its formation". Samples analyzed by us with ratios falling outside this range are with a few exceptions either oxidized, seemingly affected by hydrothermal alteration expressed by presence of disseminated pyrite, or recrystallized (Tables 1–4).

A comparison of El Laco samples from natural outcrops, min-ing exposures, and drill cores shows that the hematite content generally increases toward the original surface, faults and other zones of local oxidation in deposits resembling lava flows. Oxi-dation lowered the δ18O value of the iron oxide, as illustrated by ore that on textural grounds has been described as pyroclastic (Table 1: mine sample ELS-2 vs. surface samples FHL-82 and FHL-101). This indicates that still hot magnetite interacted with an isotopically light fluid, perhaps derived from melting snow.

Rhodes & Oreskes (1999) reported the δ18O composition for 39 samples of iron oxide from El Laco determined by laser fluorina-tion. The δ18O values range from +5.2 to –8.9 per mil, decreasing virtually linear as the hematite content increases from 0 to 96 wt percent. Disregarding the most oxidized samples, their δ18O ratios are similar to ours although shifted to somewhat higher values for unoxidized magnetite. Magmatic-textured samples

with a maximum of 5 wt percent hematite (samples LS 9404-fgmt, LS-52, SVB 9402, SVB 9403A and SVB 9403C) give a mean of 4.7±0.3‰, and the two ʻhydrothermal ̓magnetites from Cristales Grandes give 3.8±0.4‰, yielding a difference of 0.9‰. The difference coincides with our value (1.3‰; the difference between corresponding means of 3.7 and 2.4‰). Consistent with our results, the magnetites from Rodados Negros and Laquito show intermediate values, though with some overlap in the ʻhy-drothermal ̓range.

There also seems to exist a relationship between low δ18O val-ues and oxidation in the Chilean iron belt (Table 3: the Carmen and El Dorado deposits) and the Kiruna district (Table 4). The low δ18O obtained for the pyrite-rich ore breccia ROM-59 (–1.3‰) is consistent with hydrothermal activity. Fluid-rock interaction during metamorphism is another factor to consider: the δ18O ra-tios are lower for the Kiruna district that is metamorphosed at a higher grade than the Chilean iron belt. Deformation, promoting fluid-rock interaction, might be a contributing factor for the low ratios for banded ore and apatite rock at Kiruna (Table 4). In con-trast, the unmetamorphosed El Laco deposits have significantly higher δ18O ratios than Kiruna and the Chilean iron belt.

The anomalously high δ18O value for the El Tofo magnetite (+4.5), its recrystallization, and the occurrence of marble a short distance beneath the orebody are probably related. The simplest explanation is interaction between 18O-enriched metamorphic fluids derived from the marble and recrystallizing magnetite. A similar process is suggested for Bandurrias, the limestone-hosted stratiform deposit included for comparison. The recrystallized Bandurrias sample has a higher δ18O value (6.3‰) than the sam-ple with preserved primary texture (4.8‰; Table 3).

A hydrothermal vs. magmatic origin for the magnetite oresRhodes & Oreskes (1999) reported that magnetites with magmatic and hydrothermal textures at El Laco have similar δ18O ratios and therefore must be formed by the same process. They concluded that this process was hydrothermal, since it is implausible that ”magnetite intergrown with quartz in veins” crystallized from a magma. The volcanic textures of the ore are explained as fea-tures inherited from an andesite precursor replaced by ore. The replacement is believed to have taken place in several stages, with pervasive calcic metasomatism expressed by formation of diopside prior to magnetite mineralization.

We consider another interpretation more likely. The similar isotope ratios falling within the magmatic range suggest that the

deposits are what they appear to be: products of iron oxide mag-mas and hot magmatic fluids that crystallized within a relatively small temperature interval. Textural relations demonstrate that quartz is a late phase deposited after magnetite, not intergrown with it. Rhodes & Oreskes (1994) cited field and petrographic observations to suggest that the fluids precipitating quartz were in equilibrium with magnetite. Subsequently, it was realized that the two phases were not formed in equilibrium (Rhodes & Oreskes 1999).

Metasomatic replacement as described by Rhodes & Oreskes (1999) generally destroys textures (Naslund et al. 2002; Hen-ríquez et al. 2003). It is very unlikely that the volcanic textures of the ore are inherited from a silicate protolith. Vesicles are uncom-mon in the andesites at El Laco, whereas unfilled pores and open spaces are very common in the ore (Sillitoe & Burrows 2002), and they show no evidence of ever having been filled. None of the volcanic rocks hosting the orebodies have textures similar to those in the ore.

We argue that the major orebodies that resemble lava flows were emplaced at temperatures above 800°C. Laco Sur can serve as an example. Magmatic-textured ore here is crosscut by veins of intergrown pyroxene-magnetite. Hydrosaline melt inclusions in pyroxene homogenize at >800°C (Broman et al. 1999), suggest-ing that pyroxene crystallized from hot fluids in fissures during degassing and cooling of the orebody. This is supported by data of Sheets (1997) who reported homogenization temperatures of


Fig. 2. δ18O values for mag-netites from three provinces of magnetite-apatite ore in Chile and Sweden. Samples plotted as magmatic have tex-tures that without discussion would be accepted as igneous, were it not for the extremely iron-rich composition of the material. Magmatic sam-ples referred to as oxidized contain >5 percent hematite. Pyrite-disseminated samples in the iron belt are referred to as hydrothermal.

820–840°C for melt inclusions in pyroxene from San Vicente Alto and Laco Norte.

Fluid inclusion data for the Cristales Grandes dike-vein system present a more complex picture. The pyroxene contains hydro-saline melt inclusions that almost, but not completely, homog-enized at 800°C (Broman et al. 1999; Sheets (1997) recorded 710–740°C). In contrast, the fluid inclusions in apatite suggest formation at much lower temperature. Sheets (1997) reported homogenization temperatures between 111° and 401°C° with the majority of the data in the range 250–350°C for Cristales Grandes and Rodados Negros combined. According to Sillitoe & Burrows (2002), the apatite temperature (250–350°C) corresponds to the ore formation at El Laco, whereas the pyroxene temperature (710–840°C) reflects premineralization Ca metasomatism. However, a wider range of apatite crystallization temperatures is obtained if the data of Broman et al. (1999) also are taken into account. They reported a minimum temperature of 400–505°C for Rodados Negros, homogenization temperatures between 69° and 325°C for Cristales Grandes with most values clustering in the range 130–170°C, and a minimum temperature of 540°C for melt inclusions, similar to those in pyroxene, in an early genera-tion of apatite from San Vicente Bajo. It must be emphasized that these apatite temperatures are only minimum values for the ore formation since the occurrence of apatite in pore spaces and cavities in major deposits indicates that it crystallized after the main mass of magnetite. Moreover, the F-rich composition of the El Laco apatite argues against a general formation at low temperature (Naslund et al. 2002).

Given a temperature of ca. 800°C for pyroxene-magnetite intergrowths at both Laco Sur and Cristales Grandes, a rough estimate can be obtained for the crystallization temperature of Laco Sur and other major orebodies which according to cross-cutting relationships must have formed prior to the intergrowths. A calculation based on the 1.3 per mil difference between the magmatic-textured magnetite and the fluid-deposited magnetite (δ18O means of ca. 3.7 and 2.4 per mil, respectively; the latter represented by the samples from Cristales Grandes) gives a value of ≤1100°C. Using instead the 0.9 per mil difference obtained from the data of Rhodes & Oreskes (1999) gives ≤1000°C. These estimates suggest that the ore magma was about 200°C hotter than the fluids depositing magnetite at Cristales Grandes (Table 5).

In our opinion, only insignificant amounts of magnetite at El Laco are hydrothermal and formed below 500°C. However, be-fore rejecting a general hydrothermal origin for the ore deposits, it is relevant to address the question: what is the magnitude of isotopic differences expected for magnetite formed in magmatic and hydrothermal environments? The effect of a Rayleigh-type

process, e.g. depletion of the light isotope (16O) due to preferen-tial incorporation into magnetite, is likely to be small in a system where the availability of the element of concern, oxygen, is not a limiting factor. Instead, the main factor to consider in a cooling magmatic system where magnetite crystallizes is the effect of temperature on the equilibrium between magnetite and fluid.

Let us assume a situation with ore fluids of magmatic origin at a temperature of 1000°C for the magmatic case and 400–500°C as representative of hydrothermal conditions. Using the frac-tionation factors discussed earlier, magnetites formed in the hydrothermal range are likely to be ca. 3 per mil lower in δ18O than magmatic-textured magnetites. Given that the δ18O range between magmatic-textured and ʻhydrothermal ̓magnetites only is of the order of one per mil at El Laco, it suggests that the analyzed magnetites have a common origin and formed during a quite limited temperature span. Accepting that magnetites with magmatic textures have δ18O ratios consistent with a magmatic origin it seems logical to suggest that also magnetites with a more hydrothermal appearance are magmatic, although deposited from high-temperature fluids rather than melts (Table 5).

An analogous conclusion was drawn by OʼFarrelly & Rickard (1989) who analyzed oxygen isotope compositions of magnetite from Kiirunavaara and reported a unimodal distribution with a mean δ18O value of 1.5 per mil. This is identical to our results for unoxidized magnetite of magmatic texture from the same de-posit (1.5‰). OʼFarrelly & Rickard (1989) concluded that their data are consistent with a high-temperature origin (600–700°C) for the magnetite in equilibrium with water with δ18O of 7 to 8 per mil, using the fractionation equation provided by Friedman & OʼNeil (1977). The Zheng (1995) inverse spinel calibration would suggest a fluid δ18O 1 per mil higher. Blake (1992) deter-mined the oxygen isotope composition of the volcanic host rocks in the footwall and hanging wall of the Kiirunavaara orebody and reported a tight grouping of data between 6 and 9 per mil. He concluded that the fluids accompanying magnetite emplacement have not affected the narrow magmatic signature of the host rocks, and suggested a common magmatic origin of rocks and ore. The large reservoir of magmatically derived oxygen within the orebody might have acted as an internal buffer reducing the effects of any later alteration.

Recently, Edfelt et al. (2007) reported oxygen isotope data for the Tjårrojåkka Fe-Cu deposit, situated some 45 km SW of Kiirunavaara. The data range was found to be similar both for magnetite occurring in the Cu-rich ore (2.1 to –0.5 per mil) and in the spatially associated Fe-rich ore (0.9 to 0.1 per mil). Given that the estimated formation temperatures were as low as 410–660°C for the magnetite stage (based on oxygen isotope data for mineral pairs), and that there are similarities between the

Table 5. Summary of the main ore- and mineral-forming stages at El Laco.Stage I II III IVMineral Magnetite Magnetite ± Quartz a Hematite pyroxene, apatite

Occurrence Lava flows, pyroclastic Dikes (partly), veins, Veins, Oxidation deposits, dikes (partly) ore breccia coating in open spaces productb

Deposition from Iron-oxide magma Cooling magmatic fluid Hydrothermal waterc Meteoric water

Approximative temperature 1000°C 800°C <500°C <100°C

δ18O of source +9‰ +9‰ Ca. 0 and +20‰c ≤0‰a Apatite crystals are occasionally enclosed in quartz; drusy quartz may contain magnetite ʻdustingʼb Primary hematite occurs in some pyroclastic depositsc Of meteoric origin, in part modified by evaporites


two ore systems with respect to fluid composition, radiometric age and stable isotope data, it was proposed that the Tjårrojåkka Fe-Cu ores are truly hydrothermal in origin.

Implications of δ18O data for quartz and pyroxene from El LacoQuartz associated with magnetite in ore veins and breccias, and altered andesitic host rock at El Laco have significantly higher δ18O values than the range for hydrothermal quartz reported in the literature, to the best of our knowledge. Rhodes & Oreskes (1994, 1999) reported a wide range of δ18O values for quartz, from 6.0 to 27.9 per mil, with many values >20‰. In addition, there is a large spread in ratios within a given quartz vein. An important feature, not mentioned but evident from their data, is a bimodal distribution with few values in the range 12 to 18 per mil. Their results are in good agreement with the high δ18O values determined by Vivallo et al. (1994) for hydrothermally altered andesites and quartz in veins and hydrothermal breccias from El Laco, with a maximum of 35.3 per mil for an altered andesite. Rhodes & Oreskes (1999) proposed that an 18O-rich, saline fluid that has interacted with an evaporitic source would explain the high oxygen isotope ratios and be responsible for the iron mineralization. Mineralization was postulated to be one of the stages in a sequence of hydrothermal alteration and replace-ment events that transformed andesite to ore.

Considering the model of Rhodes & Oreskes (1999), magnet-ite shows a narrow δ18O range (3.5–5.2 per mil) compared to the extremely variable quartz data. In the hypothetical case of isotopic equilibrium, the δ18O values for quartz should be around 10–14 per mil higher than the values for coexisting magnetite at typical hydrothermal temperatures, say 300–500°C. For El Laco this would mean quartz data in the 14 to 19 per mil range which obviously is not the case. The lack of equilibrium means that magnetite and quartz must have been deposited from different fluids. Still, it is remarkeable that neither of the two indicated quartz-forming fluids (with δ18O calculated to be near zero and around +20 per mil, respectively) show any isotopic similarity with the inferred magnetite-forming fluid (+11 to +13 per mil) at temperatures of ca. 400˚C. We accept the opinion of Rhodes & Oreskes (1999) that a fluid that has interacted with an evaporitic source can explain the high oxygen isotope ratios of quartz and altered rocks. However, the population of low δ18O quartz would require a surface-derived fluid source with an oxygen isotopic composition that is close to zero per mil. Thus, although part of the variable δ18O quartz data could be attributed to temperature changes, it seems that quartz must have formed from fluids of different origin (Table 5).

The pyroxene analyzed by Rhodes & Oreskes (1999) has δ18O values between 7.1 and 8.9 per mil with no quantitative distinction between data for pyroxene occurring in altered an-desite and pyroxene intergrown with magnetite (values at the low end for Cristales Grandes). In two of their samples from Cristales Grandes, pyroxene with δ18O values of 7.6 and 7.3 per mil, coexists with magnetite of 3.5 and 4.1 per mil, respectively. By applying the diopside-magnetite geothermometer equation of Matthews et al. (1983), these mineral pairs suggest depositional temperatures between 720° and 850°C, in good agreement with fluid inclusion temperatures. Even if only two mineral pairs are available and analytical errors must be considered, deposition at much lower temperature, say close to 400°C, can be ruled out

because the two minerals fractionate in opposite directions for geologically reasonable temperatures. A fluid source in equilib-rium with pyroxene and magnetite at 800°C has a δ18O value between 8.5 and 9.5 per mil, i.e., at the high end of the magmatic range. It is unlikely that pyroxene formed during an alteration stage preceeding the mineralization would have oxygen isotope ratios similar to pyroxene occurring in veins cutting major ore-bodies and thus postdating the main mineralization.

Another objection against a hydrothermal scenario is the diffi-culty to raise the temperature of a nonmagmatic fluid to tempera-tures exceeding 800°C. The nature of such a process remains to be explained. If the heating is attributed to fluid interaction with a hot magmatic bedrock, one would expect the fluid to signifi-cantly change its oxygen isotope composition and approach that of the magmatic bedrock. Again, such a fluid cannot explain the variable quartz data which definitely do not resemble any kind of magmatic source, and also generate pyroxene and magnetite data within the magmatic range.

Concluding remarksThe structural-textural evidence and oxygen isotope data indicate that the El Laco iron deposits formed in a cooling magmatic sys-tem. Our interpretation is that major orebodies resembling lava flows and associated pyroclastic material crystallized directly from iron-oxide magma at temperatures close to 1000°C (Table 5). Dikes with comb layering formed from magma and fluids in channelways. Fluids from cooling magma deposited magnetite, pyroxene, and apatite in fissures and open spaces, now present as veins and ore breccia. The temperature for this pneumatolytic process was about 700–800°C or locally even lower. However, the orebodies present little evidence for significant magnetite precipitation during hydrothermal conditions.

The approximately two per mil higher δ18O values of un-oxidized magmatic-textured magnetite from El Laco compared with corresponding values for the Chilean iron belt and Kiruna district (Fig. 2) could reflect primary differences in the evolution of the iron-rich magmas with respect to temperature regimes and sources, or be due to secondary processes acting on previ-ously formed magnetite. Considering a primary magma first, the crystallization temperature of magnetite is a factor, but a two per mil difference requires a temperature contrast of ca. 400°C that seems unrealistically high. Nevertheless, it makes sense that magnetites apparently formed from ore magmas have higher δ18O values than magnetites precipitated from somewhat cooler fluids at El Laco as well as Kiruna. The average difference is 1.3‰ for El Laco and 0.6‰ for Kiruna.

The El Laco deposits have only been affected by hydrother-mal alteration. The negligible oxygen diffusion rate during this low-temperature process suggests that the unoxidized magnet-ites essentially have kept their original isotopic compositions. It is possible that the 18O-rich El Laco magnetites formed in equilibrium with an iron-rich magma enriched in 18O. There is a regional pattern of elevated δ18O values for young volcanic rocks in northern Chile (Harmon et al. 1984), consistent with crustal contamination of magmas in the thick crust here. Hypo-thetically, also the isotopic effect of fractional crystallization of 16O-enriched magnetite at depth could be of relevance. If such process is extensive, any remnant magma reaching near-surface levels would be enriched in 18O. Geophysical data support the presence of magnetite deposits at depth beneath El Laco (Alva-Valdivia et al. 2003).


The relatively low δ18O values for unoxidized magnetites from the Chilean iron belt and Kiruna (Fig. 2), at the low end of the magmatic range, may suggest a weak secondary effect on the isotope systematics. Metamorphic fluids relatively enriched in 18O may have partly equilibrated with magnetite and shifted its oxygen isotope signature. The direction and size of the shift are mainly a function of the isotopic signature of surrounding bed-rock, the temperature, the oxygen diffusion rate, and the water availability. Such re-equilibration is unlikely to be substantial due to low diffusion rate in magnetite below 400°C (Freer & Dennis 1982), the predominantly magmatic host rocks in the two provinces, and the unoxidized nature of the magmatic-textured magnetites. Thus, our interpretation is that although magmatic-textured magnetite in the districts studied here may have suffered minor secondary disturbances, their isotopic signatures essen-tially reflect crystallization from magmas with initially slightly different isotopic characteristics.

Acknowledgements. – This study was supported by the Swedish Natural Science Research Council (NFR; grant GU 04535-301), the Fondo Nacional de Desarrollo Científico y Tecnológico (FONDECYT; grants 1930061 and 1970671) and the Departamento de Inves-tigaciones Científicas y Tecnológicas (DICYT, Universidad de Santiago de Chile; grants 05-92-15HB and 05-97-15HB).

ReferencesAlva-Valdivia, L.M., Rivas, M.L., Goguitchaichvili, A., Urrutia-Fucuguachi,

J., González, J.A., Morales, J., Gómez, S., Henríquez, F., Nyström, J.O. & Naslund, H.R., 2003: Rock-magnetic and oxide microscopic studies of the El Laco iron ore deposits, Chilean Andes, and implications for magnetic anomaly modeling. International Geology Review 45, 533–547.

Blake, K.L., 1992: The petrology, geochemistry and association to ore forma-tion of the host rocks of the Kiirunavaara magnetite-apatite deposit, northern Sweden. Ph.D. thesis, Department of geology, University of Wales, College of Cardiff. 307 pp.

Bookstrom, A.A., 1977: The magnetite deposits of El Romeral, Chile. Economic Geology 72, 1101–1130.

Borthwick, J. & Harmon, R.S., 1982: A note regarding ClF3 as an alternative to BrF5 for oxygen isotope analysis. Geochimica et Cosmochimica Acta 46, 1665–1668.

Broman, C., Nyström, J.O. & Henríquez, F., 1999: Fluid inclusions in magnet-ite-apatite ore from a cooling magmatic system at El Laco, Chile. GFF 121, 253–267.

Cisternas, M.E., 1986: Stratigraphische, fazielle und lithogeochemische Un-tersuchungen in der Unterkreide der Region Atacama: Metallogenetische Bedeutung am Beispiel der schichtgebundenen Eisen-Lagerstätte Bandurrias. Heidelberger Geowissenschaftliche Abhandlungen 2, 268 pp.

Clayton, R.N. & Mayeda, T.K., 1963: The use of bromine pentafluoride in the ex-traction of oxygen from oxides and silicates for isotopic analysis. Geochimica et Cosmochimica Acta 27, 43–52.

Cliff, R.A., Rickard, D. & Blake, K., 1990: Isotope systematics of the Kiruna magnetite ores, Sweden: Pt. 1. Age of the ore. Economic Geology 85, 1770–1776.

Dobbs, F.M. & Henríquez, F., 1988: Geología, petrografía y alteración del yacimiento de hierro Ojos de Agua, III Región. Actas V Congreso Geológico Chileno (Santiago) 3, G71–G81.

Edfelt, Å., Billström, K., Broman, C., Rye, R.O., Smith, M.P. & Martinsson, O., 2007: Origin and fluid evolution of the Tjårrojåkka apatite-iron and Cu(-Au) deposits, Kiruna area, northern Sweden. In Å. Edfelt: The Tjårrojåkka apatite-iron and Cu(-Au) deposits, northern Sweden. Ph.D. thesis, Department of Chemical Engineering and Geosciences, Luleå University of Technology, Sweden. Paper 3, 45 pp.

Espinoza, S., 1986: Sobre el origen volcánico-sedimentario de los mantos fer-ruginosos Bandurrias y Manolete, al sur de Copiapó. Revista Geológica de Chile 27, 33–40.

Espinoza, S., 1990: The Atacama-Coquimbo ferriferous belt, northern Chile. In L. Fontboté, G.C. Amstutz, M. Cardozo, E. Cedillo & J. Frutos (eds.): Strata-bound ore deposits in the Andes, 353–364. Springer.

Faure, G., 1986: Principles of isotope geology (2nd ed). John Wiley & Sons. 589 pp.

Förster, H. & Jafarzadeh, A., 1994: The Bafq mining district in central Iran–a highly mineralized Infracambrian volcanic field. Economic Geology 89, 1697–1721.

Fortier, S.M., Cole, D.R., Wesolowski, D.J., Riciputi, L.R., Paterson, B.A., Val-ley, J.W. & Horita, J., 1995: Determination of the magnetite-water equilibrium oxygen isotope fractionation factor at 350°C: a comparison of ion microprobe

and laser fluorination techniques. Geochimica et Cosmochimica Acta 59, 3871–3875.

Freer, R. & Dennis, P.F., 1982: Oxygen diffusion studies I. A preliminary ion microprobe investigation of oxygen diffusion in some rock-forming minerals. Mineralogical Magazine 45, 179–192.

Friedman, I. & OʼNeil, J.R., 1977: Compilation of stable isotope fractionation factors of geochemical interest. In M. Fleischer (ed.): Data of geochemistry, 6th edition, 1–12 U.S. Geological Survey Professional Paper 440KK.

Frietsch, R., 1978: On the magmatic origin of iron ores of the Kiruna type. Eco-nomic Geology 73, 478–485.

Frietsch, R., 1984: On the magmatic origin of iron ores of the Kiruna type – a reply: Economic Geology 79, 1949–1951.

Frutos, J. & Oyarzún, J., 1975: Tectonic and geochemical evidence concerning the genesis of El Laco magnetite lava flow deposits, Chile. Economic Geology 70, 988–990.

Gardeweg, M. & Ramírez, C.F., 1985: Hoja Río Zapaleri, II Región de Antofa-gasta. Carta Geológica de Chile 66. Servicio Nacional de Geología y Minería, Santiago. 89 pp.

Geijer, P., 1910: Igneous rocks and iron ores of Kiirunavaara, Luossavaara and Tuolluvaara. Scientific and practical researches in Lapland arranged by Luos-savaara-Kiirunavaara Aktiebolag, Stockholm. 278 pp.

Geijer, P., 1931: The iron ores of the Kiruna type. Sveriges Geologiska Under-sökning C367, 1–39.

Geijer, P., 1967: Internal features of the apatite-bearing magnetite ores. Sveriges Geologiska Undersökning C624, 1–32.

Grez, E., Aguilar, A., Henríquez, F. & Nyström, J.O., 1991: Magnetita Peder-nales: a new magmatic iron deposit in northern Chile. Economic Geology 86, 1346–1349.

Haggerty, S.E., 1970: The Laco magnetite lava flow, Chile. Carnegie Institution of Washington Year Book 68, 329–330.

Harlov, D.E., Andersson, U.B., Förster, H.-J., Nyström, J.O., Dulski, P. & Bro-man, C., 2002: Apatite-monazite relations in the Kiirunavaara magnetite-apa-tite ore, northern Sweden. Chemical Geology 191, 47–72.

Harmon, R.S., Barreiro, B.A., Moorbath, S., Hoefs, J., Francis, P.W., Thorpe, R.S., Déruelle, B., McHugh, J. & Viglino, J.A., 1984: Regional O-, Sr-, and Pb-isotope relationships in late Cenozoic calc-alkaline lavas of the Andean Cordillera. Journal of the Geological Society, London 141, 803–822.

Henríquez, F. & Martin, R.F., 1978: Crystal-growth textures in magnetite flows and feeder dykes, El Laco, Chile. Canadian Mineralogist 16, 581–589.

Henríquez, F. & Nyström, J.O., 1998: Magnetite bombs at El Laco volcano, Chile. GFF 120, 269–271.

Henríquez, F., Naslund, H.R., Nyström, J.O., Vivallo, W., Aguirre, R., Dobbs, F.M. & Lledó, H., 2003: New field evidence bearing on the origin of the El Laco magnetite deposit, northern Chile – a discussion. Economic Geology 98, 1497–1500.

Hildebrand, R.S., 1986: Kiruna-type deposits. Their origin and relationship to intermediate subvolcanic plutons in the Great Bear magmatic zone, Northwest Canada. Economic Geology 81, 640–659.

Hitzman, M.W., Oreskes N. & Einaudi M.T., 1992: Geological characteristics and tectonic setting of Proterozoic iron oxide (Cu-U-Au-REE) deposits. Pre-cambrian Research 58, 241–287.

Levi, B., Vergara, M., Nyström, J.O. & Henríquez, F., 1988: Low-grade meta-morphic facies series in the Mesozoic-Cenozoic volcanic sequences of central Chile: a re-appraisal and some applications. Anais do VII Congresso Latino-Americano de Geología, Belém, Brazil 1, 243–251.

Lundberg, B. & Smellie, J.A.T., 1979: Painirova and Mertainen iron ores: two deposits of the Kiruna iron ore type in northern Sweden. Economic Geology 74, 1131–1152.

Lyons, J.I., 1988: Volcanogenic iron oxide deposits, Cerro de Mercado and vicin-ity, Durango, Mexico. Economic Geology 83, 1886–1906.

Maksaev, V., Gardeweg, M., Ramirez, C.F. & Zentilli, M., 1988: Aplicación del método trazas de fisión (fission track) a la datación de cuerpos de magnetita de El Laco e Incahuasi en el Altiplano de la Región de Antofagasta. Actas V Congreso Geológico Chileno (Santiago) 1, B1–B23.

Matthews, A., Goldsmith, J.R. & Clayton, R.N., 1983: Oxygen isotope fractiona-tion involving pyroxenes: the calibration of mineral-pair geothermometers. Geochimica et Cosmochimica Acta 47, 631–644.

Ménard, J.-J., 1995: Relationship between altered pyroxene diorite and the magnetite mineralization in the Chilean Iron Belt, with emphasis on the El Algarrobo iron deposits (Atacama region, Chile). Mineralium Deposita 30, 268–274.

Naslund, H.R., Dobbs, F.M., Henríquez, F. & Nyström, J.O., 1997: Irrefutable evidence for the eruption of iron-oxide magmas at El Laco Volcano, Chile. EOS, Transaction of the American Geophysical Union 78, no. 17, S333.

Naslund, H.R., Aguirre, R., Dobbs, F.M., Henríquez, F. & Nyström, J.O., 2000: The origin, emplacement, and eruption of ore magmas. Actas IX Congreso Geologico Chileno (Puerto Varas) 2, 135–139.

Naslund, H.R., Henríquez, F., Nyström, J.O., Vivallo, W. & Dobbs, F.M., 2002: Magmatic iron ores and associated mineralisation: example from the Chilean High Andes and Coastal Cordillera. In T.M. Porter (ed.): Hydrothermal iron oxide copper-gold & related deposits: a global perspective, vol. 2, 207–226. PGC Publishing, Adelaide.

Nyström, J.O. & Henríquez, F., 1994: Magmatic features of iron ores of the Kiruna type in Chile and Sweden: ore textures and magnetite geochemistry. Economic Geology 89, 820–839.


Nyström, J.O. & Henríquez, F., 1995: Magmatic features of iron ores of the Kiruna type in Chile and Sweden: ore textures and magnetite geochemistry – a reply. Economic Geology 90, 473–475.

Nyström, J.O., Henríquez, F. & Travisany, V., 1996: Volcanic iron ores of the Kiruna type. GFF 118, A45–A46.

OʼFarrelly, K. & Rickard, D., 1989: Stable isotope evidence for the origin of the Kiruna magnetite-apatite ore, Sweden. Terra abstracts, 344–345.

Oyarzún, J. & Frutos, J., 1984: Tectonic and petrological frame of the Cretaceous iron deposits of North Chile. Mining Geology 34, 21–31.

Oyarzún, R., Oyarzún, J., Ménard, J.J. & Lillo, J., 2003: The Cretaceous iron belt of northern Chile: role of oceanic plates, a superplume event, and a major shear zone. Mineralium Deposita 38, 640–646.

Parák, T., 1975: Kiruna iron ores are not ”intrusive-magmatic ores of the Kiruna type”. Economic Geology 70, 1242–1258.

Parák, T., 1984: On the magmatic origin of iron ores of the Kiruna type – a discussion. Economic Geology 79, 1945–1949.

Park, C.F. Jr., 1961: A magnetite ”flow” in northern Chile. Economic Geology 56, 431–436.

Park, C.F. Jr., 1972: The iron ore deposits of the Pacific Basin. Economic Geology 67, 339–349.

Rhodes, A.L. & Oreskes, N., 1994: The magnetite ”lava flows (?)”, El Laco, Chile: new evidence for formation by vapor transport. Actas VII Congreso Geológico Chileno (Concepción) 2, 1501–1505.

Rhodes, A.L. & Oreskes, N., 1999: Oxygen isotope composition of magnetite deposits at El Laco, Chile: Evidence of formation from isotopically heavy fluids. Society of Economic Geologists Special Publication 7, 333–351.

Rhodes, A.L., Oreskes, N. & Sheets, S., 1999: Geology and rare earth element geochemistry of magnetite deposits at El Laco, Chile. Society of Economic Geologists Special Publication 7, 299–332.

Rogers, D.P., 1969: The extrusive iron oxide deposits, “El Laco”, Chile. Geologi-cal Society of America, Special Paper 121, 252–253.

Romer, R.L., Martinsson, O. & Perdahl, J.-A., 1994: Geochronology of the Kiruna iron ores and hydrothermal alterations. Economic Geology 89, 1249–1261.

Ruiz, C. & Peebles, F., 1988: Geología, distribución y génesis de los yacimientos metalíferos chilenos. Editorial Universitaria, Santiago. 334 pp.

Sheets, S.A., 1997: Fluid inclusion study of the El Laco magnetite deposits, Chile. M.Sc. thesis, Dartmouth College, Hanover, New Hampshire. 94 pp.

Sillitoe, R.H. & Burrows, D.R., 2002: New field evidence bearing on the origin of the El Laco magnetite deposit, northern Chile. Economic Geology, 97, 1101–1109.

Taylor, H.P. Jr., 1967: Oxygen isotope studies of hydrothermal mineral deposits. In H.L. Barnes (ed.): Geochemistry of hydrothermal ore deposits, 109–142. Holt, Rinehart and Winston.

Taylor, H.P. Jr., 1968: The oxygen isotope geochemistry of igneous rocks. Con-tributions to Mineralogy and Petrology 19, 1–71.

Travisany, V., Henríquez, F. & Nyström, J.O., 1995: Magnetite lava flows in the Pleito-Melón district of the Chilean iron belt. Economic Geology 90, 438–444.

Vivallo, W., Henríquez, F. & Espinoza, S., 1994: Oxygen and sulfur isotopes in hydrothermally altered rocks and gypsum deposits at El Laco mining district, northern Chile. Comunicaciones, Departamento de Geología, Universidad de Chile, Santiago 45, 93–100.

Wegner, R., 1982: El Laco/Chile. Lavaströme aus Eisenerz. Lapidus 7, 15–18, 42.

Zheng, Y.-F., 1995: Oxygen isotope fractionation in magnetites: structural effect and oxygen inheritance. Chemical Geology 121, 309–316.


Fluid inclusions in magnetite-apatite ore from a cooling magmatic system at El Laco, Chile

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