Mineralogical Association of Canada Short Course, 36 MELT INCLUSIONS IN PLUTONIC ROCKS

The cover photo shows a water-rich melt inclusion in pegmatite quartz from Ehrenfriedersdorf/Germany, re-homogenized at...

Author: Mineralogical Association of Canada

51 downloads 755 Views 22MB Size Report

This content was uploaded by our users and we assume good faith they have the permission to share this book. If you own the copyright to this book and it is wrongfully on our website, we offer a simple DMCA procedure to remove your content from our site. Start by pressing the button below!
Report copyright / DMCA form

DOWNLOAD PDF

The cover photo shows a water-rich melt inclusion in pegmatite quartz from Ehrenfriedersdorf/Germany, re-homogenized at 600°C and 1 kbar using the cold-seal pressure vessel technique. The vessel was pressurized with CO2. The melt inclusion glass contains 25wt% H2O determined with Raman spectroscopy. The results are given in the paper by Thomas, Kamenetsky and Davidson - example b of Table 2. The yellow color corresponds to 25 wt.% water and the dark green color at the basis corresponds to about 0 wt.% (quartz). The blue ridge at the right side comes from a small secondary fluid inclusion trail perpendicular to the quartz chip. The water distribution image was taken with the confocal high resolution Raman device LabRam HR800 using the 488 nm excitation Laser line. The image is composed of 750 single points, each 2µm. Behind each point is a complete Raman spectrum taken in the frequency range from 2800 to 4000 cm–1. For the quantification of the water we used the integral intensity of the broad asymmetric H2O-OH Raman band between 3100 and 3750 cm–1.

Distribution

Copies of Volume 36 and preceding volumes are obtainable from: The Business Manager Mineralogical Association of Canada P.O. Box 78087 Meriline Postal Outlet 1460 Merivale Road Ottawa, Ontario Canada K2E 1B1 e-mail: [email protected] www.mineralogicalassociation.ca

Series Coordinator: R. Raeside Department of Geology Acadia University Wolfville, Nova Scotia Canada, B4P 2R6 Fax: +1 902-585-1816 e-mail: [email protected] ISBN 0-921294-36-0 Copyright 2006 – Mineralogical Association of Canada Printed in Canada

CHAPTER 1: MELT INCLUSIONS IN PLUTONIC ROCKS: PETROGRAPHY AND MICROTHERMOMETRY R. J. Bodnar Fluids Research Laboratory, Virginia Tech, Blacksburg VA 24061, USA E-mail: [email protected] and James J. Student Dept. of Geology, Central Michigan University, Mt. Pleasant, MI 48859, USA E-mail: [email protected] Erzgebirge, Germany, were smaller than 20 µm, with a mean diameter of about 10 µm. R. Thomas (personal communication, 2006) also noted a correlation between inclusion size and volatile content, with more volatile-rich inclusions reaching 200 µm in diameter. Thomas also suggested that MI in Precambrian rocks are generally smaller than those in younger rocks. MI in plutonic rocks are commonly completely crystalline and contain a distorted bubble that is not always recognizable, and are much more likely to have reequilibrated following entrapment, owing to slow cooling and contact with subsolidus hydrothermal fluids. This chapter summarizes progress in the development of techniques to study MI from plutonic/intrusive rocks and focuses on MI from more silicic (granitic) environments (Table 1-1). Veksler (2006) discusses crystallized MI from gabbroic rocks, and Veksler & Lentz (2006) consider MI in carbonatites and related rocks. The application of MI in understanding pegmatite formation is discussed by Thomas et al. (2006). In a more general sense, this chapter focuses on MI that have experienced significant devitrification or crystallization following entrapment, or subsolidus aqueous alteration. This discussion thus applies to MI from the plutonic environment, as well as inclusions from the volcanic or extrusive environment that have undergone slow cooling and/or hydrothermal alteration.

INTRODUCTION Melt inclusions (MI) are small droplets of melt now containing some combination of crystals, glass and vapor that are trapped in crystals formed in magmas. Over the past few decades, the study of MI has matured to become an accepted technique to investigate melt evolution in volcanic systems (Clocchiatti 1975, Roedder 1979, Lowenstern 1995, 2003, Sobolev 1996, Frezzotti 2001, Hauri et al. 2001, Anderson 2003, Schiano 2003). MI in volcanic rocks are commonly large (>50 µm), glassy and contained in fresh and transparent minerals, and are generally easy to identify. MI in volcanic rocks generally provide consistent and reasonable results concerning the chemistry of melts at depth and provide the best tool available for assessing the volatiles in magmas (Anderson 1973, 1974, Lowenstern 1994, 1995, 2003). The best MI are often found in the most rapidly cooled parts of pyroclastic deposits and glassy rinds of lavas, which are not so easy to find (A.T. Anderson, Jr., personal communication, 2006). Compared to studies of MI in extrusive rocks, there are many fewer studies of MI from plutonic rocks, and there is still some uncertainty concerning the interpretation of data obtained from these samples. Roedder (1984) noted, “The lack of evident silicate melt inclusions in many igneous intrusive rocks, particularly those formed at greater depths, is puzzling.” Roedder further stated, “One of the major problems in the study of melt inclusions in intrusive rocks is the difficulty in recognizing the inclusions.” MI in intrusive rocks are usually small (5–20 µm) compared to those in extrusive rocks. Thomas et al. (1996) noted that about 80% of the MI in granitic rocks from the

IN THE BEGINNING …… The first detailed discussion of MI within a rigorous geological and petrological context was by Sorby (1858). In his classic 1858 paper on fluid and melt inclusions, Sorby introduced the section on

Mineralogical Association of Canada Short Course 36, Montreal, Quebec, p. 1-25.

1

2003 1984 2004 1997 2003 2001 1979 2001 2005 2000 1992 2000 1990 1986 2003 2004 1999 1995 1997 2004 1999 2004 2000 2002 2002 2004 1975 1988 1994a 1994b 1994c 1997 2000 2002 2003 2005 1998 2001 2004 1994

2

qtz Many qtz qtz, fspar, ap, tpz and zn qtz qtz qtz qtz qtz qtz, tpz qtz qtz, fspar qtz qtz cpy qtz qtz, fspar qtz qtz qtz qtz, tpz qtz qtz qtz, pyr qtz qtz qtz, ap qtz, tpz, ap, fspar, gar, fl qtz, K-fspar qtz, fspar, ap, tpz qtz zn qtz qtz qtz qtz qtz qtz

Xl, FI Gl, Xl Xl Xl Gl Xl Xl Xl, Gl, FI Gl, Xl Dev, crypto-Xl Gl + Xl + V Xl partly Xl, FI Gl, FI Xl Xl Mix, FI Xl Mix Xl Xl Xl Xl Xl Xl Xl, Mix, FI Gl, Xl Mix Xl, V Xl, FI Xl Xl Xl Xl Xl Xl Xl Gl Xl Xl, Mix

No Yes Yes Yes No Yes Yes Yes Some Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes Yes No Yes Yes

Hmzd? (4) NA NR RQ CSVP (100) Furn NR NR NR HS HS (?) Furn HS Furn HS HS HS Furn, HS HS CSVP(?) (300) Furn CSVP (?) HS IHPV (500) HS, IHPV (500) CSPV(?) Furn, HS, CSPV (50-150) Furn, HS, CSPV (50-150) Furn, HS Furn, HS Furn Furn Furn Furn RQ CSVP (100) CSVP RQ CSVP (100) RQ CSVP (100), Furn HS NR Furn Furn

Heating Method (5) LA-ICPMS EMPA EMPA, Raman EMPA EMPA, LA-ICPMS EMPA None EMPA, SEM, Raman, LA-ICPMS EMPA, LA-ICPMS, SEM, PIXE, Raman EMPA, PIXE, SIMS EMPA EMPA EMPA EMPA PIXE SEM, PIXE SEM, EMPA, LA-ICPMS, Raman EMPA, SIMS EMPA EMPA, SEM, Raman, SIMS None EMPA EMPA, SEM EMPA, SIMS EMPA EMPA EMPA, Laser Probe-MA None None None None EPMA, SIMS EMPA, Raman EMPA, SIMS EMPA, SIMS, Raman EMPA, SIMS, Raman, SXRF SEM EMPA, SIMS EMPA, SIMS EPMA

Analyses Conducted (6)

(1). PCD = porphyry copper deposit; Peg = pegmatite; Rhy = rhyolite; Gran = granite; Monz = monzonite; QM = quartz monzonite; Dac = dacite; Alb = albite; zinn = zinnwalldite; (2). qtz=quartz; fspar = feldspar; ap = apatite; tpz = topaz; cpy = clinopyroxene; zn = zircon; pyr = pyroxene; and = andalusite; gar = garnet; fl = fluorite; (3). Xl = crystalline; Gl = glass; Dev = devitrified glass; Mix = glass + crystals; V = vapor; FI = fluid inclusions; (4). Were the inclusions heated to homogenization?; (5). Furn = 1-atmosphere furnace; HS = microscope heating stage; CSPV = cold seal pressure vessel (pressure in MPa); IHPV = internally heated pressure vessel (pressure in MPa);RQ = Rapid-quench; PCA = Piston cylinder apparatus; NR = not reported; NA = not applicable; (6). EMPA = electron microprobe; SIMS = secondary ion mass spectrometry; SEM = scanning electron microscope; Raman = Raman spectroscopy; LA-ICPMS = laser ablation inductively coupled plasma mass spectrometry; PIXE = particle induced X-ray emmission; SXRF = synchrotron X-ray fluorescence; MA = mass analyzer

Barren Gran Various LiF-rich Gran P-rich rare-metal Peg & Gran Anatectic metapelite Charnockitic orthogneiss Rapakivi Gran PCD Cu-Mo PCD PCD rhy-dac intrusive bodies Gran Topaz Gran peralkaline Gran Plutonic ejecta Quartz veins Dac porphyry Gran Syenitic dike in Cu-Au PCD Peralkaline Gran Granophyric blocks in rhy ejecta Tourmaline rich gem Peg Granodiorite Peralkaline Gran LiF-enriched Gran Peralkaline Gran PCD QM PCD QM Qtz Porphyry, Gran; Rhy Gran, porphyry, ongonite, kersanite Gran Variscan Gran Precambrian Gran Gran Tin Peg Tonalite Peraluminous Gran-Peg Tin Gran Gran, qtz monzodiorite Gran Xenoliths Alb-zinn porphyritic micro-Gran Deep crustal granitoid intrusions

MI Types (3)

Audétat & Pettke Bakumenko et al. Badanina et al. Breiter et al. Cesare et al. Chupin et al. Chupin et al. Davidson & Kamenetsky Davidson et al. Dietrich et al. Frezzotti Haapala & Thomas Hansteen & Lustenhouwer Harris Harris et al. Kamenetsky & Naumov Kamenetsky et al. Kovalenko et al. Lowenstern et al. Peretyazhko et al. Prokofiev et al. Reyf Reyf et al. Schmitt et al. Student Student & Bodnar Takenouchi & Imai Thomas Thomas Thomas Thomas Thomas & Klemm Thomas et al. Thomas et al. Thomas et al. Thomas et al. Vapnik Webster & Rebbert Webster et al. Yang & Bodnar

Host Minerals (2)

Reference

Rock Type/Environment (1)

OF STUDIES OF MELT INCLUSIONS IN SILICIC PLUTONIC ROCKS

Date

TABLE 1-1: SUMMARY

MAGMATIC PROCESSES AND VOLATILE PHASE GENERATION IN PORPHYRY-TYPE ENVIRONMENTS

MELT INCLUSIONS IN PLUTONIC ROCKS: PETROGRAPHY AND MICROTHERMOMETRY

melt inclusions with the statement “The formation of crystals from a state of igneous fusion is in every respect analogous to what takes place when crystals are formed in water”. Later he stated “There is thus a most perfect analogy between glass- and stonecavities and fluid-cavities in every respect except the nature of the included substances”. The implication of this statement is that assumptions applied to the interpretation of FI should also apply to MI. Sorby referred to MI containing glass (± a vapor bubble) as glass-cavities; those in which the melt has crystallized are referred to as stonecavities. Sorby equated the stone-cavities to

aqueous FI that are cooled to low temperature and contain ice plus various salts and hydrates that were originally in solution. Sorby examined both glassy and crystallized MI and arrived at interpretations that are still valid today. Cameras were unavailable to Sorby to photograph features observed with the microscope, and he drew illustrative sketches of inclusions that he observed in thin sections and grain mounts. Several of Sorby’s drawings of crystallized MI from plutonic environments are shown in Figure 1-1. Much of Sorby’s understanding of MI is based on studies of

FIG. 1-1. Drawings of crystallized melt inclusions from Sorby (1858). The number above each drawing refers to the figure number in the original publication. Sorby did not include a scale as in modern publications but, rather, indicated the number of times the inclusions are magnified in linear dimensions, indicated by the number in parentheses following each description. (A) Crystallized melt inclusion in a crystal of iron silicate from a copper slag (X1600); (B) Crystallized melt inclusion in pyroxene from a blast furnace slag (X400), (C) Crystallized melt inclusion in feldspar in a xenolith from Vesuvius, Italy (X500); (D) Crystallized melt inclusion in quartz from Ponza, Italy (X400); (E) Crystallized melt inclusion in trachyte from Ponza, Italy (X800); (F) Altered melt inclusion from a porphyry from Arthur’s Seat, near Edinburgh, UK (X400); (G) Crystallized melt inclusion in quartz from an elvan near Penrhyn, Cornwall, UK (X250); (H) Crystallized melt inclusion in quartz from the granite at St. Austell, UK (X1000); (I) Altered melt inclusion from a porphyry from Arthur’s Seat, near Edinburgh, UK (X200); (J) Crystallized melt inclusion in quartz from an elvan near Penrhyn, Cornwall, UK (X800); (K) Crystallized melt inclusion in quartz from a coarse-grained granite near Cape Cornwall, UK (X800); (L) Crystallized melt inclusion in quartz from the granite at St. Austell, UK, with radiating fractures (X600).

3


inclusions in slags from iron furnaces in England (Fig. 1-1A, B). Sorby also described many natural crystallized MI, including “stone-cavities” from the island of Ponza (Italy) (Fig. 1-1D, E). Fedele et al. (2003) studied similar samples from Ponza and heated the MI to produce a homogeneous glass that was then analyzed by electron microprobe. The results were used to develop a model for igneous petrogenesis in the Ponza trachyte. Sorby was among the first to recognize that coexisting FI and MI indicated entrapment in a volatile-saturated magma. Today melt-volatile immiscibility is recognized as an important process in many magmatic–hydrothermal systems (Roedder 1992), ranging from shallow granites (Frezzotti 1992) to porphyry copper deposits (Bodnar 1995, Kamenetsky et al. 1999, Davidson & Kamenetsky 2001) to orthomagmatic Au–Te deposits such as Cripple Creek, CO (Webster 2004). Referring to inclusions trapped in a volatile-saturated melt, Sorby concluded “Some crystals might be deposited from solution in the highly heated water, and catch up small portions of the fused stone [to form melt inclusions], whilst others might be formed by the crystallization of the melted stone, and catch up small portions of liquid water [to form aqueous inclusions]”. He went on to state “… we may, I think, conclude that the crystals would contain glass- [glass melt inclusions] or stone-cavities [crystallized melt inclusions], and perhaps gas- and vapour-cavities [fluid inclusions], ....” Today the occurrence of coexisting MI and FI that were trapped from the same magma is considered to be unequivocal evidence for melt–volatile immiscibility in the magma, as was documented at Ascension Island by Roedder & Coombs (1967) based on coexisting MI and halite-rich FI. In fact, one could conclude that an assumption of volatile saturation in the absence of coexisting MI and FI is highly questionable.

recognizing completely crystallized MI in plutonic rocks (Roedder 1984, Yang & Bodnar 1994), and MI that have been altered by later hydrothermal processes (Frezzotti 1992, Varela 1994, Student & Bodnar 2004). Altered, crystallized MI are especially common in the porphyry copper deposits (Student & Bodnar 2004), where the rock has undergone subsolidus alteration associated with hydrothermal activity during formation of the ore deposit. In this environment, planes of secondary aqueous FI commonly intersect the MI, and the compositions of the MI reflect alteration by the hydrothermal fluid (Frezzotti 1992, Student & Bodnar 2004). Crystallized MI usually appear as opaque, poorly defined patches at low magnification (Fig. 1-2). At higher magnification, the inclusions appear to be dark masses of crystals lacking sharp borders between the MI and the host (Fig. 1-3). Under crossed polars the inclusions often show a few bright spots representing birefringent daughter minerals, or an overall sparkly appearance in finergrained inclusions. In many cases it may be difficult to distinguish between crystallized MI and trapped solids in the host (Rapien et al. 2003). The largest MI in plutonic rocks are generally less than several 10s of micrometres in longest dimension, whereas MI up to several hundred micrometres are not uncommon in extrusive rocks.

PETROGRAPHIC ANALYSIS OF MELT INCLUSIONS IN PLUTONIC ROCKS Sorby noted that MI in plutonic rocks are often difficult to identify because “In the quartz of very coarse grained granites the stone cavities [crystallized melt inclusions] are generally obscure and of irregular shape….” and “Those in the feldspar are often so much obscured by the partial decomposition of that mineral, that it is difficult to distinguish them from small decomposed patches [of feldspar]…” (see Fig. 1-1F-L). More recent workers have also discussed the difficulty in

FIG. 1-2. Quartz phenocryst from Stage IV granodiorite from the Tyrone, New Mexico, porphyry copper deposit (Student & Bodnar 2004). The crystal is approximately 2 mm across. The phenocryst contains abundant crystallized MI, including hourglass inclusions (Anderson 1991), large, randomly distributed MI (larger opaque areas in crystal) and numerous, small MI along a growth zone near the outer edge of the crystal.

4


FIG. 1-3. Examples of melt inclusions in quartz phenocrysts from Stage IV granodiorite from the Tyrone, New Mexico, porphyry copper deposit (Student & Bodnar 2004). All of the inclusions occur in the same phenocryst, which is approximately 2 mm across (see Figure 1-2). The scale bar, shown in (Q) represents 30 µm and applies to all images. MI that occur along growth zones (G) are generally smaller than those that are randomly distributed throughout the crystal. Some MI (K, P, Q) appear to have nucleated on solid inclusions in the phenocryst; other melt inclusions (M–O) have anomalously large vapor bubbles, suggesting either entrapment of vapor along with melt or later reequilibration. Some MI (D, E) have fractures that extend into the surrounding quartz.

Many workers have discussed processes that lead to the formation of crystallized MI (Roedder 1979, 1984, Lowenstern 1995, Frezzotti 2001). In general, for a given inclusion size, the slower the cooling rate, the more likely that the melt in the inclusions will crystallize. Similarly, for a given cooling rate, larger inclusions are more likely to crystallize compared to smaller inclusions. In addition to cooling rate and size, the composition of the melt may affect the crystallization behavior of MI (Roedder 1984). Student & Bodnar (1999) noted that synthetic MI trapped under H2O-saturated conditions in the quartz-saturated haplogranite system were often partially to completely crystallized (Fig. 1-4), even though the samples were quenched from formation conditions (720°C and 200 MPa) to room temperature in a few

minutes. These workers suggested that the high H2O contents of the melts, and the fact that the melts were saturated in H2O at the time of trapping, might have promoted crystallization during cooling. Inclusion Selection The most important aspect of any MI or FI study is determining which inclusions to study and whether those inclusions record the physical and chemical conditions at the time of trapping. As such, the timing of inclusion trapping relative to formation of the host phase, and the position of the host within the overall paragenesis, must be constrained. There is surprisingly little discussion in the literature concerning the temporal classification of MI. Sobolev & Kostyuk (1975) described the use of MI in studies of magmatic crystallization,

5

R. J. BODNAR & J.J. STUDENT

and summarized the different temporal occurrences of MI. These workers distinguished between zonal and azonal primary MI. Zonal inclusions define a growth zone and there is little debate concerning their primary origin. [Note, however, that a primary origin does not guarantee that the inclusions record the physical and chemical conditions of crystal growth, as the inclusions may have reequilibrated following entrapment, as described in more detail below.] Zonal inclusions are common in some minerals, such as nepheline, plagioclase (Halter et al. 2004b) and pyroxene (Yang & Scott 2002, Rapien et al. 2003) (Fig. 1-5), but are less common in other minerals such as olivine (Anderson 1974) and zircon (Thomas et al. 2002). Isolated or randomly distributed inclusions that cannot be associated with a specific growth feature are referred to as azonal inclusions (Fig. 1-6). Such isolated inclusions are generally considered to be primary if no evidence of fracturing or mineral dissolution (that could allow secondary inclusions to form) is observed. Sobolev & Kostyuk (1975) emphasized that a negative crystal shape is not in itself sufficient evidence for a primary origin of MI, as has also been noted for FI (Roedder 1984, Bodnar 2003a). Some minerals, such as quartz (Anderson et al. 2000, Halter et al. 2004a) commonly contain both zonal and azonal MI (Fig. 1-2). Several workers (Yang & Scott 2002, Student & Bodnar 2004) have noted that within a given phenocryst zonal inclusions tend to be considerably smaller than azonal MI. Within the FI community, much effort has been devoted to the issue of inclusion selection. A procedure has been developed that allows one to be confident that the FI selected are related to the process being studied, and that the inclusions have not been affected by later events. The first step is to identify a group of FI that were all trapped at the same “time” and, by inference, at the same temperature and pressure and from fluids of the same composition. This group of inclusions, representing the most finely discriminated trapping event that can be identified based on petrography, is referred to as a fluid inclusion assemblage or FIA (Goldstein & Reynolds 1994). The amount of time represented by an FIA will vary, depending on the geologic environment. Thus, the amount of time represented by FI trapped along a growth surface of a halite crystal forming as a result of diurnal temperature variations in a sabkha environment is less than 24 hours. Similarly, the amount of time required to heal a fracture in a high temperature

FIG. 1-4. Natural (top) and synthetic (bottom) crystallized MI. The natural MI is from a quartz phenocryst in Stage IV granodiorite from the Tyrone, New Mexico, porphyry copper deposit (Student & Bodnar 2004). The synthetic MI is in quartz and was trapped at 720°C and 200 MPa under H2O-saturated conditions (Student & Bodnar 1999).

FIG. 1-5. Photomicrograph of primary MI outlining growth zones in a pyroxene phenocryst from the White Island, New Zealand, volcano (from Rapien et al. 2003).

6


the host crystal. However, MI may be preferentially trapped during episodes of anomalous and rapid crystal growth associated with changing temperature and/or pressure and/or volatile content in the magma (Roedder 1984), effectively decreasing the amount of time represented by the inclusion assemblage. It is important to emphasize that the fluid inclusion assemblage definition implies nothing about the temporal relationship between the inclusions and the host mineral. An FIA may be composed of either primary or secondary or pseudosecondary inclusions, The FIA concept does not constrain when the inclusions were trapped relative to growth of the host mineral (other than that they must have been trapped during or after the host formed). Thus, secondary (or pseudosecondary) FI along a well-defined fracture represent an FIA that was formed after precipitation of the bulk of the mineral in which those inclusions occur. While the FIA approach to selecting inclusions to study is used routinely in studies of FI, an analogous method has not been developed for MI. However, because similar petrographic techniques are (and should) be used to identify groups of MI that were trapped at the same time, we propose the term melt inclusion assemblage (MIA) to describe a group of MI trapped at (essentially) the same time and, by analogy, at the same temperature and pressure and from a melt of the same composition. Thus, MI along an individual growth zone in a phenocryst, such as those shown in Figures 1-2 and 1-5, represent MIAs containing primary MI. Yang & Bodnar (1994) studied MI in quartz from deep crustal granodiorite plutons in the Gyeongsang Basin, South Korea, and identified various MIAs based on petrographic analysis of the samples. One type of MIA occurred near the interface between plagioclase and quartz crystals that extended into vugs, while other MIAs were observed along growth surfaces near the outer portions of quartz crystals (Fig. 1-7). In both cases, petrographic evidence for simultaneous trapping of the MI in the MIA was conclusive (although there is some uncertainty as to whether the MIA at the plagioclase–quartz interface represents primary or secondary MI). Compared to FI, secondary MI are much less common. Thus, at a minimum the MI in a given phenocryst can be considered to be an MIA composed of inclusions that were all trapped during the (unknown) amount of time required for the phenocryst to grow.

FIG. 1-6. (Top) Quartz phenocryst from the Bishop Tuff (upper left) containing numerous randomly distributed (azonal) glassy MI. (Bottom) Transmitted light photographs of melt inclusions (MI) in zircons from the Quottoon Igneous Complex, British Columbia, Canada (from Thomas et al. 2002).

hydrothermal environment to produce a plane of secondary FI might be on the order of days to weeks (Brantley 1992, Sterner & Bodnar 1984, Bodnar & Sterner 1987). Dowty (1980) and Bacon (1989) compiled crystal growth rates from melts for various minerals. In general, growth rates fall in the range from about 10–2 to 10–11 cm/s. These growth rates would require 0.1 sec to 1011 sec (3.17 x 103 years) to trap a 10 µm MI. Tomiya & Takahashi (2005) estimated average growth rates for plagioclase and pyroxene in the magma beneath Usu volcano (Japan) of 0.1 to 0.7 µm/a. Thus, an FIA composed of 10 µm MI along a growth zone in a plagioclase or pyroxene phenocryst from this system would represent a minimum of about 15–100 years. This amount of time assumes a constant growth rate for 7


evaluating the quality of data obtained from MI is to identify MI that represent a MIA, as described below (Fig. 1-8). If MI do not occur in growth zones or along fractures, it is sometimes possible to group MI that show similar petrographic characteristics and, by inference, trapped a melt of the same composition. Anderson et al. (2000) noted that quartz phenocrysts from the Bishop Tuff contain clear, faceted MI in the interior, and round, brown MI close to the edge of the crystal. These workers used this zonal arrangement to infer the sequence of trapping – thus the clear MI would be assigned to one MIA, and the brown MI to a later MIA, both associated with growth of the crystal. Recently, many workers have successfully used cathodoluminescence (CL) analysis to identify FI that belong to the same FIA (Landtwing et al. 2005). Peppard et al. (2001) used CL to reveal growth zones in quartz from the Bishop Tuff, and used this information to identify MI with a common origin. CL is a powerful technique that can be used to reveal growth textures in phases that show no such features during normal petrography, and its use in MI studies is expected to increase. MICROTHERMOMETRY In FI studies, once a FIA has been identified, the next step is to determine if the inclusions record the physical and chemical conditions of trapping. To address this, the FIA must be tested to determine if the inclusions adhere to “Roedder’s Rules”, which are based on criteria first proposed by Sorby (1858). According to Roedder (1984), FI record the original trapping conditions if the following requirements are met: 1. the inclusions trap a single, homogeneous phase, 2. nothing has been added to or removed from the inclusion following trapping, 3. the inclusion represents an isochoric (constant volume) system. Note that the requirement of constant volume applies to the volume of the original cavity in which the fluid was trapped. During cooling of a MI to room temperature some (considerable?) amount of material may precipitate on the inclusion walls, resulting in a volume that is apparently smaller than the original MI volume at the trapping conditions. If this material is incorporated back into the melt when the MI is heated to homogenization (see Fig. 1-9), the condition of constant volume is satisfied. The condition would not be satisfied if the host mineral surrounding the MI deformed

FIG. 1-7. Schematic representation of the distribution of mostly crystallized MI in quartz from deep crustal granitoid intrusions from the Gyeongsang Basin, South Korea. Abundant small (12 µm) MI show random or isolated distribution (B). The outer growth zone of the quartz is decorated by numerous MI (C). Some MI in the outermost portion of the crystal contain devitrified glass (D). The scale corresponds to the crystal only – MI have been drawn larger for illustrative purposes. (from Yang & Bodnar 1994).

It is relatively easy to assign MI along growth surfaces or other interfaces to the same MIA. However, many minerals or phenocrysts contain only one or a few melt inclusions, and these are commonly distributed irregularly within the grain (Fig. 1-6). Where two or more randomly distributed MI occur within a given crystal, petrographic observations usually do not provide sufficient information to determine if the inclusions represent an MIA, i.e., were trapped at the same time and at the same temperature and pressure and from a melt of the same composition. This distinction is important because the first step in 8


Identify a Melt Inclusion Assemblage (MIA) based on petrography

Stop

No

Consistent Phase Relations?

No Why do MI have variable phase relations and/or microthermometry?

Yes

Homogenize MI: Consistent Microthermometry?

No Mixed Trapping Reequilibration in nature

Yes

Reequilibration in the laboratory

Analyze MI

MI do not represent an MIA

Interpret Results

FIG. 1-8. Flow chart showing the steps that should be followed to ensure that data obtained from MI represent conditions in the magma at the time of trapping. The first step is to identify MI that were trapped at the same time, and at the same temperature and pressure and from a melt of the same composition, and thus represent a melt inclusion assemblage (MIA). Then, petrographic, microthermometric and chemical data are evaluated to test for reequilibration of the MI after trapping.

FIG. 1-9. Series of photomicrographs showing the behavior of two crystallized MI in quartz from the Red Mountain, Arizona, USA, porphyry copper deposit during heating. Evidence of melting is first observed at about 675°C. At 790°C the inclusions contain a vapor bubble and a small feldspar crystal (not visible). At 810°C both the vapor bubble and feldspar disappear (from Student & Bodnar 2004).

9


1400°C range. These workers heated the sample at a rate of 30–40°C/minute for temperatures above 900°C and held the sample at the maximum temperature for 7 minutes before quenching. They noted that slower heating rates or longer time at the maximum temperature resulted in an increased Th and oxidation of the sample. Lowenstern (1994) heated MI in quartz phenocrysts from Pantelleria, Italy, in a one atmosphere heating stage at a rate of 50°C up to 600°C. The sample was held at 600°C for 10–20 minutes, then heated to 700°C and maintained at this temperature for 10 minutes. The sample was then heated in 50°C increments with 10 minute equilibration periods after each step up to 800°C. At higher temperatures 25°C heating steps were used. Lowenstern (1994) noted that these heating rates were likely too rapid, resulting in observed Th that were 25–75°C higher than the actual temperatures. Reyf (1997) heated small (3–5 µm) MI in quartz from a granite in a microscope heating stage, using heating steps of 30–50°C, and holding the sample at temperature for 1.5–2 hours after each step. Although this technique involved long heating experiments, the results were consistent with previous work that suggested host granites were hypersolvus and crystallized outside of the region of immiscibility in the Ab–An–Or system at 50–100 MPa. Massare et al. (2002) studied MI in olivine phenocrysts and found that when heated at one atmosphere the Th increased systematically with time as a result of deformation of the host crystal and loss of H2O from the MI. Danyushevsky et al. (2002) conducted a detailed study of the effect of heating schedule on the measured Th. They found that Th decreases with decreasing heating rate and approaches a constant value. Using this kinetic technique, the “correct” heating rate for each host phase and melt composition can be determined. Using faster heating rates will result in measured Th that are too high, while using heating rates that are slower than that at which the Th levels off increases the likelihood that the inclusion composition will change as a result of diffusion of components out of (or into?) the inclusion (Qin et al. 1992, Danyushevsky et al. 2000, Gaetani & Watson 2000). One of the most detailed studies of the effect of heating rate, inclusion size and volatile content on the homogenization behavior of MI is that of Thomas (1994a). Thomas used this information to develop a method to determine a minimum homogenization temperature for any given MI, which he interpreted to represent the trapping temperature of the MI.

plastically during heating to produce a volume that was larger than the original MI volume (Bodnar 2003b). How does one test the inclusions in a FIA to determine if “Roedder’s Rules” are satisfied? Studies of FI have shown rather conclusively that when FI re-equilibrate following entrapment the microthermometric and chemical properties of the FI show a wide range compared to inclusions that have not changed (Bodnar 2003b). Thus, if all of the inclusions within an FIA have the same number of phases and in the same volume proportions when observed at room temperature, and if the temperatures of phase changes are the same in all of the inclusions in the assemblage (indicating that the inclusions all have the same composition), one can be reasonably certain that the inclusions record the original formation conditions. While one can usually determine if all of the FI in an FIA contain the same phases and are in the same volume proportions, this is usually not possible with crystallized MI because the individual phases within the MI cannot be discerned (see Fig. 1-3). As such, crystallized MI must be homogenized and analyzed to determine if all of the inclusions in the MIA have the same composition and phase behavior during heating. Techniques that are commonly used to analyze MI include electron microprobe, SIMS (Layne 2006), FTIR, Raman spectroscopy and laser ablation ICP–MS (Pettke 2006, Halter et al. 2006). All of these techniques (except LA–ICP–MS; Halter et al. 2002) require a homogeneous glass, thus necessitating that the MI be heated to homogenization and quenched. Several different techniques have been used to homogenize MI. These techniques can broadly be divided into those in which the MI are heated at one atmosphere confining pressure and those in which the MI are heated under an elevated confining pressure. These techniques may be further divided based on whether the MI are heated continuously in one step or are heated incrementally. The most commonly used technique to homogenize MI is to heat the inclusions at one atmosphere in a microscope-mounted heating stage (Clocchiatti 1975, Magakyan et al. 1993, Lowenstern 1994, Reyf 1997, Fedele et al. 2003). The heating is usually conducted in an inert or reduced gas (He±H2) atmosphere to prevent oxidation of the sample. Magakyan et al. (1993) heated MI in clinopyroxene from boninite in a one atmosphere stage and report an error of ±20°C for homogenization temperatures (Th) in the 900– 10


Rather than heating MI in a microscopemounted heating stage, some workers heated inclusions at one atmosphere in a tube furnace (Yang & Bodnar 1994, Webster et al. 1997, Raia et al. 2000, Thomas & Webster 2000, Stockstill et al. 2005). While this technique is similar to heating at one atmosphere in a microscope-mounted stage, it does not allow MI to be monitored continuously during heating to determine temperatures of phase changes and to look for “anomalous behavior” during heating. Some workers heated the sample in one step to the final temperature (Thomas & Webster 2000). Other workers (Yang & Bodnar 1994, Student & Bodnar 2004) heated the samples incrementally and observed the sample under the microscope between heating steps to monitor the melting behavior (Fig. 1-9). Heating MI at one atmosphere pressure, either in a microscope-mounted stage or in a tube furnace, works well for MI with low concentrations of volatiles. However, if the MI contain significant amounts of H2O (or CO2?) the inclusions will commonly decrepitate before Th is reached, owing to the high internal pressures generated during heating. As a result, many workers heat crystallized volatile-rich MI only until the solid phases have dissolved, producing an inclusion that contains melt plus vapor bubble (Fig. 1-10), or heat the MI in a pressure vessel (Skirius et al. 1990, Webster & Duffield 1991, Anderson et al. 2000, Audétat et al. 2000, Thomas et al. 2002, Student & Bodnar 2004) or in a pressurized microscope stage (Massare & Clocchiatti 1987). This technique minimizes decrepitation of volatile-rich MI, although

Anderson et al. (2000) reported that some CO2 was lost from MI as a result of heating at 800–900°C under ≈200 MPa of Ar pressure for 20 hours. Student & Bodnar (1999) investigated the effect of heating technique on the observed Th, using synthetic silicate MI trapped at known P–T conditions. These workers found that, regardless of the technique used, Th increased with increasing inclusion size (Fig. 1-11), suggesting that homogenization of MI is a diffusion-controlled process, as previously argued by Thomas (1994a) and Thomas et al. (1996). MI heated in one step at one atmosphere in a tube furnace produced Th that most closely matched known temperatures. Inclusions heated continuously in a one atmosphere stage with a heating rate of 1°C/minute were about 10–15°C higher than the correct temperature, and those heated continuously at 3°C/minute were about 25°C too high (Fig. 1-12). An additional complication associated with crystallized MI from magmatic–hydrothermal ore deposits is that the host phase is often altered and/or crosscut by numerous planes of aqueous FI. Many of these planes intersect MI, and promote the decrepitation of MI during heating. Student & Bodnar (2004) tested various methods to homogenize MI in quartz phenocrysts containing abundant planes of secondary FI, including heating in a microscope heating stage, heating in a one atmosphere vertical tube furnace, and heating under pressure in a cold-seal pressure vessel. The method that proved most satisfactory to homogenize crystallized MI in phenocrysts with abundant planes of aqueous FI involved heating

FIG. 1-10. Melting sequence of a crystallized MI in plagioclase in trachyte from Ponza, Italy. At room temperature the inclusion consists of a mass of intergrown, fine-grained crystals ± glass. By 1092°C most of the silicate minerals have melted leaving only a mass of fine-grained opaque minerals, which melt between 1092 and 1197°C. After quenching, the inclusion contains a homogeneous glass and a vapor bubble (from Fedele et al. 2003).

11


870 3°C/

860

min

Homogenization Temperature (°C)

1° C / m

in

850 Frequency 5

840

4

3

2

1

3 °C /mi n Thermal Cycling

830

Th= 819-839 °C n=16

820

810

Tu b e F u r n a c e

Tra p p i n g Te m p e r a t u r e 800 ± 5 °C

800

T h = 8 0 0 ± 1 0 °C n=18

790 1 °C /mi n T h = 8 0 6 - 8 5 3 ° C n = 1 6 3 °C /mi n T h = 8 1 9 - 8 6 3 ° C n = 1 6 780

0

10

20

30

50

40

60

70

80

90 100 110

Inclusion Area (mm2) FIG. 1-11. Effect of inclusion size and heating technique on the measured homogenization temperature (Th) of synthetic silicate MI. The Th ranges for heating in a tube furnace (at one atmosphere) and the thermal cycling experiments are shown by the shaded boxes, along with the experimental conditions. These experiments did not consider inclusion size. A histogram of frequency versus Th for the thermal cycling experiment is inset along the right side of the diagram. Homogenization temperature as a function of inclusion area (as viewed through the microscope) for continuous heating experiments (1°C/min and 3°C/min) are indicated with unfilled and filled circles, respectively. Black vertical tie lines link data for the same inclusion. The solid line drawn through the data points for the continuous heating experiments shows the general relationship between inclusion size and Th, with Th approaching the trapping temperature as inclusion size approaches zero. The Th for the tube furnace experiment coincides most closely with the known trapping temperature (from Student & Bodnar 1999).

12


400

(a)

350

Pressure (MPa)

Haplogranite minimum curve

Entrapment P-T estimate

300 250

Trapping Conditions 800 °C and 200 MPa

200

e

or

uid

150

h oc

is

fl Cl

Fluid inclusion Th

100

%

.9 11

wt

Na

Melt inclusion Th range

50

200

300

400

500

600

700

800

900

1000

Temperature °C

Pressure (MPa)

230

(b)

3 °C/min Thermal Cycling 3 °C/min Continuous Heating

220

1 °C/min Continuous Heating 210

Tube Furnace

11.9 wt% NaCl fluid isochore

200

Trapping Conditions 800 ± 5 °C, 200 ± 5 MPa

190 750

800

850

Temperature °C FIG. 1-12. Estimated P–T formation conditions for coexisting melt and aqueous synthetic inclusions, calculated using vapor/melt Th for the four different heating experiments. The top diagram shows the P–T estimate for the ≈3°C/min thermal cycling heating experiment, showing the aqueous inclusion liquid/vapor curve and isochore to illustrate the "intersecting isochore" technique (Roedder & Bodnar 1980) used to determine formation pressure. The bottom diagram shows the P–T estimates for four different heating experiments shown in Figure 1-11 (from Student & Bodnar 1999).

13


samples under an elevated confining pressure. Phenocrysts were placed in a 5 mm long, 5 mm OD platinum capsule and loaded into a pressure vessel. A small hole was punched into the capsule bottom and top to permit the argon pressure medium to freely enter the capsule at run conditions to prevent collapse of the capsule and possible crushing of the sample. The vessel was sealed and pressurized with argon to 30–50 MPa, then lowered into a preheated furnace. Once the pressure reached 100 MPa during heating, argon was continuously bled off such that the sample was heated isobarically to the final run temperature. The technique for homogenizing MI in a pressure vessel involves heating in one step to some elevated temperature and holding the sample at this temperature for a sufficient amount of time to assure homogenization of a significant proportion of the inclusions. In general, the smaller inclusions in the sample homogenize first (at lower temperature and/or shorter run durations). Thomas et al. (1996) observed similar behavior. To determine the minimum run time and temperature required to homogenize a significant portion of the inclusions while at the same time minimizing the number of MI that decrepitate, a few phenocrysts from one sample were heated incrementally and examined after each heating step to determine the homogenization progress (Fig. 1-9). In most samples, a large portion of the MI homogenized over a range of a few tens of degrees – at this point the heating experiment was stopped. If heating were continued in an attempt to homogenize all inclusions in the phenocryst, those with lower Th would have decrepitated. Once the minimum Th was determined in this way, several phenocrysts from the same sample were heated to that temperature in one step and held for 24 hours. Heating under pressure significantly reduced decrepitation of MI. However, even under these conditions the larger inclusions (greater than about 30–50 µm) still decrepitated. Other workers (Sterner & Bodnar 1989, Skirius et al. 1990, Schmidt et al. 1998) have previously shown that heating FI or MI under confining pressure eliminates (or minimizes) decrepitation. Additionally, Massare & Clocchiatti (1987) reported that, when MI are heated in a pressurized microscope heating stage, Th of rhyolitic MI in sanidine decrease by 70°C/100 MPa in the temperature range 560– 850°C. As noted above, MI containing H2O-rich compositions often decrepitate, even when heated

under an elevated pressure. Decrepitation results because the internal pressure in the MI exceeds the strength of the host mineral (Bodnar 2003b). Thomas (1994b) described qualitatively the P–T path followed by H2O-rich MI during heating to the solidus and during melting, and Student & Bodnar (1996) quantified the effect of H2O on the P–T path followed by MI during cooling (or heating). The partial molar volume of H2O in hydrous melts is less than the molar volume of H2O in the vapor phase, and this difference becomes greater at lower pressures where the molar volume of H2O is large. As a result, as H2O exsolves from the melt phase in an MI of constant volume (i.e., assumes the volume of the host phase does not change with changing temperature and pressure), the pressure in the inclusion increases. For example, if an MI is trapped on the H2O-saturated solidus at 50 MPa and 782°C, H2O will begin to exsolve from the melt as the sample cools and the MI begins to crystallize feldspar and quartz (Fig. 1-13A). [This assumes that phase equilibrium is maintained during cooling. While synthetic MI do appear to maintain equilibrium during heating and cooling, it is unknown whether this applies to natural MI.] The pressure in the MI will continue to increase from 50 MPa as crystallization proceeds and the P–T path follows the H2O-saturated solidus, reaching a maximum pressure of about 150 MPa when the MI is completely crystallized at 708°C (Fig. 1-14A, point II). With further cooling the MI will follow the isochore corresponding to the H2O density in the crystallized inclusion (path II – I, Fig. 1-14A). The high pressures generated as the MI cools along the H2O-saturated solidus may cause the quartz host surrounding the MI to fracture with loss of H2O. The relative pressure increase during cooling decreases with increasing trapping pressure owing to the smaller difference between the partial molar volume of H2O in the melt and the molar volume of the free H2O phase. Thus, the pressure in a MI trapped on the H2O-saturated solidus at 200 MPa and 682°C increases to only about 230 MPa as the sample cools along the H2O-saturated solidus (Fig. 1-14B). Water-rich MI may not decrepitate during cooling in nature, owing to the elevated confining pressure. However, when heated in the laboratory to homogenize the MI in preparation for later analysis, the MI will follow the reverse of the P–T path followed during cooling in nature. Thus, the MI trapped on the H2O-saturated solidus at 50 MPa and 782°C will generate an internal pressure of about 14


FIG. 1-13. Calculated phase behavior of crystallized MI that trapped an H2O-saturated melt in the haplogranite system at 50 (A) and 200 (B) MPa. The calculated pressure in the MI during cooling from trapping conditions (and heating to homogenization) assumes a partial molar volume for H2O in the melt of 22 cm3/mole. The heating sequences A & B correspond to the P–T paths A & B, respectively, shown in Fig. 1-14, and the heating increments labeled I – V correspond to the P–T points I – V on Fig. 1-14. See text for additional details. (from Student & Bodnar 1996).

FIG. 1-14. Calculated P–T path during heating of crystallized MI that trapped an H2O-saturated melt in the haplogranite system at 50 (A) and 200 (B) MPa. The calculated pressure in the MI during cooling from trapping conditions (and heating to homogenization) assumes a partial molar volume for H2O in the melt of 22 cm3/mole. The P–T paths shown on A & B correspond to the heating sequences A & B, respectively, shown in Fig. 1-13, and the P–T points labeled I – V correspond to the heating increments I – V on Fig. 1-13. See text for additional details. (from Student & Bodnar 1996).

15


inclusions within the FIA, and is confirmed by consistent microthermometric results from all FI in the assemblage. For crystallized MI, it is generally not possible to determine the phase relations of the MI owing to the poor optics of MI and host mineral. In this case, the inclusions within the MIA should be tested for consistency in microthermometric and compositional data. As an example, Figure 1-15 shows a quartz phenocryst from the Red Mountain, Arizona, porphyry copper deposit with a zone of primary MI trapped along a growth (or possibly resorbtion) surface. A portion of this phenocryst was heated under 100 MPa confining pressure as described above, and the phase relations of 86 MI were monitored. It was previously determined (Student & Bodnar 2004) that the MI were trapped on the H2O-saturated solidus at about 810°C. Thus, when heated, the last daughter mineral (in this case a feldspar crystal) and the vapor bubble should disappear at the same temperature (i.e., similar to the MI in Fig 1-13A), which also represents the trapping temperature. Thirty-seven of the 86 inclusions from the MIA showed this behavior (Fig. 1-15), indicating that these 37 inclusions trapped only the H2O-saturated melt phase and did not reequilibrate after trapping. This conclusion is based on the fact that any change in composition and/or volume of the MI would have produced variability in the mode and temperature of homogenization (Bodnar 2003b). Thirty of the 86 inclusions contained a feldspar crystal of varying

230 MPa before melting begins (i.e., before the path intersects the H2O-saturated solidus) (point II, Fig. 1-14A). Once melting begins, the path follows the H2O-saturated solidus and the pressure in the inclusion decreases with continued heating as H2O dissolves into the melt phase. SUMMARY OF PETROGRAPHY AND MICROTHERMOMETRY The first step in any MI study is to identify a melt inclusion assemblage, or MIA, that represents a group of inclusions that were all trapped at the same time. By extension, this requirement implies that all of the MI in the assemblage trapped a melt of the same composition and at the same temperature and pressure. It is important to emphasize that a MIA is identified based initially on petrographic analysis of the sample, and not on MI compositions. Subsequent analysis of the MI in the assemblage can help to confirm that the inclusions being studied do indeed represent a MIA, and to distinguish between those MI that trapped a single, homogeneous melt and maintained the original composition during later cooling in nature and heating in the lab, and those that have not. After an MIA has been identified, the next step is to determine if the MI record the original physical and chemical conditions of trapping. For FI, this is most easily accomplished based on observation of consistent phase relations in all the

FIG. 1-15. Results of heating experiments on 86 crystallized MI from a single growth (or resorbtion) zone in a quartz phenocryst from the Red Mountain, Arizona, porphyry copper deposit. See text for details.

16


sizes after heating to 810°C. These inclusions are interpreted to have trapped H2O-saturated melt plus a feldspar crystal. Twelve of the MI contained both feldspar and a vapor bubble after heating to 810°C, indicating that these inclusions either trapped feldspar and vapor along with the melt, or reequilibrated, following entrapment. The remaining 7 MI had various combinations of melt, crystals and vapor bubble present after heating to 810°C, again suggesting mixed trapping and/or reequilibration. Thus, analyses of any of these 49 MI that did not show simultaneous dissolution of feldspar and vapor bubble at 810°C would provide a melt composition that is not representative of the melt that was present at the time of trapping. Random selection of MI in this MIA would provide erroneous (and misleading) information if the behavior of the MI during heating was not monitored before analyses were conducted. Observations of MI behavior during laboratory heating thus provide a test of assumptions regarding the timing of inclusion trapping and possible reequilibration. The protocol that one should follow in the selection of MI for study is outlined in Figure 1-8. To summarize, one should first select two or more MI that define a melt inclusion assemblage (MIA), based on petrographic examination of the sample. MI in the same growth zone or along a healed fracture are examples of MI occurrences that could be used to define an MIA. If MI cannot be related to a MIA, one should reconsider whether a MI study should be undertaken, as it is unlikely that one will be able to argue convincingly that the results obtained represent the melt that was present at the time of MI formation. After an MIA is identified, the MI should be examined to determine if they contain consistent phase relations (Fig. 1-8). If not, the MIA should be abandoned because the MI have either trapped mixtures of phases or have reequilibrated after trapping, or both. In either case, it is unlikely that the MI will provide useful information concerning the physical and chemical conditions in the magma. If the phase relations cannot be determined at room temperature because of poor optics or because the phases are too finegrained, the MI should be heated to homogenization. Assuming that the inclusions trapped a single, homogeneous melt phase, all MI should show similar temperatures of phase changes during heating. If the MI show variability in the order in which various phases disappear (i.e., solids and vapor bubble) and/or in the temperatures of

phase changes, the MIA should be abandoned because, as with the example of inconsistent phase relations described earlier, the MI do not represent the original melt that was present at the time of trapping. Those MI in the MIA that show similar phase behavior and temperatures of phase changes, such as the 37 MI in the Red Mountain sample described above, are most likely to have trapped the melt that was present and have not reequilibrated following trapping. It is these inclusions, and only these inclusions, that should be selected for further analysis and interpretation. By applying these simple tests and selecting only inclusions that satisfy the criteria outlined above, one can obtain data that accurately reflect conditions in the magma at the time that the MI were trapped. MELT INCLUSIONS IN PORPHYRY-TYPE DEPOSITS A plutonic environment in which MI have contributed significantly in recent years is in understanding ore-forming processes in porphyrytype deposits (see Student & Bodnar 2004 for review). Porphyry copper deposits are associated with epizonal siliceous intrusions emplaced at convergent margins. While FI studies have improved our understanding of the nature and role of magmatic–hydrothermal fluids in porphyry systems (Roedder 1971, Nash 1976, Bodnar 1995, Beane & Bodnar 1995, Roedder & Bodnar 1997, Davidson & Kamenetsky 2001), it has been only recently that workers have begun to study the melt MI in these systems. Student & Bodnar (2004) summarized a protocol for studying crystallized MI in samples that have undergone extensive subsolidus hydrothermal alteration, such as those in porphyry copper deposits. The technique produces glassy MI with a homogeneous composition that are amenable to analysis by a variety of techniques, including synchrotron XRF, PIXE, SIMS, electron microprobe, Raman spectroscopy, FTIR spectroscopy and laser ablation ICP–MS. Importantly, the ore metal content of MI representing different stages in the magmatic history can be determined, and relative differences between pre-, syn- and post-mineralization samples (Fig. 1-16) can be compared with models for metal partitioning between melt and coexisting magmatic–hydrothermal fluids (Candela & Holland 1986, Candela 1989, 1997, Cline & Bodnar 1991, Lynton et al. 1993: Candela & Piccoli 1995). 17


FIG. 1-16. Comparison of Zn and Cu concentrations in MI from the porphyry copper deposits at Red Mountain, Arizona, and Tyrone, New Mexico, with concentrations in MI from the White Island volcano, New Zealand. Metal concentrations in pre-mineralization quartz latite at Red Mountain and at White Island are higher than those in syn- and post-mineralization MI from Red Mountain and Tyrone. (from Student & Bodnar 2004). quartz

quartz

A

RM QL MI

B

MI RM D2 Ty QM MI ph c

ylli

ylli

ph c RMD2

potassic

albite

orthoclase

albite

orthoclase

FIG. 1-17. Compositions of MI from the porphyry copper deposits at Red Mountain, Arizona (A), and Tyrone, New Mexico (B). Hydrothermal fluids associated with phyllic alteration in porphyry copper deposits lie along the quartz–potassium feldspar (orthoclase) join near the quartz apex, whereas fluids associated with potassic alteration lie near the orthoclase apex. MI compositions at Red Mountain produce a trend (arrows, Fig. 1-17A) that projects towards the composition of fluids associated with phyllic alteration, whereas MI at Tyrone show trends (arrows, Fig. 1-17B) that project to both phyllic and potassic alteration fluids. (from Student & Bodnar 2004).

18


As noted above, MI in porphyry copper and related deposits are often intersected by one or more planes of FI. It is thought that most of these fractures occur at subsolidus conditions associated with hydrothermal alteration and may affect the compositions of MI. This interpretation is supported by the observation that compositions of MI fall along trends that project to compositions of fluids associated with potassic and/or phyllic alteration in porphyry copper deposits (Fig. 1-17). An interesting application of MI is to compare the geochemistry of magmatic systems that host (or have the potential to host) economic mineralization and those that are barren or subeconomic. As an example, Rapien et al. (2003) studied MI from the White Island, New Zealand, volcano and compared the results to bulk rock compositions in productive and barren porphyry intrusions. These workers concluded that the White Island magma has the potential to generate economic porphyry copper-type mineralization, but that the magmatic system has not evolved to the productive stage, i.e., the magmatic–hydrothermal system is too “young”. Similarly, Audétat & Pettke (2003) studied MI and coexisting FI from two barren plutons in New Mexico, USA. They concluded that the absence of mineralization was related to the low salinity of the exsolving magmatic fluids, resulting in less efficient extraction of metals from the melt. Grancea et al. (2001) studied MI in mineralized and barren intrusions in Romania and found that MI in mineralized systems were enriched in S and had a lower Al/(K+Na+2Ca) compared to barren systems. Kamenetsky et al. (1999) studied mixed silicate glass and crystalline silicate–sulfate– carbonate–sulfide–oxide inclusions from the Dinkidi Cu–Au porphyry deposit, Philippines. The inclusions are enriched in ore metals, and they interpreted the inclusions to have originally formed early in the ore-forming process as a result of immiscibility. Harris et al. (2003) observed coexisting silicate MI and high-salinity and vaporrich FI in magmatic-hydrothermal quartz veins from the Bajo de la Alumbrera porphyry copper deposit, Argentina. The close association of these three different inclusion types was interpreted to represent melt–aqueous fluid immiscibility, and compositions of the coexisting inclusions were used to calculate bulk partition coefficients for ore metals between the melt and magmatic aqueous phase. Schmitt et al. (2002) studied MI in peralkaline

granites from the Amis Complex in Namibia. The MI are enriched in Nb and REE and proved that the peralkaline composition and rare metal enrichments are primary magmatic features and not the result of later hydrothermal activity. SUMMARY This chapter attempts to summarize a methodology to recognize, select and study MI in plutonic rocks. In general, MI from this environment are small, mostly to completely crystallized, and difficult to recognize during normal petrographic observations. Crystallized MI must be homogenized to confirm that all MI in a melt inclusion assemblage show similar modes and temperatures of phase changes and, thus, likely trapped samples of the original homogeneous melt and did not reequilibrate following entrapment. The technique that proves most reliable to homogenize crystallized, volatile-rich MI is to heat the samples in one step in a pressure vessel under an elevated confining pressure. The most important take-home message from this chapter concerns the selection of MI to study. Numerous studies of natural and synthetic FI and MI show clearly that the best evidence that inclusions have trapped a single homogeneous phase and have not reequilibrated following entrapment is if all of the MI or FI in an assemblage show the same phase behavior and temperatures of phase changes. Trapping mixtures of phases, leakage, or change in the volume of the inclusion, all result in a wide range in phase relations and temperatures of phase changes. Thus, if a group of MI that were all trapped at the same time (MIA) show similar phase behavior and temperatures of phase changes (including homogenization), one can have a high level of confidence that the MI provide information on the physical and chemical environment at the time of trapping. ACKNOWLEDGEMENTS During the past decade many students, colleagues and visitors to the Fluids Research Laboratory at Virginia Tech have contributed to our understanding of melt inclusions in plutonic environments. We wish to acknowledge contributions from Andreas Audétat, Claudia Cannatelli, Benedetto DeVivo, Luca Fedele, John Mavrogenes, Maria Rapien, Nobu Shimizu, Karen Stockstill, Csaba Szabo, Jay Thomas, Rainer 19


equilibrium crystallization of magmas, according to data obtained by studies of melt inclusions. Soviet Geology and Geophysics (Geologiya i Geofizika) 25, no. 8, 73–81.

Thomas and Kyounghee Yang. Jim Reynolds led us to appreciate the importance of petrography in the selection of fluid and melt inclusions. The authors also thank Fred Anderson, Ilya Veksler and Jim Webster for comments and suggestions on an earlier version of this manuscript. Funding for recent work on FI and MI was provided by NSF Grants EAR0001168, EAR-0125918, and EAR-0337094 to RJB.

BADANINA, E.V., VEKSLER, I.V., THOMAS, R., SYRITSO, L.F. & TRUMBULL, R.B. (2004): Magmatic evolution of Li–F, rare-metal granites: A case study of melt inclusions in the Khagnilay complex, Eastern Transbaikalia (Russia). Chem. Geol. 210, 113–134.

REFERENCES

BEANE, R.E. & BODNAR, R.J. (1995): Hydrothermal fluids and hydrothermal alteration in porphyry copper deposits. In Pierce, F.W., and Bohm, J.G., eds., Porphyry Copper Deposits of the American Cordillera. Arizona Geological Society Digest 20, Tucson, Arizona, pp. 83–93.

ANDERSON, A.T., JR. (1973): The before-eruption water content of some high-alumina magmas. Bull. Volcanology 37, 530–552. ANDERSON, A.T., JR. (1974): Evidence for a picritic, volatile-rich magma beneath Mt. Shasta, California. J. Petrol. 15, 243–267.

BODNAR, R.J. (1995): Fluid inclusion evidence for a magmatic source for metals in porphyry copper deposits: In Magmas, Fluids and Ore Deposits (J.F.H. Thompson, ed.) Mineral. Assoc. Can. Short Course 23, 139–152.

ANDERSON, A.T., JR. (1991): Hourglass inclusions: Theory and application to the Bishop rhyolitic Tuff. Am. Min. 76, 530–547. ANDERSON, A.T., JR. (2003): Melt (glass ±crystals) inclusions. In Fluid Inclusions: Analysis and Interpretation (I. Samson, A. Anderson & D. Marshall, eds.) Mineral. Assoc. Can. Short Course 32, 353–364.

BODNAR, R.J. (2003a): Introduction to fluid inclusions. In Fluid Inclusions: Analysis and Interpretation (I. Samson, A. Anderson & D. Marshall, eds.) Mineral. Assoc. Can. Short Course 32, 1–8.

ANDERSON, A.T., JR., DAVIS, A.M. & LU, F. (2000): Evolution of Bishop Tuff rhyolitic magma based on melt and magnetite inclusions and zoned phenocrysts. J. Petrol. 41, 449–473.

BODNAR, R.J. (2003b): Reequilibration of fluid inclusions. In Fluid Inclusions: Analysis and Interpretation (I. Samson, A. Anderson & D. Marshall, eds.) Mineral. Assoc. Can. Short Course 32, 213–230.

AUDÉTAT, A., GÜNTER, D. & HEINRICH, C.A. (2000): Magmatic–hydrothermal evolution in a fractionating granite: A microchemical study of the Sn–W–F-mineralized Mole Granite (Australia). Geochim. Cosmochim. Acta 64, 3373–3393.

BODNAR, R.J. & STERNER, S.M. (1987): Synthetic fluid inclusions. In Hydrothermal Experimental Techniques, G.C. Ulmer and H.L. Barnes, eds., Wiley–Interscience, New York. p.423–457. BRANTLEY, S.L. (1992): The effect of fluid chemistry on quartz microcrack lifetimes. Chem. Geol. 113, 145–156.

AUDÉTAT, A. & PETTKE, T. (2003): The magmatichydrothermal evolution of two barren granites: a melt and fluid inclusion study of the Rio de Medio and Cañada Pinatebe plutons in northern New Mexico (USA). Geochim. Cosmochim. Acta 67, 97–121.

BREITER, K., FRYDA, J., SELTMANN, R. & THOMAS, R. (1997): Mineralogical evidence for two magmatic stages in the evolution of an extremely fractionated P-rich rare-metal granite: the Podlesi stock, Krusne hory, Czech Republic. J. Petrol. 38 1723–1739.

BACON, C.R. (1989): Crystallization of accessory phases in magmas by local saturation adjacent to phenocrysts. Geochim. Cosmochim. Acta 53, 1055–1066.

CANDELA, P.A. (1989): Magmatic ore-forming fluids: Thermodynamic and mass transfer calculations of metal concentrations, In Ore Deposition Associated with Magmas (J.A.

BAKUMENKO, I.T., BAZAROVA, T.YU., PANINA, L.I., SIMONOV, V.A. & KOSUKHIN, O.I. (1984): Distinctive features of the equilibrium and non-

20


Whitney & A.J. Naldrett eds.) Reviews in Economic Geology 4, Economic Geology, El Paso, TX (223–233).

DANYUSHEVSKY, L.V., MCNEILL, A.W. & SOBOLEV, A.V. (2002): Experimental and petrological studies of melt inclusions in phenocrysts from mantle-derived magmas: an overview of techniques, advantages and complications. Chem. Geol. 183, 5–24.

CANDELA, P.A. (1997): A review of shallow, orerelated granites: Textures, volatiles and ore metals. J. Petrol. 38, 1619–1633.

DAVIDSON, P. & KAMENETSKY, V.S. (2001): Immiscibility and continuous felsic melt–fluid evolution within the Rio Blanco porphyry system, Chile: Evidence from fluid inclusions in magmatic quartz. Econ. Geol. 96, 1921–1929.

CANDELA, P.A. & HOLLAND, H.D. (1986): A mass transfer model for copper and molybdenum in magmatic hydrothermal systems: Origin of porphyry-type ore deposits. Econ. Geol. 81, 1–19. CANDELA, P.A. & PICCOLI, P.M. (1995): Model oremetal partitioning from melts into vapor and vapor–brine mixtures. In Thompson, J.F.H. (ed.), Magmas, Fluids and Ore Deposits, Mineralogical Association of Canada, Short Course 23, 101– 127.

DAVIDSON, P., KAMENETSKY, V., COOKE, D.R., FRIKKEN, P., HOLLINGS, P., RYAN,C., VAN ACHTERBERGH, E., MERNAGH, T., SKARMETA, J., SERRANO, L. & VARGAS, R. (2005): Magmatic precursors of hydrothermal fluids at the Rio Blanco Cu–Mo deposit, Chile: links to silicate magmas and metal transport. Econ. Geol. 100, 963–978.

CESARE, B., MARCHESI, C., HERMANN, J. & GOMEZPUGNAIRE, M.T. (2003): Primary melt inclusions in andalusite from anatectic graphitic metapelites; implications for the position of the Al2SiO5 triple point. Geology 31, no. 7, 573–576.

DIETRICH, A., LEHMANN, B. & WALLIANOS, A. (2000): Bulk rock and melt inclusion geochemistry of Bolivian tin porphyry systems. Econ. Geol. 95, 313–326.

CHUPIN, V.P., BAKUMENKO, I.T. & SOBOLEV, V.S. (1979): Crystallized melt inclusions in quartz or rapakivi granite. Transactions (Doklady) of the U.S.S.R. Academy of Sciences: Earth Science 248, no. 1–6, 109–111

DOWTY, E. (1980): Crystal growth and nucleation theory and the numerical simulation of igneous crystallization. In Physics of Magmatic Processes (R.B. Hargraves, ed.), Princeton University Press, Princeton, 419–485.

CHUPIN, V.P., SOROKINA, G.A. & BARTON JR., J.M. (2001): Late Archean melt inclusions in zircon from charnockitic orthogneiss of Bulai Pluton (Limpopo Belt, South Africa) (abs.). Proceedings of ECROFI XVI, Porto 2001, 93–94.

FEDELE, L., BODNAR, R.J., DEVIVO, B. & TRACY, R.J. (2003): Melt inclusion geochemistry and computer modeling of trachyte petrogenesis at Ponza, Italy. Chem. Geol. 194, 81–104.

CLINE, J.S. & BODNAR, R.J. (1991): Can economic porphyry copper mineralization be generated by a typical calc-alkaline melt? J. Geophys. Res. 96, 8113–8126.

FREZZOTTI, M.L. (1992): Magmatic immiscibility and fluid phase evolution in the Mount Genis granite (southeastern Sardinia, Italy). Geochim. Cosmochim. Acta 56, 21–33.

CLOCCHIATTI, R. (1975): Les inclusions vitreuses des cristaux de quartz. Étude optique, thermooptique et chimique. Applications geologiques. [Glass (melt) inclusions in quartz. Optical, microthermometric and chemical study. Geologic applications] Memoires de la Société Géologique de France, LIV, 122, 1–96 [in French].

FREZZOTTI, M.L. (2001) Silicate melt inclusions in magmatic rocks: applications to petrology. Lithos 55, 273–299. GAETANI, G.A. & WATSON, E.B. (2000): Open system behavior of olivine-hosted melt inclusions. Earth & Planet. Sci. Letters 183 27–41.

DANYUSHEVSKY, L.V., DELLA-PASQUA, F.N., & SOKOLOV, S. (2000): Re-equilibration of melt inclusions trapped by magnesian olivine phenocrysts from subduction-related magmas: petrological implications. Contr. Min. Pet. 138, 68–83.

GOLDSTEIN, R.H. & REYNOLDS, T.J. (1994): Systematics of Fluid Inclusions in Diagenetic Minerals. Society for Sedimentary Geology Short Course 31, Tulsa, Oklahoma, 199 pp.

21


KAMENETSKY, V.S. & NAUMOV, V.B. (2004): Immiscibility between silicate magmas and aqueous fluids; a melt inclusion pursuit into the magmatic–hydrothermal transition in the Omsukchan Granite; NE Russia. Chem. Geol. 210, 73–90.

GRANCEA, L., CUNEY, M. & LEROY, J.L. (2001): Mineralized versus barren intrusions: a melt inclusion study in Romania’s gold quadrilateral. C. R. Acad. Sci. Paris, Sciences de la Terre et des planets 333, 705–710. HAAPALA, I. & THOMAS, R. (2000): Melt inclusions in quartz and topaz of the topaz granite from Eurajoki, Finland. J. Czech. Geol. Soc. 45, 149– 154.

KAMENETSKY, V.S., WOLFE, R.C., EGGINS, S.M., MERNAGH, T.P. & BASTRAKOV, E. (1999): Volatile exsolution at the Dinkidi Cu–Au porphyry deposit, Philippines: A melt–inclusion record of the initial ore-forming process. Geology 27, 691–694.

HALTER, W.E. & HEINRICH, C.A. (2005): Magmatic processes and volatile phase generation in porphyry-type environments: A laser ablation– ICP–MS study of silicate and sulfide melt inclusions. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 151-164.

KOVALENKO, V.I., TSARYEVA, G.M., GOREGLYAD, A.V., YARMOLYUK, V.V., TROITSKY, V.A., HERVIG, R.L. & FARMER, G.L. (1995): The peralkaline granite-related Khaldzan–Buregtey rare-metal (Zr, Nb, Ree) deposit, Western Mongolia. Econ. Geol. 95, 313–326.

HALTER, W.E., HEINRICH C.A. & PETTKE T. (2004b): Laser ablation ICP–MS analysis of silicate and sulfide melt inclusions in an andesitic complex II: evidence for magma mixing and magma chamber evolution. Contr. Min. Pet. 147, 397–412.

LANDTWING, M.R., PETTKE, T., HALTER, W.E., HEINRICH, C.A., REDMOND, P.A., EINAUDI, M.T. & KUNZE, K. (2005): Copper deposition during quartz dissolution by cooling magmatic– hydrothermal fluids: The Bingham porphyry. Earth & Planetary Sci. Letters 235, Issues 1–2, 229–243

HALTER, W.E., PETTKE T. & HEINRICH C.A. (2004a): Laser ablation ICP–MS analysis of silicate and sulfide melt inclusions in an andesitic complex I: analytical approach and data evaluation. Contr. Min. Pet. 147, 385–396.

LAYNE, G.D. (2006): Application of secondary ion mass spectrometry to the determination of traditional and non-traditional light stable isotopes in melt inclusions. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 27-49.

HALTER W.E., PETTKE T., HEINRICH C.A. & ROTHEN-RUTISHAUSER B. (2002): Major to trace element analysis of melt inclusions by laser ablation–ICP–MS; methods of quantification. Chem. Geol. 183, 63–86.

LOWENSTERN, J.B. (1994): Dissolved volatile contents in an ore-forming magma. Geology 22, 893–896.

HANSTEEN, T.H. & LUSTENHOUWER, W.J. (1990): Silicate melt inclusions from a mildly peralkaline granite in the Oslo paleorift, Norway. Min. Mag. 54, 195–205.

LOWENSTERN, J.B. (1995): Applications of silicate– melt inclusions to the study of magmatic volatiles. In Magmas, Fluids and Ore Deposits (J.F.H. Thompson, ed.), Mineral. Assoc. Can. Short Course 23, 71–98.

HARRIS, C. (1986): A quantitative study of magmatic inclusions in the plutonic ejecta of Ascension Island. J. Petrol. 27, 251–276.

LOWENSTERN, J.B. (2003): Melt inclusions come of age: Volatiles, volcanoes, and Sorby’s legacy. In Developments in Volcanology 5, Melt Inclusions in Volcanic Systems: Methods, Applications and Problems, (B. DeVivo & R.J. Bodnar, eds.), Elsevier, Amsterdam, 1–21.

HARRIS, A.C., KAMENETSKY, V.S., WHITE, N.C., VAN ACHTERBERGH, E. & RYAN, C.G. (2003): Melt inclusions in veins: Linking magmas and porphyry Cu deposits. Science 302, 2109–2111. HAURI, E.H., KENT, A.J.R. & ARNDT, N. (2001): Melt inclusions at the miilenium: toward a deeper understanding of magmatic processes. Chem. Geol. 183, 1–3.

LOWENSTERN, J.B., CLYNNE, M.A. & BULLEN, T.D. (1997): Comagmatic A-type granophyre and rhyolite from the Alid volcanic center, Eritrea, northeast Africa. J. Petrol. 38, 1707–1731. 22


LYNTON, S.J., CANDELA, P.A. & PICCOLI, P.M. (1993): An experimental study of the partitioning of copper between pyrrhotite and a high silica rhyolite melt. Econ. Geol. 88, 901–915.

inclusions. Am. Min. 77, 565–576. RAIA, F., WEBSTER, J.D. & DE VIVO, B. (2000): Pre-eruptive volatile contents of Vesuvius magmas: constraints on eruptive history and behavior. I. The medieval and modern interplinian activities. Eur. J. Mineral. 12, 179–193.

MAGAKYAN, R., SOBOLEV, A.V., ZAKARIADZE, G.S. & KONONKOVA, N.N. (1993): Petrology of evolved boninite magmas: evidence from the Mesozoic lesser Caucasus island arc. Petrology 1, 378–394.

RAPIEN, M.H., BODNAR, R.J., SIMMONS. S., SZABÓ. CS., WOOD, C.P. & SUTTON, S.R. (2003): Melt inclusion study of the embryonic porphyry copper system at White Island, New Zealand. Society of Economic Geologists Special Publication 10, 41– 59.

MASSARE, D. & CLOCCHAITTI, R. (1987): Influence de la pression de confinement sur les temperatures d’homogénéisation des inclusions fluids et vitreuses: premières données obtenues à l’aide d’une platine chauffante dams le domaine 0.0–400 MPa; 2008–50°C. C. R. Acad. Sci. Paris 305 1207–1210.

REYF, F.G. (1997): Direct evolution of W-rich brines from crystallizing melt within the Mariktikan granite pluton, west Transbaikalia. Mineral. Dep. 32 475–490.

MASSARE, D., METRICH, N. & CLOCCHAITTI, R. (2002): High-temperature experiments on silicate melt inclusions in olivine at 1 atm: inference on temperatures of homnogenization and H2O concentrations. Chem. Geol. 183, 87–98.

REYF, F.G. (2004): Immiscible phases of magmatic fluid and their relation to Be and Mo mineralization at the Yermakovka F–Be deposit, Transbaikalia, Russia. Chem. Geol. 210, 49–71. REYF, F.G., SELTMANN, R. & ZARZISKY, G.P. (2000): The role of magmatic processes in the formation of banded, Li,F-enriched granites from the Orlovka tantalum deposit, Transbaikalia, Russia: Microthermometric evidence. Can. Mineral. 38, 915–936.

NASH, T.J. (1976): Fluid-inclusion petrology – data from porphyry copper deposits and applications to exploration: U. S. Geological Survey Prof. Paper 907–D, 16 pp. PEPPARD, B.T., STEELE, I.M., DAVIS, A.M., WALLACE, P.J. & ANDERSON, A.T., JR. (2001): Zoned quartz phenocrysts from the rhyolitic Bishop Tuff. Am. Min. 86, 1034–1052.

ROEDDER, E. (1971): Fluid inclusion studies on the porphyry-type ore deposits at Bingham, Utah, Butte, Montana, and Climax, Colorado: Econ. Geol., 66, 98–120.

PERETYAZHKO, I.S., ZAGORSKY, V.Y., SMIRNOV, S.Z. & MIKHAILOV, M.Y. (2004): Conditions of pocket formation in the Oktyabrskaya tourmalinerich gem pegmatite (the Malkhan field, Central Transbaikalia, Russia). Chem. Geol. 210, 91–111.

ROEDDER, E. (1979): Origin and significance of magmatic inclusions: Bull. de Miner. 102, 487– 510. ROEDDER, E. (1984): Fluid Inclusions. Mineralogical Society of America, Reviews in Mineralogy, 12, 646 pp.

PETTKE, T. (2006): In situ laser ablation-ICP-MS chemical analysis of melt inclusions and prospects for constraining subduction zone magmatism. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 51-80.

ROEDDER, E. (1992): Fluid inclusion evidence for immiscibility in magmatic systems. Geochim. Cosmochim. Acta 56, 5–20.

PROKOFIEV, V., KAMENETSKY, V.S., KOVALENKER, V., BODON, S.B. & JELEN, S. (1999): Evolution of magmatic fluids at the Banska Stiavnica precious and base metal deposit, Slovakia – Evidence from melt and fluid inclusions. Econ. Geol. 97, 949– 955.

ROEDDER, E. & BODNAR, R.J. (1980): Geologic pressure determinations from fluid inclusion studies. Ann. Rev. Earth Planet. Sci. 8, 263–301. ROEDDER, E. & BODNAR, R.J. (1997): Fluid Inclusion Studies of Hydrothermal Ore Deposits. in Geochemistry of Hydrothermal Ore Deposits, 3rd ed., H. L. Barnes, ed., Wiley & Sons, Inc, New York (657–698).

QIN, Z., LU, F. & ANDERSON, A.T., JR (1992): Diffusive reequilibration of melt and fluid

23


ROEDDER, E. & COOMBS, D.S. (1967): Immiscibility in granitic melts, indicated by fluid inclusions in ejected granitic blocks from Ascension Island. J. Petrol. G. B. 8, 417–451.

Virginia Polytechnic Institute and State University, Blacksburg, VA. 119 p. http://scholar. lib.vt.edu/theses/available/etd-08192002-150340/ STUDENT, J.J. & BODNAR, R.J. (1996): Melt inclusion microthermometry: Petrologic constraints from the H2O-saturated haplogranite system. Petrology, 4, 291–306.

SCHIANO, P. (2003): Primitive mantle magmas recorded as silicate melt inclusions in igneous minerals. Earth Sci. Reviews 63, 121–144.

STUDENT, J.J. & BODNAR, R.J. (1999): Synthetic Fluid Inclusions XIV: Microthermometric and compositional analysis of coexisting silicate melt and aqueous fluid inclusions trapped in the haplogranite–H2O–NaCl–KCl system at 800°C and 2000 bars. J. Petrol. 40, 1509–1525.

SCHMIDT, C., CHOU, I-M., BODNAR, R.J. & BASSETT, W.A. (1998): Microthermometric analysis of synthetic fluid inclusions in the hydrothermal diamond-anvil cell. Am. Min. 83, 995–1007. SCHMITT, A.K., TRUMBULL, R.B., DULSKI, P & EMMERMANN, R. (2002): Zr–Nb–REE mineralization in peralkaline granites from the Amis Complex, Brandberg (Namibia): Evidence for magmatic pre-enrichment from melt inclusions. Econ. Geol. 97, 399–413.

STUDENT, J.J. & BODNAR, R.J. (2004): Silicate melt inclusions in porphyry copper deposits: Identification and homogenization behavior. Can. Min. 42 1583–1599. TAKENOUCHI, S. & IMAI, H. (1975): Glass and fluid inclusions in acidic igneous rocks from some mining areas in Japan. Econ. Geol. 70, 750–769.

SKIRIUS, C.M., PETERSON, J.W & ANDERSON, A.T. JR. (1990): Homogenizing rhyolitic glass inclusions from the Bishop Tuff. Am. Min. 75, 1381–1398.

THOMAS, J.B., BODNAR, R.J., SHIMIZU, N. & SINHA, A.K. (2002): Determination of zircon/melt trace element partition coefficients from SIMS analysis of melt inclusions in zircon. Geochim. Cosmochim. Acta, 66, 2887–2902.

SOBOLEV, A.V. (1996): Melt inclusions in minerals as a source of principal petrologic information. Petrology 4, 228–239.

THOMAS, R. (1988) Results of investigations on melt inclusions in various magmatic rocks from the northern border of the Bohemian Massif. Proceedings of the 1st international Conference on the Bohemian Massif, Prague, Czechoslovakia, Sept. 26–Oct. 3, 1988, 298–306.

SOBOLEV, V.S. & KOSTYUK, V.P. (1975): Magmatic crystallization based on a study of melt inclusions. Fluid Inclusion Research 9, 182–253 (translated from original publication in Russian). SORBY, H.C. (1858): On the microscopical structure of crystals, indicating origin of minerals and rocks. Quart. J. Geol. Soc. London 14, 453–500.

THOMAS, R. (1994a): Estimation of viscosity and water content of silicate melts from melt inclusion data. Eur. J. Mineral. 6, 511–535.

STERNER, S.M. & BODNAR, R.J. (1984): Synthetic fluid inclusions in natural quartz. I. Compositional types synthesized and applications to experimental geochemistry. Geochim. Cosmochim. Acta 48, 2659–2668.

THOMAS, R. (1994b): Fluid evolution in relation to the emplacement of the Variscan granites in the Erzgebirge region: A review of the melt and fluid inclusion evidence. In Metallogeny of collisional orogens focussed on the Erzgebirge and comparable metallogenetic settings (Seltmann, R., Kampf, H. & Moller, P., eds.) Czech Geol. Survey, pp. 70–81.

STERNER, S.M. & BODNAR, R.J. (1989): Synthetic fluid inclusions. VII. Reequilibration of fluid inclusions in quartz during laboratory-simulated metamorphic uplift. J. Met. Geol. 7, 243–260. STOCKSTILL, K.R., MCSWEEN, H.Y., JR. & BODNAR, R.J. (2005): Melt inclusions in augite of the Nakhla Martian meteorite: Evidence for basaltic parental melt. Meteor. Planet. Sci. 40, 377–395.

THOMAS, R. (1994c): Thermometric study on granites from Bornholm Island, Denmark. Z.Geol. Wiss. 22, 139–145. THOMAS, R., FOERSTER, H.-J. & HEINRICH, W. (2003): The behavior of boron in a peraluminous granite–pegmatite system and associated hydro-

STUDENT, J.J. (2002): Silicate melt inclusions in igneous petrogenesis. Unpub. Ph.D. diss., 24


thermal solutions: A melt and fluid inclusion study. Contrib. Mineral. Petrol. 144, 457–472.

VEKSLER, I.V. & LENTZ, D. (2006): Parental magmas of plutonic carbonatites, carbonatesilicate immiscibility and decarbonation reactions: evidence from melt and fluid inclusions. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 123-149.

THOMAS, R., FOERSTER, H.-J., RICKERS, K. & WEBSTER, J.D. (2005): Formation of extremely F-rich hydrous melt fractions and hydrothermal fluids during differentiation of highly evolved tingranite magmas: a complex melt/fluid inclusion study. Mineral. Dep. 148, 582–601.

WEBSTER, J.D. (2004): The exsolution of magmatic hypersaline chloride liquids. Chem. Geol. 210, 33–48.

THOMAS, R. & KLEMM W. (1997): Microthermometric study of silicate melt inclusions in Variscan granites from SE Germany: Volatile contents and entrapment conditions. J. Petrol. 38, 1753–1765.

WEBSTER, J.D. & DUFFIELD, W.A. (1991): Volatiles and lithophile elements in Taylor Creek rhyolite: constraints from glass inclusion analysis. Am. Min. 76, 1628–1645.

THOMAS, R., RHEDE, D. & TRUMBULL, R.B. (1996): Microthermometry of volatile-rich silicate melt inclusions in granitic rocks. Z. Geol. Wiss. 24, 505–526.

WEBSTER, J.D. & REBBERT, C.R. (2001): The geochemical signature of fluid-saturated magma determined from silicate melt inclusions in Ascension Island granite xenoliths. Geochim. Cosmochim. Acta 65, 123–136.

THOMAS, R. & WEBSTER, J.D. (2000): Strong tin enrichment in a pegmatite-forming melt. Mineral. Dep. 35, 570–582.

WEBSTER, J.D., THOMAS, R., FOERSTER, H.-J., SELTMANN, R. & TAPPEN, C. (2004): Geochemical evolution of halogen-enriched, granite magmas and mineralizing fluids of the Zinnwald tin–tungsten mining district, Erzgebirge, Germany. Mineral. Dep. 39, 452– 472.

THOMAS, R., WEBSTER, J.D. & DAVIDSON, P. (2006): Understanding pegmatite formation: the melt and fluid inclusion approach. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 189-210. THOMAS, R., WEBSTER, J.D. & HEINRICH, W. (2000): Melt inclusions in pegmatite quartz: Complete miscibility between silicate melts and hydrous fluids at low pressure. Contrib. Mineral. Petrol. 139, 394–401.

WEBSTER, J.D., THOMAS, R., RHEDE, D., FÖRSTER, H.-J. & SELTMANN, R. (1997): Melt inclusions in quartz from an evolved peraluminous pegmatite: Geochemical evidence for strong tin enrichment in fluorine-rich and phosphorus-rich residual liquids. Geochim. Cosmochim. Acta 61, 2589– 2604.

TOMIYA, A & TAKAHASHI, E. (2005): Evolution of the magma chamber beneath the Usu volcano since 1663: a natural laboratory for observing changing phenocryst compositions and textures. J. Petrol. 46, 2395–2426.

YANG, K. & BODNAR, R.J. (1994): Magmatic– hydrothermal evolution in the “bottoms” of porphyry copper systems: Evidence from silicate melt and aqueous fluid inclusions in granitoid intrusions in the Gyeongsang Basin, South Korea. Int. Geol. Rev. 36, 608–628.

VAPNIK, YE. (1998): Melt inclusions in granitoids of the Timma igneous complex, southern Israel. Min. Mag. 62, 29–40.

YANG, K. & SCOTT, S.D. (2002): Magmatic degassing of volatiles and ore metals into a hydrothermal system on the modern sea floor of the eastern Manus back-arc basin, western Pacific. Econ. Geol. 97, 1079–1100.

VARELA, M.E. (1994): Silicate–melt and fluid inclusions in rhyolitic dykes, Los Manantiales mining district, Argentina. Eur. J. Mineral. 6, 837–854. VEKSLER, I.V. (2006): Crystallized melt inclusions in gabbroic rocks. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 99-122.

25

SIMS IN THE DETERMINATION OF LIGHT STABLE ISOTOPES IN SILICATE MELT INCLUSIONS

CHAPTER 2: APPLICATION OF SECONDARY ION MASS SPECTROMETRY TO THE DETERMINATION OF TRADITIONAL AND NON-TRADITIONAL LIGHT STABLE ISOTOPES IN SILICATE MELT INCLUSIONS Graham D. Layne Dept. of Earth Sciences Memorial University of Newfoundland INCO Innovation Centre, Room 1047 St. John’s, NL, Canada A1B 3X5 E-mail: [email protected] accelerated into a solid sample at potentials of a few kV. The impact of these primary ions gradually erodes a shallow crater in the sample. All instruments routinely used for geochemical research use finely focused microbeams of primary ions and are thus generally referred to as “ion microprobes”. A portion of the material sputtered from the sample is ionized, and these secondary ions are the analyte species that are introduced into the mass spectrometer of the instrument, hence the name secondary ion mass spectrometry. When practitioners of SIMS refer to the “source” of the instrument, they are generally referring to the source of primary ions, for example, the duoplasmatron (DP) device commonly used to generate O– (or O2– or O2+). However, in concept, SIMS instruments are solid source mass spectrometers, with the sputtering crater comprising the “source” of ionized sample for mass spectrometry – analogous to the filament source in thermal ionization mass spectrometry (TIMS), or the plasma source in inductively coupled plasma mass spectrometry (ICP–MS). The mass spectra of secondary ions produced by sputtering are often complex. In addition to the monatomic singly charged ions most commonly used as analytes for light stable isotope determinations (e.g., 11B+/10B+ for δ11B, 37Cl–/35Cl– for δ37Cl), there are numerous possible isobaric interferences at each integral mass. For example, 10 BH+ interferes with 11B+ at 11 Da, and 34SH– interferes with 35Cl– at 35 Da. The appropriate level of mass resolving power (∆M/M; MRP) may be used to separate these interfering species with the magnetic prism of the instrument according to small differences in their mass. Alternatively, energy filtering (Shimizu et al. 1978) may sometimes be used to suppress interferences. This latter technique exploits differences in the initial energy spectrum of analyte and interfering species as they leave the

INTRODUCTION In the context of analyzing silicate melt inclusions, or similar materials, secondary ion mass spectrometry (SIMS) has the potential to achieve sub-per mil reproducibilities for a wide variety of light stable isotope ratio determinations. These can generally be accomplished with a lateral spatial resolution of better than 10µm where necessary, and with sputtered pit depths of less than a few µm. For glassy materials, this represents a total sample consumption of less than 10 ng for a single analysis. Sample preparation is normally fairly simple. Sample mounts must be compatible with the ultra-high vacuum of the sample chamber (ideally, better than 10–8 torr), and present a flat polished surface of the objects of interest. For analysis of insulating samples (such as silicate melt inclusions) mounts must generally be coated with a thin conductive layer of gold or carbon (300–500 Å). Most SIMS instruments used for light stable isotope determinations limit sample size to a maximum 25.4 mm diameter. Overall sample requirements are, therefore, quite similar to those for electron probe microanalysis (EPMA) of major or trace elements. SIMS shares the advantage with other forms of microanalysis such as EPMA, LA–ICP– MS and PIXE of allowing in situ analysis of low destructivity, preserving the textural context of the objects analyzed within the rock. This is, of course, invaluable when studying melt inclusions, where the extremely small size of an individual object severely limits study of their elemental or isotopic composition by other means. Abbreviations used repeatedly in this chapter are defined in Table 2-1. PRINCIPLES OF SIMS SIMS relies on the physical phenomenon of “sputtering”. A primary beam of ions is

Mineralogical Association of Canada Short Course 36, Montreal, Quebec, p. 27–49.

27

GRAHAM D. LAYNE

TABLE 2-1. ABBREVIATIONS CHAPTER TEXT. CCD cps Da

DP ECL EEF em EPMA eV fc IMF kV LA–ICP– MS MRP nA NEG pA PIXE SIMS TIMS

USED WITHIN

mass spectrometric techniques, the isotope ratio measured by the mass spectrometer is finitely different from the actual ratio in the sample. This effect increases with descending mass of the analyte ions and is of substantial magnitude. For example the measured ratio of 11B+/10B+ in many materials in the IMS 3f instrument is 4.8 % lighter than the actual sample ratio (Chaussidon et al. 1997). IMF is a consequence of multiple effects in the mass spectrometer, most of which are consequent to the relatively small differences in initial energy spectra of secondary ions of the various isotopes of the same element. Since these differences vary with the bulk chemical composition of the sample matrix, careful attention must be paid to characterizing appropriate reference materials for calibrating IMF. However, although IMF effects are of large magnitude in relationship to the (sub-) per mil differences to be measured in isotope ratios of scientific interest, IMF may be effectively held constant at these levels, allowing very reproducible results. It is commonly valuable to determine the concentration of the element itself, in tandem with its light stable isotope ratio, for example S with δ34S. This is often as simple as measuring a single additional “reference” peak related to the inclusion matrix, in this example, 30Si–. The concentration of S is then calculated simply by comparing 32S–/30Si– * wt.% SiO2glass in the sample to that in a reference glass of similar bulk composition. The term for Si concentration normalizes the denominator of the measured ratio. Determinations (e.g., δ34S, δ37Cl) that utilize Cs+ primary beams for the production of negative secondary ions from insulators like melt inclusion glasses also require additional charge compensation of the sample during analysis. Since this combination of primary and secondary ions produces an electron deficit at the sputtering site, an electron flood gun such as the Cameca NEG (see Instrumentation, below) may be effectively used for this purpose. One useful strategy is to use the minimum primary beam density necessary to produce sufficient signals for the analyte ions of interest, minimizing the sample charging effect, and reducing the dependence of IMF on charge compensation. This can be useful even in analysis of positive secondary ions (e.g., δ11B), where charge compensation depends exclusively on the conductive sample coating. Where primary currents of less than 10 nA are sufficient, O2– primary beams can sometimes be used to advantage for sputtering

charge-coupled cevice, as used as a sensor in digital cameras counts per second equivalent to the unified atomic mass unit (u), defined as 1/12 the mass of one atom of 12C. Preferred by many journals to the older term amu (atomic mass unit). duoplasmatron emitter-coupled logic extreme energy filtering electron multiplier detector electron probe microanalysis electron volt Faraday cup detector instrumental mass fractionation kilovolt laser ablation – inductively coupled plasma – mass spectrometry mass resolving power, 10% peak height definition used throughout text nanoampere normal incidence electron gun picoampere proton induced X-Ray excitation secondary ion mass spectrometry thermal ionization mass spectrometry

sample surface. In SIMS instruments that possess an energy-dispersed crossover in their mass spectrometer, the manipulation of an energy slit position, or of the initial accelerating potential of the secondary ions, allows preferential sampling of the monatomic species. However, many of the important interfering species in light stable isotope determinations are, in fact, hydride species like 10 BH+ and 34SH–, which do not have sufficiently different energy spectra from the monatomic analyte species to be suppressed effectively by energy filtering. As a consequence, most techniques for light stable isotope determination discussed herein involve the use of appropriate MRP to separate interfering species. A central concern in the quantitative determination of isotope ratios by SIMS is instrumental mass fractionation (IMF). As in many 28


positive secondary ions, even though they generally focus to somewhat larger diameters. This is because the singly charged oxygen dimer produces approximately double the sputtering rate for a given charging current. Further, although poorly documented in the published literature, it appears that collecting a broad band of secondary ion energies, through the use of wide energy slits (>50eV) in the mass spectrometer, may also be beneficial in minimizing the effect of sample charging on IMF drift in many applications. As light stable isotope measurements with SIMS become possible with precisions and reproducibilities approaching a few tenths of a per mil, as is now the case with δ18O using multicollection instruments, additional subtle physical effects of sputtering ionization may need to be incorporated in data reduction. For example, the phenomenon of quasi-simultaneous arrivals (QSA) of secondary ions at the conversion dynode of pulse counting ems – an effect related to the ejection of more than one analyte secondary ion in response to a single primary ion – may require a correction procedure at the sub-per mil level (Slodzian et al. 2001, Slodzian et al. 2004). The preparation of melt inclusions as polished mounts creates a thin (10–100 nm) layer of altered or damaged material on the surface of the sample. Exposure to air, cleaning solvents and the process of conductive coating, all add to this contamination burden. The presence of surface contaminants, if not addressed, can introduce strong biases into melt inclusion analyses. However, SIMS is intrinsically capable of eliminating surface contamination by pre-sputtering the site of analysis before data accumulation – essentially depth profiling through these layers. Most instruments also allow the insertion of a field aperture in the mass spectrometer (see Instrumentation, below), or other means to control the effective field of view for collecting secondary ions from the sample. This can be used to eliminate the collection of secondary ions from areas outside the melt inclusion boundaries or beyond the pre-sputtering preparation. It is, however, usually important to collect secondary ions from the entire diameter of the sputtered crater itself during analysis, avoiding irreproducibility in IMF caused by local spatial variations in the relative initial energies of the analyte ions. For this reason the total field of view is generally maintained somewhat larger than the area of the sputtered crater.

A second source of biased results can originate from contaminants that occur dispersed in three dimensions; mineral inclusions, microlites, fluid inclusions and contaminant-laden cracks or fissures. Fortunately, these too may often be detected, and their effect eliminated, by monitoring and discarding short-term excursions in the depth resolved data stream for analyte peak signals. INSTRUMENTATION Cameca IMS 3f/4f/5f/6f/7f Many important advances in the use of SIMS for light stable isotope determinations have been made using instruments from the Cameca IMS 3f/4f/5f series. A useful summary of the detailed features of these and other SIMS instruments is found in Ireland (1995). The Cameca f-series was first manufactured as the IMS 3f, beginning in 1978. A simplified synoptic of the original IMS 3f design is shown in Figure 2-1. Primary ions, usually O– or O2–, are produced in a duoplasmatron ion source (DP), and extracted into the primary column through potentials of up to 12.5 kV. A pair of Einzel-type electrostatic lenses, lens 1 (L1) and lens 3 (L3), is used to focus the primary ions to a small diameter spot on the sample surface. Simple fourplate deflectors are used to center the primary ion beam through these lenses. The L3 deflector plates can also be enabled to deliver a square-rastered primary beam to the sample, or to deflect the primary beam into a Faraday cup for measurement of the primary current. An eight-plate stigmator (L3 stigmator) is also associated with L3, to enable further shaping of the beam. L1 and L3 are used in tandem to deliver a de-magnified image of the exit aperture of the DP (source aperture) to the sample surface (critical illumination). An aperture placed before lens 3 (L3 aperture) is used as a beam limiter to control aberration of the focused spots. Subsequent f-series instruments added a Cs+ ion source, twinned with the DP on the primary column by means of a primary beam magnetic filter (PBMF). The PBMF selects explicitly between 16O– or 16O2–, or 133Cs+, depending on the desired mode of operation. They were also equipped with an additional electrostatic lens (L2) immediately after the PBMF, to improve the current density of the primary beam spot. This third lens, in tandem with an additional aperture (PBMF aperture), allows alternate strategies for primary beam focusing. These include using L2 to deliver maximum beam intensity to illuminate the PBMF aperture. An 29

GRAHAM D. LAYNE

Electrostatic Analyzer Energy Slit Duoplasmatr on

Deflectors Magnet

Field Aperture Contrast Aperture

Spectrometer Lens

Entrance Slit

Deflector 1 Exit Slit

Transfer Lenses L3 Stigmator Secondary Projector Beam L3 Lenses Deflectors

Lens 1 Deflector 2 L3 Aperture

Projection Deflector Stigmator Faraday Electron Cup Multiplier

Transfer Deflector Sample Lens 3 Airlock

Electrostatic Analyzer Immersion Lens

Detection Deflector

Channel Plate/ Phosphor Screen

FIG. 2-1. A simplified synoptic showing the major optical elements of the Cameca IMS 3f ion microprobe.

image of the PBMF Aperture is then focused on the sample by L3 in a manner analogous to that utilized on SHRIMP-type instruments (Kohler illumination, Ireland 1995). Secondary ions sputtered from the sample are extracted through a potential of ±4.5kV, and focused through the immersion lens into the transfer optics of the instruments. The transfer optics comprise three electrostatic lenses, used to control the maximum field of view for secondary ion collection from the sample, and the effective magnification for direct ion imaging. The transfer optics focus the secondary ions on the entrance slit of the mass spectrometer. The contrast aperture may be used to limit the cross section of the ion beam entering the mass spectrometer for high spatial resolution ion imaging, but is usually maintained at its maximum diameter for light stable isotope determinations. Cameca f-series instruments utilize a double focusing (Nier-type) mass spectrometer, which will accommodate either energy filtering or modest mass resolution approaches to the elimination of isobaric interferences. Energy filtering approaches, as discussed above, are

enabled by the energy-dispersed crossover at the energy slit located between the electrostatic analyzer (ESA) and the magnet of the mass spectrometer (Fig. 2-1). Alternatively, mass resolution can be increased from a base value of approximately MRP 300 by narrowing the entrance slit and exit slit. Secondary ions leaving the exit slit are delivered through the first projection lens, a smaller ESA, and the detection deflector to either a Faraday cup (fc) or electron multiplier (em) detector for quantitative ion detection. The isotopic ratios of different analyte masses are counted sequentially by cyclical switching of the magnetic field. The f-series design also allows for ion microscopy, either by direct ion imaging, or by scanning ion imaging using rastered primary beams registered to the pulse-counting em detector. For direct ion imaging, both projection lenses are used in tandem (and the detection ESA de-energized) to project a secondary ion image of the sample plane onto either a channel plate/phosphor screen/CCD camera combination, or a resistive anode encoder (RAE) detector. This imaging capability can be particularly useful for the very exact location and 30


centering of the analytical spot within small melt inclusions. For example, most melt inclusions within olivine hosts are significantly lower in Mg, and consequently register as a recognizable dark area in CCD imaging of 24Mg+. Cameca f-series instruments may also be equipped with the self-balancing normal incidence electron gun (NEG) designed by Slodzian for charge compensation when analyzing insulators (Migeon et al. 1990). This device uses a Wehnelt electron gun, and additional optics, to supply a cloud of electrons with an opposite balanced potential to that used to extract secondary ions from the sample. Any charging of the sample during analysis effectively changes this potential balance, drawing down a small compensating current of electrons to the sample surface. Direct CCD imaging for optimization of a bright spot signal (e.g., 16O–) from the sample is valuable for the routine alignment of compensating electron flux from the NEG for successive analytical spots. The IMS 6f and 7f are more recent models of the Cameca f-series. In addition to substantially increased automation and computerization of the instrument controls and data acquisition, they include several updates to the ion optics that benefit light stable isotope determination. Most significantly, they have an improved NEG design, with an Einzel lens for more uniform and stable focusing of the charge compensating electrons over a 150 µm diameter on the sample. The extraction potential for secondary ions has been increased from ±4.5 kV to ±10 kV, with incremental benefits for ion collection efficiency and em response. A laminated mass spectrometer magnet slightly reduces the waiting time for peak switching for analyses not involving the fc detector. The most recent IMS 7f instrument has an additional lens for the duoplasmatron source that substantially increases O– (or O2–) primary beam densities, advantageous for the analysis of positive secondary ions (e.g., Li+, B+). The IMS 7f also includes electrostatic deflection for rapid switching between fc and em detection for high dynamic range measurements. The automated controls for the ion optics allow easy alignment and programming of higher current coincident primary beams that can be used for fast and efficient pre-sputtering for the removal of surface contamination. Versions of the DP lensing system, improved NEG, and electrostatic detector switching are all available for retrofitting to the older f-series instruments.

In laboratories attempting sub-per mil reproducibility for light stable isotope analyses, it has been common to retrofit more sophisticated detector preamplifier and counting systems to all but the very newest instruments. Many different, customized approaches have been taken to these improvements but the objectives are in common: i) for the fc detector, to reduce and stabilize background noise (current) while maintaining a reasonably fast response rise time for peak switching analyses, ii) for the em detector, to maintain a stable pulse height distribution from the detector and stable overall dead time for the pulse counting system. Cameca IMS 1270/1280 The Cameca IMS 1270, first manufactured in 1993, is a large format high resolution, high transmission ion microprobe originally conceived and designed for U–Pb geochronology. It will accommodate a 5 device multicollector array with sufficient nominal mass dispersion to collect simultaneously 206Pb+, 207Pb+, 208Pb+, 232Th+ and 238 + U (Fig. 2-2). The primary ion column, sample chamber, immersion and transfer optics, and NEG are virtually identical to the IMS 6f. It is the much larger and more powerful mass spectrometer of the IMS 1270 that distinguishes these architectures. The mass spectrometer performance necessary for U–Pb geochronology has also translated into a highly versatile instrument with many features valuable for light stable isotope analysis. The multicollection array of the IMS 1270 is valuable in many light stable isotope determinations. Each of the five detector chariots can be used to deploy either an fc or miniaturized em detector. The ability to collect simultaneously two or more analyte peaks has obvious and readily calculable benefits for precision/time since the duty cycle for data collection on every peak can be virtually 100% (excluding counting on mass-free positions to establish background signals). It must be borne in mind, however, that considerable time must often be devoted to detector gain matching and calibration, sometimes severely diluting the precision/time efficiency gained over the monocollection peak switching approach. The fundamental strength of multicollection for many natural materials is in eliminating the cycle to cycle bias induced by primary ion current and sample-induced fluctuations in signal during peak switching analyses. This is of enormous value in increasing analytical precision 31

GRAHAM D. LAYNE

FIG. 2-2. Detailed synoptic of the Cameca IMS 1270/1280 multicollection array (adapted from Cameca, SA product catalog literature).

in samples where the analyte element is heterogeneously distributed within a glassy matrix. However, it is not always viable to use SIMS on very phase-heterogeneous inclusions. For some isotope systems, IMF is too highly dependent on bulk composition of the individual phases, introducing substantial errors. The IMS 1270 has a maximum mass dispersion (∆M/MMEAN) across the multicollection focal plane of 15.6%. This will nominally support simultaneous detection for any of the ratio measurements discussed herein. In some cases, combined analyses of more than one ratio (e.g., δ7Li and δ11B) might be accommodated using a single magnetic peak switch, and appropriate detector positions. However, several significant issues require attention for quantitative analysis with multicollection: i) periodic matching of detector gains is essential to maintain accuracy and reproducibility. Electron multipliers, in particular, may age quite rapidly, especially when incident ion currents exceed 105–106 cps. These effects are substantial, even on an intraday basis, when attempting to maintain sub per mil reproducibility while using miniaturized ems.

ii) the use of multicollection necessitates disengaging the first projection lens in the IMS 1270, disabling the capability for direct ion imaging of the sample. This reduces the capability for accurate spot centering, and eliminates the ability to objectively monitor NEG alignment on a spot to spot basis. The former capability will be partially restored, though in a less rapid and well resolved form, in future instruments, which will have scanning ion imaging capabilities tied to one or more multicollector ems. In addition to multicollection capability, the IMS 1270 offers fundamental advantages for some analyses in terms of superior transmissions at high MRP. The IMS 1270 maintains almost 100% transmission through the mass spectrometer for MRP as high as 3000 (10% definition), while the IMS 6f maintains 50%. This superior transmission offers no substantial advantage over the f-series instruments for major element analytes such as O (for δ18O) in silicates, where MRP of 2300 is sufficient to resolve the significant isobaric

32

GRAHAM D. LAYNE

Relative Transmission

10.00

1.00

IMS 6f

0.10

IMS 1270/1280 NanoSIMS 50

0.01 100

1000

10000

100000

Mass Resolution

FIG. 2-3. Relative transmission versus mass resolution curves for the Cameca IMS 1270 and IMS 6f instruments (adapted from Cameca, SA product catalog literature).

interferences, and signals are substantial even on the minor isotope (18O). However, for trace-element analytes such as B, it is a tremendous advantage for maintaining precise analyses even for the B concentrations of 10 ppm, or less, which are common in igneous melt inclusions. One other convenient feature of the IMS 1270 optics is a continuously variable square field aperture. This allows perfect matching of the desired field of view for secondary ion collection from the sample for a given spot size or raster presputtered area. The newest, IMS 1280 instruments were first delivered in 2005. They feature the same ion optics as the IMS 1270, supported with modernized “e7”-type electronics. Most of their “improved features” are a consequence of more sophisticated software for the automated control and centering of ion optics, and are gradually being made available for existing IMS 1270 installations through software upgrades. Foremost among these are automated precentering of secondary ions in the field aperture using deflectors in the transfer optics. Optionally, the entrance slit of the mass spectrometer may be mechanically re-centered for each analysis as well. The object is to maintain a uniform flight path (and energy cross section) of secondary ions through the transfer optics to the mass spectrometer from spot to spot on the sample. Field aperture re-centering, in particular, appears useful in refining reproducibility of light stable isotope analyses at the sub-per mil level. Other newly automated features include monitoring of pulse height distributions for em

detectors, conceived to enable correction of small changes in em gain during successive spot analyses. The simple permanent magnets traditionally used to correct flight path differences for the lightest ions (due to the Earth’s vertical magnetic field in high latitude installations) have been supplanted in some versions of the IMS 1280 by a series of Helmholtz coils mounted along the flight tubes of the mass spectrometer. These are supposed to provide tunable compensatory fields to address concerns about low order effects of stray magnetic fields on IMF at the sub-per mil level. SPECIFIC TECHNIQUES AND APPLICATION EXAMPLES δ18O: Stable isotope ratio analysis of a major element Quantitative determination of δ18O by SIMS has been in use for some time (e.g., McKeegan 1987, Hervig et al. 1992, Riciputi & Paterson 1994, Valley et al. 1998). The major challenge in achieving routine sub-per mil reproducibility has been to stabilize IMF, particularly for insulators (like most melt inclusions) and to calibrate for it accurately across a wide range of major element compositions. Also, the large dynamic range of this ratio (16O/18O ≈ 487) requires careful attention to detector characteristics and stability. Eiler et al. (1997), closely following Riciputi & Paterson (1994), summarized an approach that is typical of one that has been generally successful in obtaining reproducibilities better than ±1‰ (1σ) for δ18O in many laboratories 33

GRAHAM D. LAYNE

equipped with Cameca f-series instruments. They used an IMS4f, with a primary ion beam of 8 nA of Cs+, accelerated through a nominal potential of 10 kV and focused to a 20–30 µm spot on the sample. Extreme energy filtering (EEF, Schauer & Williams 1990, Hervig et al. 1992) was used to eliminate 17 OH– interference effectively from the 18O– peak, and to reduce sensitivity to variations in sample charge compensation with the NEG. The use of EEF is based on the observation that IMF becomes insensitive to variations in charge compensation, even of several tens of volts, for secondary ions with initial energies in excess of 350 eV as they leave the sample surface. In this case, EEF was performed using a sample offset potential of 350V (±25eV energy slit). Despite the high degree of signal loss to the energy filter, the substantial primary beam current used allows typical count rates on silicate matrices to be maintained at 106 cps on 16O– (≈2*104 cps on 18O–) or more, with detection of both peaks by a single regular format multiple-dynode em. Each spot analysis involved 120 cycles of counting by magnetic peak switching on 16O– (1s) and 18O– (5s) (or 200 cycles of 16O– (1s) and 18O– (10s)) with a 50 ms settling time between peak switches. Typical analyses therefore take between 726 and 2210 s (12–37 minutes), not including any pre-sputtering. Pre-sputtering times of a few minutes remain invaluable in removing surface contamination, and allowing steady state charging of the sample surface to be attained. Eiler et al. (1997) reported that actual inrun precisions (standard error of the mean for the 120 cycles) of ±0.92 ‰ (±0.60‰ for 200 cycle analyses) closely conform to the theoretical precisions based on Poisson counting statistics, which were typically ±0.90‰ (±0.50‰ for 200 cycle analyses). This confirms that instrumental instabilities and drifts during individual analyses were an order of magnitude lower than the counting precision limit. Riciputi & Paterson (1994), using a similarly equipped IMS 4f, had earlier demonstrated that the above method using 200 cycles of data collection would yield external reproducibilities routinely better than ±1‰ (1σ), typically ±0.7‰ (1σ), for quartz and carbonate materials. They also showed that the reproducibility for IMF (18O–/16O–) on their Brazilian quartz standard (–7.26 %) was consistent to ±2.1‰ (1σ) over a 15-day period. With periodic intraday calibration for IMF, this procedure can therefore yield overall reproducibilities that approach ±0.5 ‰ (1σ) for chemically

simple matrices. These, and other studies of δ18O and other light stable isotope ratios, have benefited from improved em preamplifier and counting systems, such as those manufactured by Pulse Counting Technologies, coupled with improved electron multipliers such as the ETP 133H. These ECLbased systems are indeed faster, maintaining overall system dead time as low as 9 ns. This attribute is beneficial in allowing quantification of count rates in excess of 106 cps, as is the improved linearity of the low resistance ETP 133H. However, the core benefit of such systems for high dynamic range measurements, such as δ18O, actually stems from improved stability of the pulse height distribution for individual ion strikes on the em, as processed by the em preamplifier and its contribution to improved reproducibility of the 18O/16O measurement. These systems can effectively limit the intraday effects of aging of the large format em to a few tenths of a per mil, even at count rates of 2*106 on 16O–. Gurenko et al. (2001) presented one of the first published studies of δ18O in phenocryst-bound melt inclusions, using the IMS 1270 at CRPG– Nancy. They used 1–10 nA primary Cs+ beams focused to 5–30 µm spot diameters. The NEG was employed for charge compensation, in an analogous manner to that used in earlier studies with f-series instruments. No energy filtering was used. The energy slit of the instrument was instead set to an extremely wide bandpass (120 eV; with energy slit center offset –25eV from the 0eV axis), and MRP 5000 was used to eliminate isobaric interferences on O analyte species. Monocollection mode measurements of δ18O used a combination of electron multiplier (18O–) and Keithley 642-supported Faraday cup (16O–) with gain intercalibration of these paired detectors monitored using 18O–/16O– ratios measured on standard glasses. Gain calibration was observed to drift by 1‰ per day, or more, an effect attributed to gradual aging of the electron multiplier. These measurements yielded precision and reproducibility similar those produced using f-series instruments with EEF, typically ±0.4‰ (1σ) and ±0.7‰ (1σ), respectively. Data for 34S–/32S– were collected during the same analysis as for 18O–/16O– (see discussion below, in δ34S, δ37Cl section). Additional δ18O determinations were also accomplished using paired Faraday cup multicollection, with ion intensities on 18O– maintained at 3–5 million cps. This enabled internal precision of better than ±0.5‰ (1σ) for only 3 34


of charge compensation by NEG for each spot, overall reproducibility for δ18O in flat, well-coated samples can be reduced to better than ±0.5‰ (1σ). A wide energy slit (>50eV) seems to provide a benefit in improving the reproducibility of IMF. This recent experience with IMS 1270/1280 instruments operated in monocollection mode suggests that the tradition of using EEF for δ18O on f-series instruments might bear careful reexamination. As discussed above, the primary column and transfer optics of the IMS 1270/1280 are virtually identical to those of the IMS 6f/7f. The major architectural changes from the IMS 4f instruments originally used in the studies of Riciputi & Paterson (1994) and Eiler et al. (1997) are i) the increase in secondary ion extraction potential from 4.5 kV to 10 kV, and ii) the improvement of the NEG to include an Einzel lens for electron focusing. It appears possible, therefore, that any f-series instrument equipped with the newer style of NEG, particularly 6f/7f instruments, should be capable of highly reproducible δ18O analysis using MRP to eliminate 17OH– interference. This would allow substantially lower primary beam currents, which would act in concert with the improved NEG charge compensation to increase the reproducibility of IMF, along with higher lateral spatial resolution (10–15 µm versus 20–30 µm). With multicollector IMS1270/1280 instruments, several detector pairings are possible for δ18O. However, the most commonly used are either fc for 16O– and em for 18O– , or fc for both peaks, utilizing an increased primary beam current. This latter arrangement allows precision of better than ±0.2‰ (1σ) to be developed for spot analyses as short as 5 minutes, while maintaining lateral spatial resolution of 10–15 µm. At this level of precision, overall reproducibility then becomes dominated by variations in IMF caused by sample behavior, small drifts or variations in the gain matching between the detector pairs, and matrix effects on IMF for chemically variable materials (such as natural glasses in melt inclusions). The first two effects are intrinsically combined within measurements of IMF on a standard material, which must be performed periodically to maintain intraday reproducibility. Given that several means exist to make SIMS determinations of δ18O at the sub-per mil level of precision and reproducibility, the lingering issue for the analysis of melt inclusions is that of accurate calibration for IMF for these chemically

minutes of data accumulation. Cross-calibration of Faraday cup gains showed intraday drifts limited to less than 0.1‰. Typical precision and reproducibility for δ18O were improved in this manner to ±0.2‰ and ±0.3‰ (1σ), respectively. In all cases, matrix-dependent changes in IMF of 18O–/16O– were calibrated using a set of eight natural glasses ranging in composition from 45.3 to 72.3 wt.% SiO2. Negative correlations of IMF with SiO2 and (Na+K)/Al atomic ratios, as well as positive correlations with FeO and CaO were also observed in this study. No statistical correlation of IMF with either Al2O3 or H2O (0.16 to 1.55 wt.%) was noted. In practice, a simple linear regression of observed IMF versus SiO2 for these materials was used to correct for IMF in the melt inclusion measurements. Overall uncertainty in δ18O, including that derived from the IMF correction scheme, was estimated at ±1.7 ‰ for monocollection and ±0.7 ‰ (1σ) for multicollection measurements. Approaches to δ18O determination with IMS 1270/1280 instruments have generally used mass resolution (MRP >2275) to resolve 17OH– from 18O–. This is sensible, since these instruments still retain virtually 100% transmission at this MRP. The consequently higher transmission achieved over EEF techniques allows the option of substantially reducing primary beam current, current density, and spot size., for example, 200 pA of Cs+ rastered over 15µm2 versus the 8 nA in 25 µm spots commonly used in the f-series analyses described above. Pre-sputtering of surface contamination becomes critical at these lower currents, given the drastically reduced rate of material removal. A 200pA primary Cs+ beam will produce the equivalent of >5*107 cps on 16O– (≈105 cps on 18 O) using the above conditions. Using a combination of fc (16O– (1s)) and em (18O– (4s)) detectors in monocollection mode with magnetic peak switching (and electrostatic detector switching) the IMS 1270 can develop a typical inrun precision of better than ±0.5‰ (1σ) for a fast 30 cycle analysis (3.5 minutes, including (1s) wait times for magnetic switching, and 1s background measurement cycle for fc noise); with external reproducibilities (σ/√n) of better than ±2‰ (1σ) for 10 spots and ±1.5‰ (1σ) per mil for 20 spots, using fully automated analysis of points within 15 mm of the center of a homogeneous sample such as SiO2. With longer counting times (12–15 minutes), and some manual attention to the centering and stability 35

GRAHAM D. LAYNE

variable materials. Hervig et al. (1992), Riciputi & Paterson (1994), Eiler et al. (1997) and Gurenko et al. (2001) have presented various semi-empirical models for the variation of IMF of 18O/16O in a variety of natural and synthetic minerals and glasses. None of these models, however, currently provide a basis to extrapolate IMF calibration adequately at a sub-per mil level across a wide range of melt inclusion chemistries. Limited data for natural glass compositions in Eiler et al. (1997) illustrate that IMF variations between basalt and rhyolite can exceed 4‰, and may vary by 1‰ even over restricted ranges of basalt compositions. If working within limited compositional ranges (e.g., high SiO2 rhyolite), it may suffice to calibrate against a single wellcharacterized standard of similar bulk composition. However, for studies involving systems with more variable melt inclusion chemistry, it will be necessary to develop and utilize a series of IMF standards that cover a wider compositional range.

plus supporting trace element evidence (determined separately by SIMS), they concluded that the inclusions represent 18O-enriched phonolite melts produced by low degree partial melting of a highly depleted peridotite that was metasomatized by slabderived fluids. On the basis of this interpretation, these inclusions provide the first direct demonstration of the existence of mantle fluids or melts with strongly elevated δ18O values. δ34S, δ37Cl: Stable isotope ratio analysis of minor elements with excellent sputtering ionization yields Unlike O, S and Cl exist only at minor to trace element levels in most glassy melt inclusions. However δ34S and δ37Cl are both readily amenable to SIMS determination, due to the excellent yields of S– and Cl– secondary ions under sputtering by Cs+ primary beams. Natural glasses routinely contain significant S, producing the isobaric interference of 34 SH– on 35Cl– in the secondary ion spectrum. Attempts to utilize SIMS for δ37Cl were initially limited, using f-series instruments, because hydride interferences like SH– are difficult to eliminate effectively using energy filtering. However, Layne et al. (2004) have recently published a detailed technique for determining δ37Cl in natural glasses using high MRP with the IMS 1270. Figure 2-4 displays the mass spectrum at 35 Da using an IMS 1270 at MRP 5250 (10% definition) for a sample of synthetic basalt glass with an extreme enrichment in Cl (2.79 wt.% Cl).

Applications: Using procedures similar to those of Riciputi & Paterson (1994) and Eiler et al. (1997) for the IMS 4f, Eiler et al. (1998) reported δ18O data for two vein-bounded silicate glass inclusions of phonolite composition, hosted by a sample of mantle olivine from the island arc setting of Simberi Island, New Guinea. These two inclusions showed δ18O of +11.3 ±1.3‰ (1σ, n=4) and +8.8 – 9.2‰ (n=2), respectively – values much heavier than the δ18O of +5.7±0.5‰ inferred for most mantlederived magmas. Based on these measurements,

FIG. 2-4. Secondary ion spectrum (em detection) at nominal mass 35 Da for a synthetic basalt glass (“Basalt glass B” of Godon et al. 2004, 2.79 wt.% Cl) to illustrate isobaric interferences at MRP 5250 (10% definition) (after Layne et al. 2004).

36

GRAHAM D. LAYNE

systems where individual samples span a broad range of melt compositions, a series of standards is required. Godon et al. (2004) have demonstrated that IMF is a highly correlated function of SiO2, Al2O3, CaO and FeO, permitting correction of the measured δ37Cl value to ±0.5 ‰ over this wide compositional range, using a simple linear fit based on a set of synthetic standard materials (Fig. 2-5). It is also notable that the IMF of 37Cl–/35Cl– for many SiO2 rich glasses is positive, as also noted for 7 + 6 + Li / Li (see δ7Li, δ11B, below). Analyses accumulated in 12 minutes routinely yield internal precision better than ±0.45‰ (1σ). Individual analyses of homogeneous materials have internal precision values (0.30– 0.45‰, 1σ) that approach the theoretical limits calculated from Poisson counting statistics (0.25– 0.30‰, 1σ). Overall reproducibility of individual spots is better than ±0.7‰ (1σ), even for sessions spanning several days, and can be maintained for 12 minute analyses of samples with as little as 250 ppm Cl. The multicollector of the IMS 1270/ IMS1280 will permit simultaneous collection of 35 – Cl and 37Cl– signals, with obvious advantages for reducing total analysis time and/or improving precision. However, any expected increase in overall counting time efficiency will be partially offset by the necessity of periodic intercalibration of the multiple em detectors. Further, the acquisition

The closest isobaric interference apparent is indeed that of 34SH– on 35Cl–. The average S/Cl ratios documented for many oceanic basalt examples are in fact 103–104 times larger than those for the Clenriched basalt of Figure 2-4, and so would produce a 34SH– peak of 10–100% the height of the 35Cl– peak. Total Cl contents of these same basalts are also commonly 120eV), and a resolution of 1800 MRP is applied to separate the interference of 10BH+ on 11B+ (Fig. 2-6). This level of MRP is also sufficient to eliminate any other possible interferences, such as 9 BeH+ on 10B+. Ion counting is by an em detector, using cyclical magnetic peak switching between 10 + B and 11B+. However, the Chaussidon et al. (1997) study was adapted in many ways to the ultra-low concentrations in their samples (100 ppb–1 ppm), as approached using an IMS 3f instrument. This included using primary beam currents as high as 100 nA to produce sufficient B+ signal, counting times of up to 90 minutes, and completely separate measurement of B concentration via 11B+/30Si+ using a classical energy filtering technique (±10 eV energy slit, –60 V sample offset potential). A useful revelation of the Chaussidon et al. (1997) study was the relative independence of B+ ionization efficiency from matrix effects over a

MRP 1800 11B+

105

I (cps)

104 10BH+

103 102 101 100 10-1

10.95

10.975

LC 19 (4.5 wt%)

log (Bppm /SiO2wt% )

1

JV 1 (1920 ppm)

RM 6 (252 ppm)

0 DK 89 (14 ppm)

-1

UTR 2 (17.7 ppm)

BCR 1 (3.1 ppm) NBS 614 (1.3 ppm)

-2

NBS 617 (0.2 ppm)

-3

-5

-4

-3

-2 -1 log (11B+/ 30Si+)

11.05

wide range of natural and synthetic glass compositions including natural basalt and pantellerite, and synthetic compositions that approximated high-silica rhyolite. In practice this allows a single universal working line for determinations of B concentration (Fig. 2-7). They also demonstrated that the IMF for 11B+/10B+ is extremely consistent over a very wide range of matrix compositions (silicate glasses, marine salts and even boric acid),

LQ 24 (1.0 wt%) GB4 (970 ppm)

11.025

FIG. 2-6. Secondary ion spectrum at nominal mass 11 Da for the GB4 glass reference material showing separation of the 10BH+ on 11B+ peaks at MRP 1800 (10% definition) (after Chaussidon et al. 1997). Originally published in Geostandards Newsletter: The Journal of Geostandards and Geoanalysis (1997, Vol. 21 No. 1, pp. 7-17), reproduced by permission of the copyright holder - Association Scientifique pour la Geologie et ses Applications (ASGA).

3 2

11

Da

0

39

1

FIG. 2-7. Calibration line for B concentration measurements – B/SiO2 versus 11B+/30Si+ (data from Chaussidon & Libourel 1993). The linear fit obtained for standards of diverse composition shows that the background for B is low (120 eV). Magnetically switched peaks were collected for 120 cycles of 6Li+ (6 s) and 7Li+ (3 s), with the magnetic field position for the peaks scanned and re-centered every 10 cycles. A wide range of standard compositions plotted on a single IMF calibration line (Fig. 2-9). Consequently, it was possible to use a single sample of fused natural basalt (BHVO, 5 ppm Li) to calibrate IMF for glass and mafic mineral samples. Chaussidon et al. (1997) and subsequent studies, including Decitre et al. (2002), have noted a seemingly unusual feature of the IMF for 7Li+/6Li+ – it is positive (+4% to +6%), meaning that the heavier ion (7Li+), is enriched in the measured ratio. The reproducibility of δ7Li for replicate BHVO analyses was better than ±1‰ (1σ) during single analytical sessions. Decitre et al. (2002) determined Li concentrations separately using energy filtering, counting 7Li+ and 30Si+ at an MRP of 500, with an energy slit of ±10 eV and a sample offset potential of 80 V. Olivine, pyroxenes, biotite, amphibole, and basalt plot on a single working line for Li concentration (Fig. 2-10), demonstrating the absence of large matrix effects on Li ion yield for 41

GRAHAM D. LAYNE

these bulk compositions. The serpentine standard UB-N scatters significantly from this line, due to an inherently inhomogeneous content of Li. Gurenko & Schmincke (2002) used an IMS 1270 to emulate the same conditions established by Chaussidon & Robert (1998) in their earlier studies of δ11B and δ7Li using the IMS 3f. Their precision and reproducibility for these two quantities in replicate analyses of the GB-4 standard glass were consequently similar to this earlier study. Since the natural glasses examined in this newer study contained less than 1 ppm of total B, the precision of individual analyses of δ11B was limited to ±2.2 to ±3.8‰ (1σ). Kobayashi et al. (2004) performed both δ11B and δ7Li determinations using an IMS 1270 in multicollection mode, for their study of olivinehosted melt inclusions from Hawaiian lavas. Primary ion beams were 15 nA O– for Li, and 20 nA O– for B, producing final pit diameters of ~20, and ~30 µm, respectively. MRP was ~2000 for both determinations, near the intrinsic minimum MRP of the IMS 1270. Secondary ions were extracted through the standard 10 kV potential. A 50 eV energy window was utilized, with no sample offset

FIG. 2-9. Accepted value of δ7Li by TIMS (Bristol University) versus δ7Li measured by SIMS (IMS 3f). High correlation of both mafic mineral and basalt glass standards (BHVO) illustrates a negligible difference in matrix effect on IMF for these compositions (after Decitre et al. 2002). Copyright 2002 American Geophysical Union, reproduced in modified form by permission of American Geophysical Union.

FIG. 2-10. Ion yield calibration lines for mineral and basalt standards (four different analytical sessions): Li/Si (as normalized from 7Li+/30Si+) measured by IMS 3f SIMS versus an accepted Li/Si measured conventionally on bulk powders. Standards include biotite, clinopyroxene MC and olivine MC (Massif Central, France), clinopyroxene BZ29 and olivine BZ29 (Zabargad lherzolite), amphibole (Kipawa) and basaltic glass (Nazca). Serpentine UB-N (Vosges) appears intrinsically inhomogeneous for Li/Si. This line illustrates the independence of Li+ ion yield under sputtering from matrix effects for these minerals and glasses (after Decitre et al. 2002). Copyright 2002 American Geophysical Union, reproduced in modified form by permission of American Geophysical Union.

42

GRAHAM D. LAYNE

potential. All peaks were counted with miniaturized ems positioned on the multicollector array. For both δ7Li and δ11B determinations a single analysis consisted of 550 s counting after 5 min of presputtering. Typical sensitivities were ~2000 cps/ ppm/nA for 7Li+ and ~50 cps/ppm/nA for 11B+. Concentrations of Li and B in the glass inclusions were subsequently determined independently, using a Cameca IMS 5f (after Nakano & Nakamura 2001). The reference materials for IMF were a series of five synthetic glasses, all fused from a homogeneous powder of a basaltic andesite from Izu-Oshima, Japan, but spiked with increasing concentrations of Li and B. IMF and differences in em gain were calibrated using linear regressions of these five glass standards as measured both before and after analysis of unknown samples. Typical precisions for individual determinations were ±0.6‰ (1σ) for δ7Li (for ~25000 cps on 7Li+, equivalent to 0.83 ppm Li) and ±0.5‰ (1σ) for δ11B (for ~23000 cps on 11B+; equivalent to 23 ppm B). Reproducibility, based on repeat measurements of the set of reference glasses was better than ±1.0‰ (1σ) for both δ7Li and δ11B. The uncertainty related to the accuracy of the IMF calibrations was estimated to be ±0.4–0.6‰ (1σ) for δ7Li, and ±0.3–0.8‰ (1σ) for δ11B. Kasemann et al. (2005) assessed the δ7Li of available basalt glass reference materials using several techniques, including SIMS. They determined δ7Li with an IMS 4f, using conditions very similar to those of Decitre et al. (2002), with the exception that the energy slit was narrowed to 50 eV. They concluded that USGS glasses GSD-1G (δ7Li 31.1±0.4‰, 1σ) and BCR-2G (δ7Li 4.1±0.5‰, 1σ) are useful standards for the calibration of IMF in δ7Li determinations of basaltic glass. However, they also noted that SIMS analyses of the NIST 610-612-614 series of microanalytical glass standards (~72 wt.% SiO2) imply a substantially different magnitude of IMF in high silica glasses. For SIMS determinations of δ7Li (IMS 4f) in USGS reference glass GSD-1G (37 ppm Li), they achieved ±0.4‰ (1σ) precisions for individual spots (~960 s analyses), external reproducibilities (σ/√n) of ±0.6‰ (1σ, n =10) for single day sessions, and ±0.2‰ (1σ) for four days. In subsequent trials of GSD-1G, using an IMS 1270 in monocollection mode, they achieved precisions of ±0.2‰ (1σ) for individual spots (~800 s analyses), external reproducibilities (σ/√n) of ±0.5‰ (1σ) for single day

sessions, and of ±0.4‰ (1σ) over two days. Applications. Gurenko & Chaussidon (1997) successfully used the analytical approach of Chaussidon et al. (1997) to establish a uniform δ11B for the Icelandic mantle (–11.3 ±1.9‰), based on their analysis of olivine-hosted melt inclusions of primary mantle melt. Gurenko & Schmincke (2002) studied the late Pliocene (~2 Ma) orthopyroxene-bearing tholeiite of the Iblean Plateau, which is believed to represent a large volume, high degree (16–30%) melting event of the mantle beneath Sicily. They determined δ7Li and δ11B in olivine- and orthopyroxene-hosted glass inclusions and their host pillow-rim glasses. Major element variations, H2O (0.2–0.5 wt.%), Cl (100–350 ppm), as determined by EPMA, and Li (4.6–5.8 ppm), Be (0.5–0.8 ppm) and B (0.6–1.1 ppm), as determined by SIMS, were deemed to reflect heterogeneity of the Iblean magma source. The δ7Li (–3.4 to +1.2‰) and lightest δ11B (–17.1‰ to –12.9‰) values observed in the glass inclusions (n=8) and their host glasses (n=3) are lighter than the average MORB, but very similar to OIB magmas. This is attributed to the presence of isotopically light mantle domains beneath the Iblean Plateau, most probably resulting from previous subduction of crust. Further, the wider variations of δ11B (as heavy as –3.1‰) over restricted ranges in δ7Li in a small subset (n=3) of the measured samples are explained by the addition of less than 2 wt.% of altered basaltic rocks to the crystallizing magma in shallow crustal reservoirs. Schmitt et al. (2002) determined δ11B and B in quartz phenocryst-hosted primary melt inclusions (n=31) from the calc-alkaline ignimbrites and lavas of the Neogene-Pleistocene central Andean Altiplano-Puna Volcanic Complex. Strongly devitrified inclusions were rehomogenized in a hydrothermal bomb apparatus for 20 hours at 800°C and 100 MPa with a CO2 pressurizing medium before SIMS analysis. The average δ11B of both melt inclusion and matrix glasses from individual ignimbrite units are relatively uniform with an overall average δ11B of 3.8 ±2.8‰ (1σ). This range overlaps that of local basement rocks (δ11B, –5 to –11‰), implying a dominantly crustal source for these magmas. Straub & Layne (2002) determined δ11B and B concentrations in matrix glasses (n=24) and glassy plagioclase-hosted melt inclusions (n=12) from Neogene fallout tephra of the Izu arc volcanic 43

GRAHAM D. LAYNE

δD, δ13C: Stable isotope ratio analysis of volatile trace element analytes with strong exotic contribution issues The primary challenge in both δD and δ13C determinations by SIMS is controlling the contribution of exotic or contaminant H and C to the secondary ion signals. In the case of δD, the extremely high dynamic range of the measured ratio (H/D >6000) provides an additional degree of difficulty. Enhanced sample chamber pumping with cryopumps or N2 cold fingers can help to reduce the partial pressure of H and C-based gas species near the sample. However, most exotic contributions to H and C signal appear to originate from the sample surface itself, with localized migration to the site of sputtering during analysis. Special mounting procedures, such as replacing epoxide mounting media with indium metal, may have some benefit in reducing surface contamination (Hauri et al. 2002). Pre-preparation of samples in vacuum ovens, and/or high vacuum chambers for periods of up to days before analysis, has a marked impact on contaminant signals. However, careful attention to pre-sputtering of the sample, and the exact conditions of secondary ion collection, are crucial in performing accurate analyses. The first detailed published application of SIMS to the determination of δD in terrestrial materials was that of Deloule et al. (1991) who explored the δD analysis of hydrous minerals. They utilized a primary beam of 2–5 nA of O– rastered over an approximately 30 µm2 area. The dynamic transfer optical system (DTOS) of the IMS 3f was used to maintain MRP, even with this larger rastered beam spot. Mass resolution was set at MRP 1300 to ensure separation of H2+ interference from D+ (Fig. 2-11). The energy slit was maintained wide open (>120 eV) and no sample offset was applied. Ion detection was accomplished with a regular format em in pulse counting mode. Peak switching analyses were run for 90 to 120 minutes, to achieve precision of approximately ±5‰ (1σ). To mitigate against contaminant H, both samples and the IMS 3f sample chamber were baked at 120ºC. A liquid nitrogen cold finger was also active within the sample chamber during analysis. The ratio of H2+/H+ appears diagnostic of moisture contamination of the sample, and thus analyses were only carried out on samples that had reached a condition where the measured H2+/H+ was 50% of the B in the Izu VF rocks was derived from an underlying mantle wedge. Straub & Layne (2002) thus concluded that down-dragged serpentinized mantle wedge is an important reservoir in the eventual recycling of B through arc front volcanism and it has an important effect on the long term recycling of crustal constituents in the subduction factory. Kobayashi et al. (2004) determined δ7Li 11 and δ B (and Pb isotopes) in olivine-hosted melt inclusions (n=28) from lava samples of the Hawaiian volcanoes Kilauea Iki, Mauna Loa, and Koolau. Host olivine grains were heated in Pt crucibles (1 atm, QFM buffer) for 30 minutes at 1190ºC to rehomogenize all melt inclusions to glass prior to analysis. Overall, these melt inclusions showed substantial variations in both δ7Li (–10.2 to +8.4‰) and δ11B (–10.5 to +5.2‰), markedly exceeding those previously measured in whole-rock samples from these same volcanoes. The authors interpreted the lowest observed values for δ7Li (minimum –10.2‰ in Koolau) and δ11B (minimum –10.5‰ in Mauna Loa) as indicative of an isotopically light Li and B source. They considered this source most likely involved the recycling of crustal materials that had experienced near-surface alteration and then dehydration during subduction, and that it played a potentially important role in creating the geochemical and isotopic heterogeneity observed in Hawaiian lavas. 44


analytical approach to δD and δ13C may be the minimization of H– desorbed directly from the sample surface due to the impact of “stray” or continuum flux electrons consequent to small imperfections in the potential matching of the electron flood. Once good vacuum was established in the sample chamber (106 cps. This required primary beam currents as high as 10– 15 nA for 0.1 wt.% H2O. IMF drift for D/H was usually limited to 5‰ per day. For samples with >0.5 wt.% H2O, 2–3‰ (1σ) precision was generally achieved for a 20–40 min analysis, and reproduceibility for replicate analyses of the same sample approached that governed by the precision of the individual spots. IMF calibration relied on a series of basalt, andesite and rhyolite standards with known δD. Pending a more detailed presentation of matrix effects on IMF, Hauri et al. (2002)

103

MRP 1250

H2+ 2

I(cps)

10

101

D+

100

10-1 2.014

Da

2.016

FIG. 2-11. Spectrum at nominal mass of 2 Da showing H2+ and D peaks resolved at MRP 1250. Primary beam was 5nA of O– focused to a 15µm spot. (after Deloule et al. 1991).

Deloule et al. (1991) used their results to undertake a detailed examination of the variation of IMF with major element chemistry in amphiboles and micas. The most detailed published procedures for δD and δ13C by SIMS in natural glasses are those of Hauri et al. (2002) and Hauri (2002) using an IMS 6f. In contrast to most previous studies of δD and H2O concentrations in glassy samples, Hauri et al. (2002) used negative secondary ions. This approach is inherently well suited to δ13C determination, since C ionizes quite efficiently as C– under Cs+ primary ion bombardment. The use of H– as the analyte for δD is a strategy that appears beneficial in reducing exotic contamination. The H+ analyte used in previous studies (e.g., Deloule et al. 1991) seems inherently prone to contamination effects, since 1H+, a simple proton, is expected to be omnipresent in sample surfaces. Cs+ primary ion beams (10 kV) were collimated using Kohler illumination (see Instrumentation, above) of the final primary beam aperture, to produce spots of 40 µm diameter by slightly defocusing the beam on the sample. Charge compensation with the NEG was employed, using procedures essentially similar to those commonly employed for δ18O determinations with f-series instruments. The total electron current generated by the NEG was reduced to the minimum value effective for charge compensation. Hauri et al. (2002) presented this as a precaution against H– desorption from or within the transfer/electron gun optics. However, the larger benefit for many types of

45

GRAHAM D. LAYNE

demonstrated that IMF for D/H is –20 to – 26% and is correlated with the H2O and Fe contents of their reference glasses. IMF for rhyolite was often 2–4% greater than for basalt and andesite. Overall uncertainty for δD analyses was cited as ±5‰ (1σ, Hauri 2002). Hauri et al. (2002) also performed δ13C analyses with the IMS 6f using similar conditions to those for δD, except that mass resolution was increased to MRP 3200 to eliminate the interference of 12CH– on 13C–. IMF calibration was based on andesite and basalt standards. Assorted basalt glass standards gave the same IMF to ±1‰, on the order of –3.3 to –3.9% for D/H. IMF was lower in andesite glasses (–2.8 to –3.2%). The precision of individual (60 min) δ13C analyses was ±2–3‰ (1σ), limited largely by counting statistics. Overall uncertainty for δ13C analyses was stated as ±5‰ (1σ, Hauri 2002).

H2O values back along open system degassing curves using estimates of the primary H2O content of each inclusion. The author concluded that there are significant contrasts in the δD of the primary magmas for each of the Hawaiian volcanoes in this study, ranging from approximately –48‰ for Kilauea to –120‰ for Koolau. This implies a δD heterogeneity in the underlying Hawaiian mantle plume that is spatially extensive (> tens of km) and/or relatively young (200 ppm CO2) were also determined for δ13C, displaying light values (–12 to –29‰). The author interpreted the broad correlation of δ13C with δD in these data as consistent with open-system degassing of CO2-rich magmas followed by mixing with less degassed magmas.

Applications: Using the techniques described in Hauri et al. (2002), Hauri (2002) used an IMS 6f instrument to measured δD as well as H2O, CO2, F, S and Cl in olivine-hosted melt inclusions from five lava samples of the Hawaiian volcanoes Loihi, Kilauea, Mauna Loa and Koolau. All melt inclusions were first briefly reheated to 1275ºC in a 1 atm furnace (QFM buffered) to rehomogenize quench crystals in the glass. Heating time was limited to 10 minutes to minimize diffusive loss of hydrogen to the host olivine. After SIMS analysis, several chemical criteria were applied to filter out inclusions containing melts that might have suffered modification of their primitive δD compositions through natural processes of H diffusion, or assimilation, which preceded or accompanied lava emplacement. These criteria included discarding data for: i) inclusions suspected of suffering significant diffusive loss of H to the olivine host (all samples with positive δD values), ii) inclusions that may have experienced assimilation of seawater components as evidenced by apparent Cl contamination (all samples with Cl/K ratios in excess of those considered “normal” for each specific volcano), iii) inclusions that may have experienced shallow degassing (all samples with CO2 750

>800

Thom, oC

ap

ol

mon

carbonatite

Host mineral

carbonatite phoscorite

Rock type

Palabora, RSA

Magnet Cove, USA

Pluton

TABLE 6-1. MODAL COMPOSITION OF CRYSTALLIZED CARBONATITIC MELT AND FLUID INCLUSIONS (VOL.%). MAGMATIC PROCESSES AND VOLATILE PHASE GENERATION IN PORPHYRY-TYPE ENVIRONMENTS

ILYA V. VEKSLER & DAVID LENTZ

TABLE 6-2. MANTLE-DERIVED CARBONATED SILICATE LIQUIDS REPRESENTED BY MELT INCLUSIONS AND KIMBERLITIC GROUNDMASS (WT.%). Pluton

Kugda, Russia

Krestovskiy, Guli, Russia Russia meliliteolivineRock olivinite melilitolite monticellite melilite rock rocks Host perovskite perovskite melilite perovskite olivine Quench T, °C 1100 1050 1050 1170 1221 35.66 SiO2 34.78 36.64 34.50 35.06 Ti°O2 5.53 3.85 4.99 6.40 5.43 Al2O3 4.07 8.44 8.76 7.70 11.23 FeO 7.9 8.61 9.07 9.40 8.17 MgO 10.2 4.75 5.02 6.10 4.66 CaO 22.46 17.05 18.99 15.10 21.0 Na2O 4.24 4.93 4.87 4.20 8.48 K2O 3.18 2.24 1.56 4.90 3.14 P2O5 0.53 2.96 2.36 2.40 0.78 SrO n.d. 0.65 0.31 0.40 n.d. BaO n.d. 0.36 0.66 0.40 n.d. SO3 n.d. n.d. n.d. 0.60 n.d. Cl n.d. n.d. n.d. 0.20 n.d. CO2 n.d. n.d. n.d. n.d. n.d. Total 92.89 90.50 90.65 92.30 98.2a Ref.

Gardiner, East Greenland

Nielsen et al. 1997

Panina 2005

Rass & Plechov 2000

Zeughaus dike, Germany

Lac de Gras, Canada lherzolitic olivine melilitite xenoliths in kimberlite olivine ap, cpx, ol Cr-diopside 982 855 unheated 34.52 21.57 30.0 4.22 4.36 0.26 13.0 13.27 1.3 10.35 10.83 6.0 4.9 0.8 22.0 17.83 37.68 16.0 5.61 0.16 0.05 3.15 0.04 0.2 2.72 0.07 0.03 n.d. n.d. 0.03 n.d. n.d. 0.04 0.8 1.10 0.9 0.53 0.05 0.01 2.0b 9.8b 13 c c 97.96 90.18 89.82 Seifert & Thomas 1995

van Achterbergh et al. 2004

UdachnayaEast, Russia kimberlite groundmass unheated 26.10 1.30 2.00 7.20 28.90 12.40 4.30 2.10 0.40 0.11 0.13 0.70 2.80 10.3 98.74 Kamenetsky et al. 2004

a – homogenized glassy inclusions contain 1.4 wt.% Cr2O3; b – volatile content estimations based on low totals of electron microprobe analyses; c – excluding volatiles. Abbreviations same as in Table 6-1.

TABLE 6-3. COMPOSITIONS OF CARBONATITIC MELT INCLUSIONS AND CHLORIDE-CARBONATE GLOBULES IN THE KIMBERLITIC GROUNDMASS. ABBREVIATIONS FOR MINERALS ARE THE SAME AS IN TABLE 6-1. Pluton

Magnet Palabora, Cove,USA RSA Rock carbonphos- carbonatite corite atite Host mon ol ap 15.7 11.6-17.4 7.28 SiO2 TiO2 n. d. n. d. 0.07 Al2O3 n. d. n. d. 1.11 FeO 4.38 8.2-13 0.74 MnO n. d. 0.1 0.09 MgO 0.98 13.9-19.8 3.19 CaO 49.67 17.5-23.3 48.77 Na2O n. d. n. d. 0.06 K2O n. d. 2.9-4.3 0.77 P2O5 1.13 n. d. 11.46 SrO n. d. n. d. 0.36 BaO n. d. n. d. 0.1 SO3 n. d. 0.02-0.7 n. d. F n. d. 2.3-3.4 3.41 Cl n. d. n. d. n. d. CO2b 16.71 20.9-27.9 22.59 Total 88.57 100.0 Ref.

Nesbitt & Kelly 1977

Lac de Gras, Kovdor, Russia Gardiner, E. Krestovskiy, Russia Udachnaya, Canada Greenland Russia kimberlite peridotite phoscorite melilitolite melilitolite, pyroxenite kimberlite Cr-diopside 20.4 16.4 0.24 0.22 0.85 0.67 4.44 2.3 0.14 0.14 22.6 11.3 26.4 37.2 0.03 0.12 0.72 0.48 0.1 0.13 0.86 1.18 1.35 1.8 n. d. n. d. 0.02 0.01 n. d. n. d. 21.6 27.9 99.75 99.85

ol 10.65 0.03 0.34 3.50 0.06 6.86 36.25 5.74 1.42 2.0 n. d. n. d. n. d. n. d. n. d. 33.56 100.43

ol 7.87 0 0 3.57 0.51 16.09 18.5 16.2 1.22 4.18 n. d. n. d. n. d. n. d. 0.4 30.38 99.10

mel 0.26 0.11 0.04 0.44 0.01 0.84 44.59 8.73 1.30 2.01 0.34 0.01 n. d. n. d. n. d. 39.83 98.49

Aldous Solovova van Achterbergh Nielsen & Veksler Nielsen et al. 1980 et al. 1998 et al. 2004 2002, and unpublished 1997

mel groundmass pvta mon 6.97 1.41 1.88 1.32 2.09 0.0 0.0 0.05 0.96 0.0 0.13 0.09 1.80 0.56 8.1 0.37 n. d. n. d. n. d. n. d. 2.30 1.68 0.45 2.47 19.01 32.84 35.66 31.9 5.68 13.09 5.42 18.87 10.59 7.8 10.2 2.41 2.98 5.35 2.73 0.23 0.56 3.5 0.9 0.25 0.90 0.35 0 0.05 7.89 5.32 0.63 1.45 n. d. n. d. n. d. n. d. 1.08 0.06 4.59 4.32 n. d. n. d. n. d. 35.27 62.82 71.96 70.69 99.05 Panina 2005

Kamenetsky et al. 2006

a – average composition of immiscible globules in silicate-carbonate inclusions; b – CO2 estimations are based on modal proportions of carbonate daughter minerals. Abbreviations same as in Table 6-1.

130

PARENTAL MAGMAS OF PLUTONIC CARBONATITES

TABLE 6-4. THE EXAMPLES OF ALKALI-RICH CARBONATITIC FLUID INCLUSIONS AND IMMISCIBLE SALT MELTS ASSOCIATED WITH CARBONATITES; COMPOSITIONS IN WT.%. Pluton

Kovdor, Russia

Gardiner, East Greenland

Rock

carbonatite

melilitolite

Kenya ijolite

urtite

Krestovskiy, Russia

Mushugai-Khuduk, Mongolia

melilitolite, pyroxenite

celestite-fluorite rock

Host Thom., C SiO2 TiO2 Al2O3 FeO MnO MgO CaO Na2O K2O P2O5 SrO BaO SO3 F Cl CO2a H2O

apatite melilite apatite apatite perovskite fluorite 620-870 970 n. d. n. d. 1030 670 600 0.0 0.73 2.0 3.7 0.26 0.0 0.0 0.0 0.04 0.0 0.0 0.96 0.0 0.0 n. d. 0.10 1.0 1.8 0 0.0 0.0 n. d. 0.95 0.9 0.0 0.06 0.0 0.0 n. d. 0.03 n. d. n. d. n. d. 0.0 0.0 6 - 7.2 0.68 0.0 0.0 0.08 0.0 0.0 20 - 30 19.19 19.1 25.0 2.95 35.0 26.0 12.5 - 14 28.83 26.3 20.9 10.78 2.7 11.1 2.3 - 5.2 0.48 6.1 5.9 19.99 3.9 3.7 4.5 - 9.5 0.84 n. d. n. d. 0.42 0.0 0.0 n. d. 1.46 0.5 0.0 0.19 9.9 5.7 n. d. 12.05 4.4 3.9 0.11 1.2 1.9 n. d. n. d. 4.4 1.8 20.07 33.1 29.0 n. d. n. d. n. d. n. d. n. d. 5.5 2.1 n. d. n. d. 6.9 2.6 0.16 0.1 10.9 n. d. 39 29.2 34.9 n. d. 2.1 4.5 n. d. n. d. n. d. n. d. n. d. 6.0 4.7 Veksler et al. Nielsen et al. Le Bas & Aspden Ref. Panina 2005 Andreeva et al. 1998 1998a 1997 1981 a , CO2 estimations are based on modal proportions of carbonate daughter minerals.

a

mt

Kalkfeld, Namibia quartzite at contact with carbonatite quartz n. d. n. d. 1-1.4 n. d. 3-4.1 0.6 1 - 1.7 13 - 16 18 - 21 7 - 10.8 n. d. up to 3.2 up to 0.9 n. d. n. d. n. d. ~20 ~20 Bühn & Rankin 1999

b

mel cpx

sht mt

cpx

cpx

50 mm

40 mm ol

c

d

ol pvt pvt

ol

phl ol 50 mm

50 mm

FIG. 6-2. Photomicrographs of crystallized melt inclusions from carbonatites and associated silicate rocks. a. Crystallized melt inclusion in forsteritic olivine, Kovdor peridotite. b. Melilite-hosted inclusion from Kovdor olivine-melilite rock. c. Large silicate-carbonate inclusion in forsteritic olivine from Kodor phoscorite surrounded by swarms of smaller inclusions. d. Small crystallized melt inclusion in perovskite (arrow), phoscorite, Vuoriyarvi carbonatite intrusion, Kola peninsula, Russia. See text for more details. Abbreviations for phases: cpx – clinopyroxene, mel – melilite, mt – magnetite, ol – olivine, pvt – perovskite, sht – shortite.

131


solid and fluid inclusions in monticellite, apatite, and calcite with the employment of scanning electron microscopy (SEM) and energy-dispersive X-ray analysis (EDX). Inclusions in apatite, similar to those studied at the volcanic and sub-volcanic localities in East Africa (see above), were extremely abundant, but were difficult to characterize because of the extreme variation in phase proportions (solids, liquid and vapor). The variations were ascribed to necking-down and/or heterogeneous entrapment of immiscible fluids. In contrast, primary and secondary inclusions in early-formed monticellite exhibited a greater uniformity in their modal composition with over 50 vol.% solids (Table 6-1), water and a vapor bubble (presumably CO2-rich). Attempts to homogenize the inclusions were unsuccessful, even after heating them to 800oC for a week. Bulk compositions of the inclusion phase assemblages were calculated on the basis of the EDX analyses of solid phases and statistically representative modal proportions of the daughter minerals (Table 6-3). Calcite was the predominant daughter phase among those identified, and the inclusions were interpreted as trapped droplets of carbonatite magma with an estimated density of 2.2 to 2.3 g.cm–3. The enrichment of silica and Fe oxides in the melt relative to the calciocarbonatite, was attributed to crystal settling and relative enrichment of calcite at shallower levels. The calcite from Magnet Cove was found to contain abundant secondary and pseudo-secondary fluid inclusions, but none appeared to be unequivocally primary. Palabora, Republic of South Africa The 2012–2047 Ma Palabora complex of Limpopo Province comprises a large oval-shaped body of pyroxenite, syenite and carbonatite intruded into Archean gneiss (Eriksson 1989, Woolley 2001). The central carbonatite zone of the complex is famous for extensive economic Cu-sulfide mineralization. Three different carbonatites, emplaced in the following order, have been recognized (Aldous 1980, see also Rankin 2005): 1) phoscorite (olivine-apatite-magnetite rock) 2) banded apatite–magnetite carbonatite 3) cross-cutting, transgressive, apatite-magnetite carbonatite. The late-stage, transgressive carbonatite is the dominant host for the primary Cu-sulfide mineralization, where it occurs in a series of discontinuous veinlets.

FIG. 6-3. Back-scattered electron images of carbonatesilicate crystallized melt inclusions. a, b – olivinehosted inclusions from Kovdor phoscorite. c – inclusion in clinopyroxene from the Gardiner complex, east Greenland. Abbreviations for phases: brad – bradleyite, Ca-Ba – Ca-Ba carbonates, c-hum – clinohumite, dol – dolomite, eit – eitelite, mag – magnesite (?), Na-sul – Na sulfate, she – scheelite; other abbreviations as in Fig. 6-2.

132


FIG. 6-4. Back-scattered electron image and X-ray elemental dot maps of a silicate-carbonate inclusion in olivine from peridotite, the Kovdor complex; nye, nyererite, other abbreviations same as in Fig. 6-2.

because of oxidation in air, precluding further observations. On the basis of electron microprobe analyses of the daughter minerals, and an estimate of their relative proportions, Aldous (1980, see also Rankin 2005) determined the bulk composition of melt inclusions (Table 6-3), and suggested that the olivine-hosted inclusions probably represented the earliest-formed carbonatite melt at Palabora. Importantly, melt trapped in the olivine-hosted inclusions with up to 0.3 wt.% Cu was already exceptionally enriched in this metal.

Aldous (1980) studied multi-solid, carbonate melt inclusions in olivine from the early-forming phoscorites. These inclusions typically contained magnesian calcite, tetraferriphlogopite (phlogopite characteristic of evolved carbonatite showing an extensive Fe3+–Al3+ substitution at the tetrahedral sites), low-Ti magnetite, and sulfide, but no obvious vapor bubble or associated aqueous fluid (Tables 6-1 and 6-3). In homogenization runs, some melting of the daughter phases was observed around 750oC, but above that temperature olivine darkened 133


Solovova et al. (1998) studied solid and highly concentrated fluid (salt melt) inclusions in apatite from phoscorite, and pyroxene from pyroxenite. Apatite was found to contain numerous solid inclusions up to 20–50 µm in size. Calcite is the predominant mineral of the solid inclusions, although clinopyroxene, amphibole, phlogopite, zeolites and some accessory phases are also abundant. A few thin hexagonal plates of portlandite [Ca(OH)2] were found in apatite. Portlandite was clearly identified by its refractive indexes in immersion oil, and analyzed by electron microprobe. The portlandite inclusions typically associate with a vapor bubble, and the importance of portlandite is discussed below. The multi-solid, salt melt inclusions in pyroxene and apatite are less abundant and usually very small (2–3 µm). Some salt melt inclusions in pyroxene are 20 µm across and large enough for freezing and homogenization experiments. The proportion of solids in the pyroxenehosted inclusions is 80–95 vol.%. The first signs of melting were observed at 700–730oC, but inclusions decrepitated above these temperatures. Modal proportion of daughter phases and bulk-chemical composition of the trapped liquid were calculated for a large (50 µm) inclusion in apatite (Tables 6-1 and 6-3). The composition is dominated by calcite, and phlogopite is the only identified alkali-bearing phase. In common with the melt composition calculated by Aldous (1980), the liquid is Ca-rich and very low in alkalis.

to the nahcolite- and nyerereite-bearing inclusions in apatite from sub-volcanic complexes in east Africa (Rankin 1975, 1977, Roedder 1984, Le Bas & Aspden 1981, see also Table 6-1), alkali sulfates or chlorides, which would suggest a hydrothermal origin, are absent from the Guli inclusions. Thus, the inclusions were interpreted as trapped solids, or melt inclusions. Regardless, the study demonstrated that water-soluble, Na–K–Ca carbonate can be preserved as micro-inclusions in a typical calciocarbonatite. The Guli complex of Triassic age (240 Ma) covers an area of 1500–1600 km2, and it is the largest intrusion within the Maimecha–Kotui province of ultramafic-alkaline-carbonatitic rocks (Kogarko et al. 1991, 1995). The Guli pluton includes a variety of rock types, which were intruded in the following order: dunite, pyroxenite, melilitic rocks, melteigite and jacupirangite, melanephelinite, syenite, and carbonatite. The carbonatite bodies outcrop in the northern (3 km2) and southern (5 km2) parts of the complex. Melilitic rocks were found to form small intrusive bodies in dunite. Rass & Plechov (2000) studied crystallized melt inclusions in early, forsteritic (Fo86-92) olivine from one small (120 m across) body of olivinemelilite plutonic rock (kugdite). The principal daughter phases of the inclusions are monticellite, Ca-rich pyroxene, nepheline, Cr-spinel; some inclusions also contain small amounts of apatite and calcite. The authors were able to homogenize and quench the inclusions to glass, although many inclusions quenched at 1221–1320oC contained quench (or residual?) monticellite. A typical composition of the quenched glass is presented in Table 6-2. The composition is close to the calculated bulk of unheated carbonate-bearing multi-solid inclusions in perovskite from early olivine cumulates of Kugda, another multiphase intrusion of the Maimecha-Kotui province. The inclusions in olivine from the Guli kugdite have been interpreted as samples of high-temperature, mantle-derived liquids, which probably had experienced a significant loss of CO2 by shallowlevel magma degassing, and presumably were parental to the olivine-melilite rock, as well as some other, more evolved rock types of the Guli complex.

Guli, Maimecha-Kotui Province, north-eastern Russia Kogarko et al. (1991) found Na-rich and SrBa-rich carbonate inclusions in accessory perovskite (CaTiO3) and calzirtite (Ca2Zr5Ti2O16) crystals from calciocarbonatite of the Guli complex of ultramafic-alkaline rocks and carbonatites at the northern border of the Siberian platform. The Narich phase was compositionally close to nyerereite Na2Ca(CO3)2, which is a typical phenocryst mineral of the Oldoinyo Lengai natrocarbonatite lava (see above). The inclusions are very abundant, irregular, and small (normally less than 30 µm). The discovery of the Na-rich inclusions has important implications for the composition of carbonatitic liquids parental to calciocarbonatite, because it demonstrates that although the Guli carbonatitic magma had obviously crystallized large masses of calcite, some derivatives evolved to very alkali-rich (and also Sr- and Ba-rich) compositions. In contrast

The Zeughaus dike, south-eastern Saxony, Germany Seifert & Thomas (1995) described crystallized melt inclusions form the Tertiary olivine melilitite (polzenite) Zeughaus dike in 134


Saxony. They found inclusions in olivine, pyroxene, apatite, phlogopite, and magnetite. The inclusions in olivine quenched to glass, and two types of average compositions (A and C, according to Seifert & Thomas 1995) are presented in Table 6-2. The more refractory olivine-hosted melt composition A has been interpreted as parental mantle-derived melilititic liquids, the other two compositions were proposed to result from silicate-carbonate liquid immiscibility.

of melilite- and perovskite-bearing melteigite and ijolite, and the younger comprises melilitolite with associated small dikes and veins of calciocarbonatite and peralkaline nephelinite. Detailed petrography of the multi-solid inclusions in Kovdor and Gardiner revealed numerous assemblages of daughter minerals, which include all the principal mineral species observed in the host rocks, but also unstable, water-soluble Na– Ca–K and Na–Mg carbonates, carbonate-chlorides, and carbonate-phosphates (Table 6-1). The widespread occurrence of shortite, Na2Ca2(CO3)3, in the inclusion assemblages was taken as a strong indicator of the primary origin and closed-system behavior of the inclusions, because shortite is a subsolidus phase not stable above 400oC (Cooper et al. 1975) and thus cannot be trapped as a solid in association with refractory silicate daughter phases. One of the puzzling features of the inclusion petrography, which was encountered early in the study, was the observation that, commonly, daughter mineral assemblages were very different in host minerals from the same rock. For example, olivine-hosted inclusions from Kovdor peridotite (olivine–clinopyroxene cumulate rock) are composed of clinopyroxene, phlogopite, magnetite, abundant Na–Ca carbonates, minor calcite, and apatite (Fig. 6-4), whereas the pyroxene-hosted inclusions from the same sample did not contain daughter olivine, and are composed of a richteritic amphibole, phlogopite, monticellite, magnetite, calcite, and only very minor Na–Ca carbonates (Veksler et al. 1998a). The difference in mineral assemblages was attributed to the following reaction:

Gardiner, East Greenland and Kovdor, Kola Peninsula, Russia Nielsen et al. (1997), Veksler et al. (1998a), and Sokolov (2002) have carried out a systematic study of crystallized melt inclusions from all the principal rock types of the Kovdor and Gardiner plutonic complexes. The study built upon the discovery of Na-Ca carbonates in the Guli complex (see above), and reported many new alkali-rich phases in the inclusions from calciocarbonatite and associated silicate rocks. The Kovdor complex is a concentrically zoned subvolcanic pluton emplaced at about 370 Ma into Archean granites and gneisses of the Baltic shield (see Kogarko et al. 1995 and Krasnova et al. 2004 for an overview of Kovdor geology). The earliest rocks, olivinite, occupy the central part of the intrusion, and are surrounded by irregular zones of peridotite, pyroxenite, melilite-bearing rocks, and rocks of the ijolitejacupirangite series. The youngest suite of rocks in Kovdor includes phoscorite and carbonatite, which form a stock in the southwestern part of the complex. Carbonatite evolves from calciocarbonatite to dolomite carbonatite. The carbonatitic-phoscoritic stock has been mined for magnetite, apatite and baddeleyite, and mapped and studied in great detail (Krasnova et al. 2004). The Gardiner complex in east Greenland formed at about 50 Ma during the early Tertiary break-up in the North Atlantic. The complex is about 5 km in diameter and dominated by early ultramafic cumulates (Nielsen 1981). They are interpreted as in situ cumulates formed on the walls of an open, shallow-level (50–100 MPa) subvolcanic magma chamber of a highly alkaline, carbonatite-bearing volcano (Nielsen et al. 1997). The ultramafic cumulates are intruded by a radial and ring dike system of increasingly alkaline composition (Nielsen 1980). Two major ring dikes, more than 100 m in width, are emplaced in the center of the complex. The older dike is composed

4 CaMgSi2O6 + Mg2SiO4 + Na2Ca(CO3)2 + H2O + CO2 = Na2CaMg5Si8O22(OH)2 + CaMgSiO4 + 3CaCO3 (1) With the excess of diopside component in pyroxene-hosted inclusions, nyerereite and olivine may be totally consumed by the reaction, and richterite becomes the main sink for Na. Another host-inclusion reaction has been considered to be responsible for the lack of daughter mineral assemblages with coexisting monticellite and diopsidic pyroxene: CaMgSi2O6 + Mg2SiO4 + 2 CaCO3 = 3 CaMgSiO4 + 2 CO2

(2)

Three types of crystallized melt inclusions were distinguished in Gardiner melilitolite (Nielsen 135


1999), and Sr-Ba carbonates and sulfates. Apart from the host olivine, tetraferriphlogopite is the only silicate phase stable in the evolved inclusion assemblages. Notably, the characteristic shortite and northupite were found as major constituents of carbonatitic inclusions and groundmass in kimberlite of the Udachnaya pipe (Kamenetsky et al. 2004, 2006, and Kamenetsky 2006).

et al. 1997). These are similar to those observed in Kovdor melilite-bearing rocks. Large primary silicate-carbonate inclusions (30–100 µm, see Fig. 6-2b) were interpreted as crystallized droplets of carbonated melilititic melt. Commonly, a significant interstitial volume of the inclusions is occupied by gas bubbles. Two additional types of the inclusions are not readily distinguished morphologically, but have different bulk compositions, and are believed to represent coexisting immiscible alkali carbonatite and natrocarbonatite melts (Tables 6-3 and 6-4). The larger melilite-hosted inclusions of the first type from Kovdor olivine melilitite are composed of diopside, phlogopite, apatite and magnetite, contain gas vesicles, but no carbonate daughter phases were found in them (Veksler et al. 1998a). The absence of carbonate daughter minerals was explained by the decarbonation reaction involving daughter pyroxene and the melilite host:

Sulfate-bearing and halogen-rich carbonatites from Mongolia and Transbaikalia, Russia Andreeva et al. (1998) reported unusual salt melt inclusions in fluorite crystals from calcitic carbonatite of the Mushugai-Khuduk alkalinecarbonatitic volcano-plutonic complex in southern Mongolia. The complex is composed of potassic volcanic and subvolcanic rocks (nepheline melaleucitite, melanephelinite, phonolite, trachyte, trachydacite, nepheline syenite, quartz-bearing syenite, and shonkinite). The carbonatite contains variable amounts of fluorite, chalcedony, celestine, barite, bastnaesite, dolomite or ankerite, and show gradual transition to a series of vein rocks, in which the modal proportion of calcite goes down, and the compositions are dominated by fluorite, celestine, magnetite, apatite, phlogopite, barite, and other minerals. The inclusions were studied mostly in the celestine–fluorite vein rock. Andreeva et al. (1998) called the rock carbonatite; however, it may represent a post-magmatic product of hydrothermal fluids, an example of “carbothermal” vein rocks (see Mitchell 2005b for discussion). The inclusions were sub-divided into 5 compositionally different types, and interpreted as portions of sulfate, phosphate and halogen-rich salt melts and/or very concentrated hydrothermal fluids. Some representative compositions of the inclusions are presented in Table 6-4. Doroshkevich et al. (2003) studied primary and secondary salt melt and concentrated fluid inclusions in bastnaesite, pyroxene and titanite from sulfate-rich carbonatite occurrences of western Transbaikalia, Russia. Late Mesozoic carbonatites in that area are related to potassic mafic alkaline igneous rocks, and commonly contain large amounts of sulfates (barite and celestine). Bulk rock SO3 content in some carbonatites exceeds 10 wt.% (e. g., up to 13.6 wt.% at the Khalyuta occurrence). Based on the inclusion evidence, the authors proposed that sulfate-rich inclusions implied immiscible separation of sulfate-carbonate liquid from a silicate melt.

CaMgSi2O6 + CaCO3 = Ca2MgSi2O7 + 2CO2 (3) Perovskite from Gardiner melilitolite contains abundant crystallized primary melt inclusions (Tables 6-2 and 6-3). Nielsen et al. (1997) and Veksler et al. (1998a) studied dozens of the inclusions, but daughter melilite was never found. The daughter minerals include wollastonite, pectolite and hydrogrossular, which are the typical products of low-temperature, fluid-present alteration of melilite. In addition, the inclusions contain clinopyroxene, nepheline, cancrinite and phlogopite. Calcite and Na–Ca carbonate (mostly shortite) are present in minor amounts. The most striking feature of the inclusions in both complexes is the strong variation in the proportion and the compositions of daughter carbonate minerals (Figs. 6-2 to 6-4). Two types of Na–Ca–K carbonate, compositionally close to nyerereite and shortite, are present in olivine-hosted inclusions from early dunite olivinite, and peridotite. The most evolved olivine-bearing rocks in Kovdor are phoscorite and forsterite-calcite carbonatite. Daughter calcite is nearly absent from olivine-hosted inclusions in phoscorite and carbonatite, although host rocks are dominated by calcite. Nyerereite and shortite in these inclusions are joined by dolomite and complex Na–Mg carbonates and mixed salts, such as eitelite Na2Mg(CO3)2, bradleyite Na3Mg(CO3)(PO4), and northupite Na3Mg(CO3)2Cl (Figs. 6-2c, 6-3a and 6-3b). Minor characteristic daughter minerals include K–Fe sulfides, such as djerfisherite K6(Fe,Cu,Ni)25S26Cl (see also Henderson et al. 136


sulfate- and chloride-rich inclusions were interpreted as products of carbonate–sulfate, or carbonate–chloride liquid immiscibility, but alternatively may represent “carbothermal” fluids (see previous section).

The Krestovskiy Complex, Polar Siberia, Russia The Krestovskiy complex is a multiphase silicate-carbonatite intrusion of the MaimechaKotui province of ultramafic–alkaline–carbonatitic igneous rocks in the northern part of the Siberian platform, 50 km southwest off the large Guli pluton (see above). The massif is oval in shape, and represents a layered body of ultramafic rocks consisting of thick (tens to hundred metres) layers of olivinite, wehrlite, and pyroxenite. Small (0.12 and 0.54 km2) bodies of melilitolite and monticellite-bearing rocks crop out at the western and eastern margins of the intrusion. Host rocks of the Krestovskiy plutonic complex are basalt and melanephelinite. The intrusion and the hosting effusive series are cut by dikes of dolerite, nepheline and melilite lamprophyres, and alkali picrite. Panina (2005) carried out a detailed study of melt and fluid inclusions in olivine, pyroxene, monticellite, melilite, perovskite, and garnet from all principal rock types of the complex. The most interesting results were obtained for inclusions in different growth zones of large euhedral perovskite crystals from the olivine–monticellite–melilite rocks. Dark-colored perovskite I in the core of the crystals contain rare, oval primary melt inclusions; the inclusions in the light-colored perovskite II at the rim are more abundant, and commonly have elongated or irregular shapes. The size of the inclusions ranges from 20 to 30 x 150 µm. The inclusions are composed of multiphase crystal aggregates comprising pyroxene, kalsilite, phlogopite, combeite, rankinite, apatite, djerfisherite, titanite, pectolite, and other accessory minerals. On heating, the inclusions exhibit liquid immiscibility between silicate and carbonate–sulfate melt in the temperature interval from 1130 to 1250oC. Some inclusions homogenize at 1230–1250oC, others do not, because they decrepitate on further heating. The inclusions that remained heterogeneous up to very high temperatures were interpreted as products of heterogeneous trapping of two immiscible liquids. The compositions of the immiscible silicate glass and sulfate–carbonate globules from the quenched perovskite-hosted inclusions are listed in Tables 6-2 and 6-3. Melilite, monticellite, pyroxene, and other silicate minerals host a broad variety of other melt and fluid inclusion types (Tables 6-3 and 6-4), and some compositions are much enriched in sulfates, phosphates, or chlorides. Six inclusion types have been recognized by Panina (2005) in total. The

Carbonatitic liquids in kimberlites: Lac de Gras, Canada, and the Udachnaya East pipe, Russia Van Achterbergh et al. (2002) reported quenched carbonate–silicate inclusions in clinopyroxene megacrysts (crystal size 1–20 cm) presumably excavated from 200 km depth beneath the Slave craton in northern Canada by kimberlite eruptions at Lac de Gras. Globular carbonate– silicate inclusions occur along distinct planes cutting across the clinopyroxene crystals, and show a bimodal size distribution (0.8–1.3 mm and 3.2–5.8 mm in diameter). Orthopyroxene exsolution lamellae in the host clinopyroxene extend to the edges of the inclusions, indicating that the entrapment of the globules preceded exsolution and therefore probably occurred at a temperature above 1240oC. Numerous planes of small fluid inclusions (1–20 µm), presumably filled with CO2, radiate from the carbonate inclusions into the clinopyroxene host. The inclusions consist of Tirich phlogopite, elongated forsteritic olivine (Fo89), Cr-spinel, and minor perovskite and sulfides, set in a matrix of calcite. Porous poikilitic clinopyroxenes line the walls of some inclusions, and enclose small grains of calcite and phlogopite. The inclusions were sub-divided into high-silica and low-silica groups, and the average compositions of the groups listed in Table 6-3 were calculated from the compositions of individual minerals, weighted by modal abundance. The carbonate-rich globular inclusions were interpreted as primary mantle carbonatite melts trapped and quenched at shallow depths during or after the kimberlite eruption. The authors proposed a temporal link between the entrapment of carbonatite and the Pliocene eruption of the kimberlite. Low bulk concentrations of incompatible Na and REE in the carbonatitic melt were thought to be inconsistent with an origin by low-degree partial melting of carbonated peridotite with low carbonate content (van Achterbergh et al. 2004). The low alkali and REE abundances were explained by melting of a source with a large modal proportion of Ca-Mg carbonate minerals, and the involvement of recycled crustal material has been inferred based on the oxygen, carbon, and strontium isotope data (van Achterbergh et al. 2002). 137


The composition in Table 6-2 represents another inclusion type found in Cr-diopside from lherzolite xenoliths brought to the surface by the Lac de Gras kimberlite bodies. They are concentric ultramafic silicate–carbonate globules, varying from 0.1 to 2 mm in diameter (van Achterbergh et al. 2004). The carbonate component, which crystallized as calcite, separates an outer layer of silicate glass (silicate-a) from a silicate-glass core (silicate-b). Small grains of pyrite and pyrrhotite occur in the silicate glass of some inclusions. The calcite is low in Sr and Ba (less than 1 wt.%), and enriched in rare earth elements. There is almost complete separation of chemical elements between the concentric layers, with very little CaO in the silicate glass (0.3-0.4 wt.%), and very little MgO and SiO2 in the carbonate. The composition in Table 6-2 is the calculated bulk average composition of the concentric silicate–carbonate globules, which are believed to represent trapped droplets of silicate-carbonate melt generated by partial melting of carbonated peridotite at high pressure. A very different composition has been found in olivine-hosted inclusions and the groundmass of group I kimberlite from the Udachnaya East diamondiferous pipe in Siberia, Russia (Kamenetsky et al. 2004, 2006, see also Kamenetsky 2006). In contrast to the melts at Lac de Gras, the inclusions and groundmass of the Udachnaya pipe (Tables 6-2 and 6-3) are extremely enriched in water-soluble alkali chlorides, alkali carbonates, and sulfates, and show high concentrations of REE and other incompatible trace elements. The composition in Table 6-2 is the bulk of the kimberlitic groundmass (Kamenetsky et al. 2004), while the composition in Table 6-3 represents a chloride–carbonate globule of presumably immiscible origin (Kamenetsky et al. 2006). Although it was argued that alkali carbonates, sulfates and chlorides may result from contamination by crustal evaporites (Golovin et al. 2003), the detailed petrography, and oxygen, carbon, and strontium isotopes (Kamenetsky et al. 2004, 2006) strongly support the mantle origin of water-soluble salt components. The mantle origin of the exotic melt composition at Udachnaya is further corroborated by findings of Na–Ca carbonates and K-rich brines as abundant and characteristic inclusions in diamonds (Izraeli et al. 2001, 2004).

IMPLICATIONS OF MELT INCLUSION STUDIES FOR CARBONATITE PETROGENESIS Compositions of primary mantle-derived liquids Melt inclusions in early liquidus minerals from ultramafic–alkaline–carbonatitic plutonic complexes worldwide document silica-deficient (34–37 wt.% SiO2), carbonated silicate melts with high TiO2 contents (~4–6 wt.%) and CaO/Al2O3 weight ratios up to about 5 (Table 6-2). The melts are also characterized by rather high alkali contents, and they are commonly peralkaline (molar (Na2O+K2O)/Al2O3 > 1). Broad variations in MgO/(FeO+MgO) of the compositions listed in Table 6-2 probably result from fractionation of early Mg–Fe silicate minerals, such as olivine and diopsidic clinopyroxene. More evolved inclusion compositions (mostly from melilite-bearing rocks) are close to silicate–carbonate rocks such as bergalite or okaite (Barker 2001). The highest MgO contents and Mg# are documented in the inclusions from kimberlite. Notably, the kimberlitic compositions having about the same Mg# show very different alkali contents, and the alkali-rich composition at the Udachnaya pipe (Kamenetsky et al. 2004) is also several times to orders of magnitude higher in Ti, P, Ba, Sr, and Cl than the composition at Lac de Gras (van Achterbergh et al. 2004). The compositions in Table 6-2 are in general agreement with the diversity of liquids produced in partial melting experiments on a carbonated mantle peridotite at pressures above 2 GPa (Wyllie & Huang 1976, Eggler 1978, Wallace & Green 1988, Thibault et al. 1992, Dalton & Wood 1993, Dalton & Presnall 1998, Wyllie & Lee 1998). However, compositions with Cl contents at a weight percent level have never been considered in experimental work to date. Recent discoveries of KCl- and NaCl-rich brines in diamond- and olivine-hosted melt and fluid inclusions from different kimberlite localities worldwide (Izraeli et al. 2001, 2004, Kamenetsky et al. 2004, 2006) imply that chloriderich liquids may be important metasomatic agents in the mantle, and deserve more attention in the future. The composition of perovskite-hosted melt inclusions from early olivinite cumulate of the Kugda intrusion, the Maimecha-Kotui province, Russia (Nielsen et al. 1997), appears to be the leastevolved of the silicate–carbonate liquid compositions from ultramafic–alkaline–carbonatitic complexes listed in Table 6-2. It also seems to retain volatile components, and has not degassed as much as similar inclusions in olivine from the Guli 138


complex (Rass & Plechov 2000). It appears to be a good candidate for a “primitive” (not much affected by shallow-level crystallization and magma degassing) mantle-derived parental liquid of Kugda and the whole clan of similar ultramafic–alkaline– carbonatitic complexes.

(Fig. 6-5). The join was later revised and amended by Schairer et al. (1962), but the earlier, incomplete version in Fig. 6-5 better demonstrates the essence of phase equilibria expressed by equation 4. It should be noted that in the synthetic system, and in the majority of natural ijolitic and melilitolitic rocks, the albite component on the right side of the equation does not form a separate feldspar phase, but resides in melt and/or nepheline solid solution. Natural rocks and crystallized melt inclusions imply that carbonates and CO2 are important constituents of ijolitic and melilititic magmas. Thus, the reaction between the diopside + nepheline association and melilite transforms into a series of decarbonation reactions. The first reaction, which involves diopside and akermanite components of clinopyroxene and melilite solid solutions, is presented above by equation 3. Another decarbonation reaction involves nepheline and Namelilite components:

Mineral reactions and relationships between ijolite, melilitite and carbonate At any scale, from relative abundances of principal volcanic and plutonic rock types (Woolley 2003, and Fig. 6-1 of this chapter) to variations in modal proportions of daughter minerals in microscopic crystallized melt inclusions (e.g., Veksler et al. 1998a, and the equations 1 to 3 above), mineral reactions vividly manifest themselves in ultramafic–alkaline–carbonatitic complexes. Bowen (1928), who introduced the fundamental reaction principle, was also probably the first to point out that ijolite (nepheline + clinopyroxene) and melilitolite rock associations are related to each other by the following reaction: 3CaMgSi2O6 + 2NaAlSiO4 = diopside nepheline Ca2MgSi2O7 + NaCaAlSi2O7 + Mg2SiO4 + melilite solid solution olivine NaAlSi3O8 albite

NaAlSiO4 + CaCO3 + SiO2 = NaCaAlSi2O7 + CO2

In effect, reactions 3, 4 and 5 show that the same Ca-rich and silica-deficient melt can crystallize either as a calcite-bearing ijolite or as a melilitolite. The decarbonation reactions, which produce CO2, obviously strongly depend on pressure, and consequently the formation of melilite-bearing rocks appears to be favored by magma degassing. This may be one of the reasons why melilitic rocks,

(4)

Bowen (1922) encountered the reaction in natural rocks and on the synthetic diopside–nepheline join o

T, C 1500 Diopside+Liquid

1400 1300

Olivine+Diopside +Liquid 1200 1100

Carnegieite +Liquid

Olivine+Liquid

Nepheline+Liquid Melilite+ Olivine +Liquid

Melilite+Diopside+Liquid

CaMgSi2O6 Diopside

20

(5)

40

Nepheline+Melilite +Olivine+Liquid

60

Weight %

80

NaAlSiO4 Nepheline

FIG. 6-5. The diopside – nepheline phase diagram, after Bowen (1922).

139


as pointed out by Woolley (2003), are more common in volcanic rock associations than in silicate–carbonate plutonic complexes (Fig. 6-1). Some reactions (like reaction 5) also involve silica, and Barker (2001) recently used such reactions for calculations of silica activities in natural carbonatitic liquids. Veksler et al. (1998a) described numerous examples of decarbonation and other mineral reactions in crystallized melt inclusions from Kovdor and Gardiner (see above). They proposed two scenarios of shallow level magma evolution depending on the stability of liquidus melilite. Melts crystallizing early liquidus melilite, but retaining a significant amount of CO2, were expected to evolve towards the silicate–carbonate miscibility gap, and produce immiscible silicate and carbonate melts. Three groups of melt inclusions in melilite and perovskite from Gardiner melilitolite (Nielsen et al. 1997, Veksler et al. 1998a) were considered as examples of such types of melt evolution. It was proposed that when melilite is not stable, the crystallization paths may not intersect the twoliquid field, but reach the calcite saturation boundary via fractionation of the nepheline plus clinopyroxene (± phlogopite) liquidus assemblage. Further evolution along the calcite-saturated cotectic is expected to produce a series of silicate– carbonate liquids with decreasing SiO2 and increasing alkali contents. A sequence of inclusion compositions from Kovdor (Table 6-3) may represent such a liquid evolution trend along the silicate–carbonate multiply-saturated cotectic toward late-stage alkali-rich compositions sampled by melt inclusions in the Kovdor phoscorite.

halogen-rich, or sulfate-rich compositions reported for some evolved crustal carbonatites (e.g., Krestovskiy, Mushgai-Khuduk, or above-mentioned occurrences in Transbaikalia) may also form in the deep mantle. Many compositions observed in the inclusions are not properly covered by experiments in synthetic systems. The available melt inclusion record is still too sketchy and incomplete, and thus universal models explaining the apparent diversity of natural carbonatitic liquids would probably be premature. Water-soluble daughter phases, including Na–Ca carbonates are hard to detect, and could have been overlooked in early studies of carbonatitic inclusions. Host–melt reactions (Veksler et al. 1998a, see also equations 1-4 above) may also destabilize Na–Ca carbonates in inclusions. However, the evidence for low-alkali, calciocarbonatitic liquid at Palabora (Tables 6-1 and 6-3) has been recently supported by the occurrences of portlandite in apatite-hosted inclusions (Solovova et al. 1998, see above). The occurrence of portlandite links low-alkali melt inclusions at Palabora with the well-characterized CaO–CO2– H2O and CaO–SiO2–CO2–H2O synthetic systems (Wyllie 1989, Wyllie & Tuttle 1959, 1960), where low-temperature calciocarbonatitic liquids are stabilized by the presence of dissolved water, and portlandite appears as an important component. The occurrence of daughter fluorite in inclusions at Palabora (Solovova et al. 1998) also suggests a potential role of F as a fluxing component (Jago & Gittins 1990). The role of silicate-carbonate liquid immiscibility Direct observations of silicate–carbonate immiscibility in natural melt inclusions (Rankin & Le Bas 1974, Romanchev & Sokolov 1980, Panina 2005) support the views that this process can be responsible for the origin of some natural carbonatites (Kjarsgaard & Hamilton 1989, Bailey 1993, Kjarsgaard 1998). The experimental constraints on silicate–carbonate unmixing at pressures up to 0.5 GPa are especially robust (e. g. Kjarsgaard & Hamilton 1989, Brooker 1998, Kjarsgaard 1998, Veksler et al. 1998b and references therein). For instance, the partitioning of key alkalis and alkaline earth elements between the immiscible silicate and carbonate minerals is well constrained experimentally for a broad range of bulk compositions and P–T conditions. Nernst partition coefficients (D = CLCi/CLSi, where CLCi and CLSi are weight concentrations of element i in

Parental liquids of plutonic carbonatites According to experimental phase equilibria constraints (e.g., Lee & Wyllie 1998), lowtemperature carbonatitic liquids at shallow, crustal depth can theoretically range from water-bearing calciocarbonatite compositions to dry natrocarbonatite fluxed by high contents of alkali carbonate components. Melt and fluid inclusion compositions presented in Tables 6-3 and 6-4 appear to support the notion that chemical types of natural carbonatitic liquids are diverse. The inclusions from Magnet Cove, USA, and Palabora, South Africa, seem to represent the liquids at the low-alkali end of the spectrum. The inclusions at Kovdor, Gardiner, and Krestovskiy are in general more alkali-rich. The example of the Udachnaya pipe (Tables 6-2 and 6-3) implies that exotic 140


carbonate liquid LC and silicate liquid LS) are the simplest and most straightforward measure of liquid–liquid element partitioning. When plotted against ionic potentials Z/r (where Z is the nominal charge of the metal cation, and r is the ionic radius), the D values of the alkali and alkaline earth elements form a convex curve with the maximum at Sr (Fig. 6-6a). Two examples presented in Fig. 6-6a and Table 6-5 are a peralkaline nephelinitenatrocarbonatite liquid pair modeling the Oldoinyo Lengai compositions (Veksler et al., unpublished; later, and refined experiments of the same type as described by Veksler et al. 1998b), and a low-alkali liquid pair in equilibrium with melilite, melanite, and nepheline (Kjarsgaard 1998). The examples show that although melt composition, pressure, and temperature do affect element partitioning, the effects are not very great, and amount to a shift of the convex curve to the right or to the left from the maximum (see also Brooker 1998 for the effects of pressure up to 2.5 GPa). In a broader range of experimental conditions, the absolute D values may increase up to 10 for Sr (e.g., Kjarsgaard 1998, sample BK316), but the general convex trend does not change. In fact, D values of the mono- and divalent cations plot along the convex upward trend against Z/r not only for carbonate-silicate, but any other salt–silicate immiscible pair of liquids (Veksler 2004). Earlier inclusion studies (Rankin & Le Bas 1974, Romanchev & Sokolov 1980) did not report the compositions of the conjugate liquids. With a few exceptions (Le Bas & Aspden 1981, Nielsen et al. 1997, Panina 2005), which are discussed below, electron microprobe data (Tables 6-3 and 6-4) reveal relatively Ca-rich immiscible carbonatite compositions, closer to the experimental liquids in equilibrium with a melilitic melt (Kjarsgaard 1998) than peralkaline compositions at Oldoinyo Lengai. Average compositions of two presumably immiscible groups of inclusions were reported by Nielsen et al. (1997), and Panina (2005) analyzed quench silicate glass and carbonate–sulfate globules in individual perovskite-hosted inclusions (Tables 6-2 and 6-3). The calculated D values for these examples are presented in Fig. 6-6b. Sr and Ba were rather low for reliable electron microprobe determination in the inclusions studied by Nielsen et al. (1997), but the remaining D values in the Gardiner samples are in general agreement with the experimental convex curve. Although the inclusions in Krestovskiy appear to represent a straightforward case of liquid immiscibility, the D values calculated

D carbonate/silicate

10

Ba

Sr

a

Ca

Na Mg

K

Mn

1

Rb Cs

OL41

BK254

0.1 0

10

Z/r

20

30

D carbonate/silicate

10

K Na

b

Ba Sr Ca

1

Mg

0.1

Gardiner Krestovskiy

Mn

0.01 0

10

Z/r

20

30

FIG. 6-6. Carbonate-silicate liquid-liquid Nernst partition coefficients in synthetic systems (a) and melt inclusions (b). See Table 6-5 and text for more detail.

from the available electron microprobe data (Panina, 2005, see also Tables 6-2, 6-3 and 6-5) do not fit well with the convex D–Z/r curve. In particular, the D value for K appears to be too high, and DSr is too low. The reasons for the discrepancy are not clear, and surely more microprobe analyses of immiscible inclusions would be very welcome. Immiscible silicate and carbonate liquids are traditionally plotted on the so-called Hamilton projection (e.g., Kjarsgaard & Hamilton 1989 and references therein), and such a plot for the inclusion compositions in Tables 6-2, 6-3 and 6-4 is presented in Figure 6-7. The extent of the experimentally determined miscibility gap in the high-Ca nephelinite–carbonatite system at 0.2 GPa and a few tie-lines (Kjarsgaard 1998) are shown for 141


TABLE 6-5. LIQUID-LIQUID CARBONATE/SILICATE NERNST PARTITION COEFFICIENTS D’S (SEE TEXT FOR THE DEFINITION) IN EXPERIMENTS WITH SYNTHETIC MIXTURES, AND NATURAL IMMISCIBLE MELT INCLUSIONS. Ionic properties Sample T, C P, GPa H2Oa, wt.% Ion Cs+ Rb+ K+ Na+ Li+ Ba2+ Sr2+ Ca2+ Mg2+ Mn2+ Fe3+ Al3+ Ti4+ Si4+ Nb5+ P5+ S6+ FClReference

CN 8 6 6 6 5 6 6 6 6 6 6 4 5 4 6 4 4 -

r, pm 188 166 152 116 82 149 132 114 86 81 69 53 66 40 78 17 14 133 181

Z/r x 1000 5.3 6.0 6.6 8.6 12.2 13.4 15.2 17.5 23.3 24.7 43.5 56.6 60.6 100.0 64.1 294.1 428.6 -7.5 -5.5

Shannon 1976

Experiments with synthetic mixtures

Natural inclusions Gardiner, Krestovskiy, OL41 OL2 BK254 BK316 melilite perovskite 850 850 950 900 1050 1170 0.1 0.1 0.2 0.5 n. d. n. d. dry 11 dry dry n. d. n. d. Carbonate/silicate partition coefficients (D) 0.69 0.65 n. d. n. d. n. d. n. d. 0.87 0.64 n. d. n. d. n. d. n. d. 1.29 0.96 0.32 0.28 0.83 2.16 2.02 1.01 0.93 1.25 1.79 1.35 2.51 1.35 n. d. n. d. n. d. n. d. 4.03 2.00 2.47 7.79 0.02 2.25 4.24 1.58 4.10 10.08 1.10 1.40 2.77 2.34 3.72 18.38 2.35 1.26 0.49 0.62 1.17 4.93 0.17 0.38 0.30 0.35 0.67 2.87 0.04 n. d. 0.04 0.34 0.42 0.47 0.05 0.19 0.006 0.19 0.013 0.014 0.005 0.13 0.09 n. d. 0.066 0.221 0.02 0.33 0.03 0.33 0.079 0.012 0.007 0.20 0.21 0.42 n. d. n. d. n. d. n. d. n. d. 6.04 4.93 n. d. 0.85 1.24 n. d. n. d. n. d. n. d. n. d. 13.2 6.11 4.15 3.83 7.93 n. d. n. d. 5.94 2.05 2.24 5.27 n. d. 5.38 Nielsen Veksler et al. Panina 2005 Kjarsgaard 1998 et al. 1997 unpublished

a

– bulk water content in starting mixtures. CN – coordination number to oxygen anions; r – ionic radius; Z – nominal ionic charge; Z/r – ionic potential; n. d. – not determined. Ionic radii are according to Shannon (1976).

comparison. The projection in Figure 6-7 shows that a number of melt inclusion compositions plot inside the experimental miscibility gap. This is probably not surprising, and one should not expect perfect agreement with the experimental data in this case, because the inclusions represent a broad diversity of liquid compositions and P–T conditions. All of those factors are known to affect the extent and exact position of the silicate–carbonatite miscibility gap. The projection in Figure 6-8 attempts to use a few key reference inclusion compositions for tracing the main topological elements of the Hamilton projection, that is, the liquid miscibility gap, and the carbonate saturation curve. The miscibility gap is constrained by the immiscible inclusions from Krestovskiy (Panina 2005), and the silicate–calcite cotectic is traced by the inclusion compositions from Magnet Cove, Palabora, and Kovdor (Nesbitt & Kelly 1977, Aldous 1980, Veksler et al. 1998a, see Tables 6-2, 6-3 and 6-4).

The miscibility gap constrained by the natural inclusion compositions appears to be somewhat narrower than the maximal extent reported in experimental systems (Fig. 6-7), but more analyses of the inclusions are needed for a more accurate comparison. Melt-fluid immiscibility Melt and fluid inclusion studies have provided crucial information about the nature and composition of aqueous fluids exsolved from carbonatites and coexisting silicate rocks. The best fluid data come from inclusions in apatite (Rankin 1975, 1977, 2005, Roedder 1984, and references therein), but also from rare examples of quartzhosted inclusions from fenitization zones around carbonatite bodies (Bühn & Rankin 1999). Fluid compositions are in general very alkali-rich; some inclusions may contain weight percent concentrations of halogens and SO3, up to 3 wt.% 142


Na2O + K2O

nephelinitic-melilititic liquids kimberlitic liquids low-alkali carbonatitic melts

Kr

phoscoritic melt high-alkali melts and fluids Kf Ic Gr Uc

900

Kr

950 Gu Gr

Kr

Ko

Gr

960 Ko

Kr

Pa

Ud

Ku

SiO2 + TiO2 + Al2O3

Ko

Pa

LG MC

LG

CaO + MgO + FeO

FIG. 6-7. Compositions of melt and fluid inclusions from Tables 2, 3 and 4 plotted onto the Hamilton projection. The silicatecarbonatite miscibility gap for low-alkali compositions at 0.2 GPa (Kjarsgaard 1998) is shown by the heavy solid curve; experimental tie-lines are shown by dashed lines with numbers indicating the experimental temperatures. Abbreviations for the locations: Gr – Gardiner, Gu – Guli, Ic and Uc – inclusions from ijolite and urtite in Kenya (Le Bas & Aspden 1981), Kr – Krestovskiy, Ko – Kovdor, Kf – Kalkfeld, MC – Magnet Cove, LG – Lac de Gras, Pa – Palabora, Ud – Udachanaya pipe.

Na2O + K2O Kr

Kf Ic

Two liquids

?

Gr cc-nye

Uc Kr

70

Ko

Ca

10

lc

Kr

Pa MC

SiO2 + TiO2 + Al2O3

e

S i l i c a t e s

it

Ko

CaO + MgO + FeO

FIG. 6-8. Main topological elements of the Hamilton projections inferred from melt and fluid inclusions. See text for discussion; abbreviations same as in Fig. 6-7.

143


ΣREE (Bühn & Rankin 1999), and up to a few weight percent of Sr and Ba (Tables 6-1 and 6-4). Some compositions show no detectable silica (e.g., apatite-hosted inclusions in Kovdor, Veksler et al. 1998a); others have SiO2 concentrations at a level typical for immiscible carbonatite liquids (Le Bas & Aspden 1981). In the Hamilton projection (Figs. 6-7 and 6-8), the inclusion compositions form a group of points stretching from the calcite–nyerereite cotectic of the Na2CO3–K2CO3–CaCO3 system (Cooper et al. 1975) to the alkaline corner. The topology of the alkali-rich portion of the projection is poorly constrained. It is not easy to distinguish between two-fluid and three-fluid equilibria (Nielsen et al. 1997, Nielsen & Veksler 2002, Veksler 2004). It is not quite clear how fluid compositions cross the calcite-nyerereite cotectic, although some explanations involving the effect of fluorine addition have been offered (Jago & Gittins 1990; see discussions by Nielsen & Veksler 2002). The positions of critical curves in relevant multicomponent silicate–carbonate–H2O systems have not been defined. Some very preliminary experimental constraints on the topology are provided by the studies of the Na2CO3–H2O system (Ravich 1974, Koster van Groos 1990), in which the upper critical end-point has been located at ~480oC and 160 MPa (Ravich 1974). The partitioning of Ca, Mg, and Na between hydrothermal fluids and carbonatitic melts was experimentally studied by Veksler & Keppler (2000), and, as one would expect, the study confirmed a higher solubility of Na in the fluids relative to Ca and Mg. In view of the unclear general topology, the alkali-rich compositions of apatite-hosted inclusions from ijolite and urtite in Kenya (Le Bas & Aspden 1981, see also Table 6-4 and Figs. 6-7 and 6-8) can be interpreted either as immiscible natrocarbonatite melt, or as highly concentrated hydrothermal brines. As discussed by Nielsen & Veksler (2002, and references therein), a similar dilemma stands for the natrocarbonatite lava at Oldoinyo Lengai. Some geochemical features of natrocarbonatite agree with the long-proposed origin by dry silicatecarbonate liquid immiscibility, however, other significant details do not comply with the one-stage exsolution model, and hint towards the involvement of hydrothermal fluids. Clearly, more experimental and melt inclusion studies are needed to clarify the issues.

Other types of liquid unmixing Direct observations on melt and fluid inclusions during heating experiments reveal liquid immiscibility between halogen-rich brines and molten carbonates (Fulignati et al. 2001, Webster & De Vivo 2002, Kamenetsky et al. 2004, 2006, and Kamenetsky 2006). Chloride–carbonate liquid unmixing has not been reproduced experimentally in synthetic systems, and Mitchell (1997) presented evidence for this type of liquid immiscibility in the groundmass of the Oldoinyo Lengai natrocarbonatite. The chloride–carbonate unmixing, and other types of immiscibility in salt melts and brines are discussed in detail by Kamenetsky (2006). This phenomenon requires better experimental characterization in the future, because it may play an important petrogenetic role not only in crustal carbonatites and skarns, but also in kimberlite from the deep mantle. CONCLUSIONS Melt and fluid inclusion studies have been indispensable for the development of petrogenetic models for the origin of carbonatites and the carbonatite–silicate igneous rock association. Melt inclusions in early cumulus silicates constrain the compositions of silica-undersaturated, carbonated melts, which may represent parental magmas of ultramafic–alkaline–carbonatitic plutonic complexes. Silica content at about 34–37 wt.%, high TiO2 concentrations (4–6 wt.%), CaO/Al2O3 up to 5 by weight, and total alkalis at about 7–8 wt.% are the main characteristic features of early liquids. Melt and fluid inclusions serve as direct examples of silicate–carbonate liquid immiscibility, alkalirich fenitizing fluids, and complex mineral reactions controlling the evolution of natural carbonatitic magma. Melt and fluid compositions in the inclusions appear to evolve to alkali-rich carbonatitic liquids with high contents of Cl, F, SO3, Sr, Ba, and REE, approaching the composition of the Oldoinyo Lengai natrocarbonatite lava. ACKNOWLEDGEMENTS We thank Jim Webster for his outstanding organizational and editorial efforts in preparing this short course volume. Reviews by Roger Mitchell and Dima Kamenetsky helped to improve the earlier version of the chapter. I.V.V. acknowledges the support from the research fellowship program at the American Museum of Natural History (New York, N.Y.), the CERCAMS grant (Natural History Museum, London), and the European SYNTHESYS 144


COOPER, A.F., GITTINS, J. & TUTTLE, O.F. (1975): The system Na2CO3–K2CO3–CaCO3 at 1 kilobar and its significance in carbonatite petrogenesis. Am. J. Sci. 275, 534–560.

program. The research by DL was supported by a NSERC Discovery Grant. REFERENCES ALDOUS, R. (1980): Ore genesis in copper bearing carbonatites: a chemical, mineralogical and fluid inclusion study. Unpublished Ph.D. thesis, Imperial College, Univ. London.

DALTON, J.D. & PRESNALL, D.C. (1998): The continuum of primary carbonatitic-kimberlitic melt compositions in equilibrium with lherzolite: data from the system CaO–MgO–Al2O3–SiO2– CO2 at 6 GPa. J. Petrol. 39, 1953–1964.

ANDREEVA, I.A., NAUMOV, V.B., KOVALENKO, V.I. & KONONKOVA, N.N. (1998): Fluoride–sulfate and chloride–sulfate salt melts in the carbonatite-bearing complex Mushugai-Khuduk, Southern Mongolia. Petrology 6, 284–292.

DALTON, J.D. & WOOD, B.J. (1993): The composition of primary carbonate melts and their evolution through wall rock reaction in the mantle. Earth Planet. Sci. Lett. 119, 511–525.

BAILEY, D.K. (1990): Mantle carbonatite eruptions: crystal context and implications. Lithos 26, 37–42.

DALY, R. A. (1910): Origin of alkaline rocks. Geol. Soc. Am. Bull. 21, 87–118.

BAILEY, D.K. (1993): Carbonate magmas. J. Geol. Soc. (Lond.) 150, 637–651.

DAWSON, J.B. (1962): Sodium carbonate lavas from Oldoinyo Lengai, Tanganyika. Nature 195, 1075– 1076.

BARKER, D.S. (2001): Calculated silica activities in carbonatite liquid. Contrib. Mineral. Petrol. 141, 704–709.

DEINES, P. (1989): Stable isotope variations in carbonatites. In Carbonatites: Genesis and Evolution. (K. Bell, ed.), London, Unwin Hyman, p. 301–359.

BELL, K. & BLENKINSOP, J. (1989): Neodymium and strontium isotope geochemistry of carbonatites. In Carbonatites: Genesis and Evolution (Bell, K., ed.). London, Unwin Hyman, 278–300.

DOROSHKEVICH, A.G., KOBYLKINA, O.V. & RIPP, G.S. (2003): Role of sulfates in the formation of carbonatites in the Western Transbaikal region. Doklady Earth Sci. 388, 131–134.

BELL, K. & KELLER, J. (1995): Carbonatite Volcanism. Oldoinyo Lengai and Petrogenesis of Natrocarbonatites. Springer, Berlin, Heidelberg, New York.

DREW, L.J., QUINGRUN, M. & WEIJUN, S. (1990): The Bayan Obo iron–rare-earth–niobium deposits, Inner Mongolia, China. Lithos 26, 43–65.

BOWEN, N.L. (1922): Genetic features of alnöitic rocks at Isle Cadieux, Quebec. Am. J. Sci. 203(3), 1–13.

EGGLER, D.H. (1978): The effect of CO2 upon partial melting of peridotite in the system Na2O–CaO– Al2O3–MgO–SiO2–CO2 to 35 kb, with an analysis of melting in a peridotite–H2O–CO2 system. Am. J. Sci. 278, 305–343.

BOWEN, N. L. (1928): The Evolution of the Igneous Rocks. Princeton University Press, Princeton, NJ.

ERIKSSON, S.C. (1989): Palabora: a saga of magmatism, metasomatism and miscibility. In Carbonatites: Genesis and Evolution (K. Bell, ed.). Unwin Hyman, London, 221–254.

BREY, G. & GREEN, D.H. (1977): Systematic study of liquidus phase relations in olivine melilitite– H2O+CO2 at high pressures and petrogenesis of an olivine melilitite magma. Contrib. Mineral. Petrol. 61, 141–162.

FOLEY, S.F., VENTURELLI, D.H., GREEN. D.H. & TOSCANI, L. (1987): The ultrapotassic rocks: characteristics, classification, and constraints for petrogenetic models. Earth Sci. Rev. 24, 81–134.

BROOKER, R.A. (1998): The effect of CO2 saturation on immiscibility between silicate and carbonate liquids: an experimental study. J. Petrol. 39, 1905–1915.

FULIGNATI, P., KAMENETSKY, V.S., MARIANELLI, P., SBRANA, A. & MERNAGH, T.P. (2001): Melt inclusion record of immiscibility between silicate, hydrosaline, and carbonate melts: applications to skarn genesis at Mount Vesuvius. Geology 29, 1043–1046.

BÜHN, B. & RANKIN, A.H. (1999): Composition of natural, volatile-rich Na–Ca–REE–Sr carbonatitic fluids trapped in fluid inclusions. Geochim. Cosmochim. Acta 63, 3781–3797. 145


JOHNSON, R.L. (1961): The geology of the Dorowa and Shawa carbonatite complexes, southern Rhodesia. Geol. Soc. S. Africa Trans. 64, 101– 145.

GENGE M.J., JONES A.P & PRICE G.D. (1995): An infrared and Raman study of carbonate glasses: Implications for the structure of carbonatite magmas. Geochim. Cosmochim. Acta 59, 927– 937.

M.B., SOBOLEV, A.V., KAMENETSKY, KAMENETSKY, V.S., MAAS, R., DANYUSHEVSKY, L.V., THOMAS, R., POKHILENKO, N.P., & SOBOLEV, N.V. (2004): Kimberlite melts rich in alkali chlorides and carbonates: a potent metasomatic agent in the mantle. Geology 32, 845–848.

GIRAUT, J. (1966): Genèse et géochemie de l’apatite et de la calcite dans les roches liées ou complex carbonatique et hyperalcalin d’Oka (Canada). Bull. Soc. Franç. Min. Crist. 89, 496– 513. GITTINS, J. (1989): The origin and evolution of carbonatite magmas. In Carbonatites: Genesis and Evolution (K. Bell, ed.). Unwin Hyman, London, 580–600.

KAMENETSKY, V.S. (2006): Melt inclusion record of magmatic immiscibility in crustal and mantle magmas. In Melt Inclusions in Plutonic Rocks (J.D. Webster, ed.) Min. Assoc. Can. Short Course 36, 81–98.

GOLOVIN, A.V. SHARYGIN, V.V. POKHILENKO, L.N., MAL’KOVETS, V.G., KOLESOV, B.A. & SOBOLEV, N.V. (2003): Secondary melt inclusions in olivine from unaltered kimberlites of the Udachanya East pipe, Yakytia. Doklady Earth Sci. 388, 93–96.

KAMENETSKY, V.S., KAMENETSKY, M.B., SHARYGIN, V.V., FAURE, K. & GOLOVIN, A.V. (2006): Chloride and carbonate immiscible liquids at the closure of the kimberlite magma evolution (Udachnaya East kimberlite, Siberia). Chem. Geol. (in review)

GUDFINNSSON, G.H. & PRESNALL, D.C. (2005): Continuous gradations among primary carbonatitic, kimberlitic, melilititic, basaltic, picritic and komatiitic melts in equilibrium with garnet lherzolite at 3–8 GPa. J. Petrol. 46, 1645– 1659.

KEPPLER, H. (2003): Water solubility in carbonatite melts. Am. Mineral. 88, 1822–1824. KJARSGAARD, B.A. (1998): Phase relations of a carbonated high-CaO nephelinite at 0.2 and 0.5 GPa. J. Petrol. 39, 2061–2075.

HARMER, R.E. & GITTINS, J. (1998): The case for primary, mantle-derived carbonatite magma. J. Petrology 39, 1895–1903.

KJARSGAARD, B.A. & HAMILTON, B.A. (1989): The genesis of carbonatites by immiscibility. In Carbonatites: Genesis and Evolution (K. Bell, ed.). Unwin Hyman, London, 388–404.

HENDERSON, C.M.B., KOGARKO, L.N. & PLANT, D.A. (1999): Extremely closed system fractionation of volatile-rich, ultrabasic peralkaline melt inclusions and the occurrence of djerfisherite in the Kugda alkaline complex, Siberia. Mineral. Mag. 63, 433–438.

KOGARKO, L.N., KONONOVA, V.A., ORLOVA, M.P. & WOOLLEY, A.R. (1995): Alkaline Rocks and Carbonatites of the World. Part 2: Former USSR. Chapman & Hall, London.

HÖGBOHM, A.G. (1895): Über das Nephelinsyenotgebeit auf der Insel Alnö. Fören. Stockholm Förh. 17, 100–160.

KOGARKO, L.N., PLANT, D.A., HENDERSON, C.M.B. & KJARSGAARD, B.A. (1991): Na-rich carbonate inclusions in perovskite and calzirtite from the Guli intrusive Ca-carbonatite, Polar Siberia. Contrib. Mineral. Petrol. 109, 124–129.

IZRAELI, E.S., HARRIS, J.W. & NAVON, O. (2001): Brine inclusions in diamonds: a new upper mantle fluid. Earth Planet. Sci. Lett. 187, 323–332.

KOSTER VAN GROOS, A.F. (1990): High-pressure DTA study of the upper three-phase region in the system Na2CO3–H2O. Am. Mineral. 75, 667–675.

IZRAELI, E.S., HARRIS, J.W. & NAVON, O. (2004): Fluid and mineral inclusions in cloudy diamonds from Koffiefontein, South Africa. Geochim. Cosmochim. Acta 68, 2561–2575.

KOSTER VAN GROOS, A.F. & WYLLIE, P.J. (1968): Liquid immiscibility in the join NaAlSi3O8– Na2CO3–H2O and its bearing on the genesis of carbonatites. Am. J. Sci. 266, 932–967.

JAGO, B. & GITTINS, J. (1990): The role of fluorine in carbonatite magma evolution. Nature 349, 56–58. 146


KRASNOVA, N.I., BALAGANSKAYA, E.G. & GARCIA, D. (2004): Kovdor – classic phoscorites and carbonatites. In Phoscorites and Carbonatites from Mantle to Mine: the Key example of the Kola Alkaline Province (F. Wall & A.N. Zaitsev, eds.), Mineralogical Society Series 10, 99–132.

MITCHELL, R.H. (1997): Carbonate-carbonate immiscibility, neighborite and potassium iron sulphide in Oldoinyo Lengai natrocarbonatite. Mineral. Mag. 61, 779–789. MITCHELL, R.H. (2005a): Mineralogical and experimental constraints of the origins of niobium mineralization in carbonatites. In Rare Element Geochemistry and Mineral Deposits (R.L. Linnen, and I.M. Samson, eds.). Geological Association of Canada Short Course 17, 201–215.

LARSEN L.M. & REX, D.C. (1992): A review of the 2500 Ma span of alkaline-ultramafic, potassic and carbonatitic magmatism in west Greenland. Lithos 28, 367–402. LE BAS, M.J. (1977): Carbonatite–Nepheline Volcanism. Wiley, New York.

MITCHELL, R.H. (2005b): Carbonatites and carbonatites and carbonatites. Can. Mineral. 43, 2049–2068.

LE BAS, M.J. & ASPDEN, J.A. (1981): The comparability of carbonatitic fluid inclusions in ijolites with natrocarbonatite lava. Bull. Volcanol. 44, 429–438.

MITCHELL, R.H., SCOTT SMITH, B.H. & LARSEN, L.M. (1999):Mineralogy of ultramafic dikes from the Sarfartoq, Sisimiut and Maniitsoq areas, West Greenland. In Proc. Seventh Int. Kimberlite Conf. 2. Red Roof Designs, Cape Town, South Africa, 574–583.

LE MAITRE, R.W., compiler (2002): Igneous Rocks: a Classification and Glossary of Terms. Cambridge University Press, Cambridge, UK.

NESBITT, B.E. & KELLY, W.C. (1977): Magmatic and hydrothermal inclusions in carbonatite of the Magnet Cove Complex, Arkansas. Contrib Mineral. Petrol. 63, 271–294.

LEE, W.-J. & WYLLIE, P.J. (1998): Processes of crustal carbonatite formation by liquid immiscibility and differentiation, elucidated by model systems. J. Petrol. 39, 2005–2013.

NIELSEN, T.F.D. (1980): The petrology of melilitolite, melteigite, carbonatite and syenites ring dyke system in the Gardiner complex, East Greenland. Lithos 13, 181–197.

LEE, W.-J., WYLLIE, P.J. & ROSSMAN, G.R. (1994): CO2-rich glass, rounded calcite crystals, and no liquid immiscibility in the system CaO–SiO2– CO2. Am. Mineral. 79, 1135–1144.

NIELSEN, T.F.D. (1981): The ultramafic cumulate series, Gardiner complex, East Greenland. Contrib. Mineral. Petrol. 76, 60–72.

LENTZ, D. R. (1998): Late-tectonic U–Th–Mo–REE skarn and carbonatitic vein-dyke systems in the southwestern Grenville Province: A pegmatiterelated pneumatolytic model linked to marble melting. In Mineralized Intrusion-related Skarn Systems (Lentz, D.R., ed.) Min. Assoc. Can. Short Course 26, 519-657.

NIELSEN, T.F.D., SOLOVOVA, I.P. & VEKSLER, I.V. (1997): Parental melts of melilitolite and origin of alkaline carbonatite: evidence from crystallized melt inclusions, Gardiner complex. Contrib. Mineral Petrol. 126, 331–344.

LENTZ, D.R. (1999): Carbonatite genesis: a reexamination of the role of intrusion-related pneumatolytic skarn processes in limestone melting. Geology 27, 335–338.

NIELSEN, T.F.D. & VEKSLER, I.V. (2002): Is natrocarbonatite a cognate fluid condensate? Contrib. Mineral. Petrology 142, 425–435.

MARIANO, A.N. (1989): Nature of economic mineralization in carbonatites and related rocks. In Carbonatites: Genesis and Evolution (K. Bell, ed.). Unwin Hyman, London, 149–176.

OTTO, J.W. & WYLLIE, P.J. (1993): Relationships between silicate melts and carbonate-precipitating melts in CaO–MgO–SiO2 –CO2 –H2O at 2 kbar. Mineral. Petrol. 48, 343–365.

MELCHER, G.C. (1966): The carbonatite of Jacupiranga, São Paulo, Brazil. In Carbonatites. (Tuttle, O.F. & Gittins, J. eds.). Interscience, New York, 169–181.

PANINA, L.I. (2005): Multiphase carbonate–salt immiscibility in carbonate melts: data on melt inclusions from the Krestovskiy massif minerals (Polar Siberia). Contrib. Mineral. Petrology 150, 19–36. 147


RANKIN, A.H. (1975): Fluid inclusion studies in apatite from carbonatites of the Wasaki area of western Kenya. Lithos 8, 123–136.

SHAND, S.J. (1922): The alkaline rocks of the Transport Line, Pretoria District. Geol. Soc. S. Africa Trans. 25, 85.

RANKIN, A.H. (1977): Fluid-inclusion evidence for the formation conditions of apatite from the Tororo carbonatite complex of eastern Uganda. Mineral. Mag. 41, 155–164.

SHANNON, R.D. (1976): Revised effective ionic radii and systematic studies of interatomic distances in halides and chalcogenides. Acta Crystallog. A32, 751–767.

RANKIN, A.H. (2005): Carbonatite-associated rare metal deposits: Composition and Evolution of OreForming Fluids – The Fluid Inclusion Evidence. In Rare-Element Geochemistry and Mineral Deposits (R.L. Linnen & I.M. Samson, eds.). Geological Association of Canada Short Course 17, 299–314.

SOKOLOV, S. (2002): Melt inclusions as indicators of the magmatic origin of carbonatite rare metal and rare earth minerals. Chem. Geol. 183, 373– 378. SWEENEY, R.J., FALLOON, T.J. & GREEN, D.H. (1995): Experimental constraints on the possible mantle origin of natrocarbonatite. In Carbonatite Volcanism. Oldoinyo Lengai and Petrogenesis of Natrocarbonatites (K. Bell & J. Keller, eds.) Springer, Berlin, Heidelberg, New York, 191–207.

RANKIN, A.H. & LE BAS, M.J. (1974): Liquid immiscibility between silicate and carbonate melts in naturally occuring ijolite magma. Nature 250, 206–209.

TAPPE, S., JENNER, G.A., FOLEY, S.F., HEAMAN, L., BESSERER, D., KJARSGAARD, B.A. & RYAN, B. (2004): Torngat ultramafic lamprophyres and their relation to North Atlantic alkaline province. Lithos 76, 491–518.

RASS, I.T. & PLECHOV, P.Y. (2000): Melt inclusions in olivines from the olivine-melilitite rock of the Guli Massif, northwestern Siberian platform. Doklady Earth Sci, 275A, 1399–1402. RAVICH, M.I. (1974): Water–Salt Systems at Elevated Temperatures and Pressures. Nauka, Moscow, (in Russian).

THIBAULT, Y., EDGAR, A.D. & LLOYD, F.E. (1992): Experimental investigation of melts from a carbonated phlogopite lherzolite: implications for metasomatism in the continental lithospheric mantle. Am. Mineral. 77, 784–794.

ROEDDER, E. (1973): Fluid inclusions from the fluorite deposits associated with carbonatite at Amba Dongar, India, and Okorusu, South west Africa: metasomatic feldspar rocks (potash fenites) associated with the fluorite deposits and carbonatites of Amba Dongar, Gujarat, India: discussion and contributions. Trans. Inst. Mining & Metallurgy B-82, B35-B39.

VAN ACHTERBERGH, E., GRIFFIN, W.L., RYAN, C.G., O’REILLY, S.Y., PEARSON, N.J., KIVI, K. & DOYLE, B.J. (2002): Subduction signature for quenched carbonatites from the deep lithosphere. Geology 30, 743–746.

ROEDDER, E. (1984): Fluid Inclusions. In (Ribbe, P.H., ed.), Reviews in Mineralogy 12, Washington, D.C., 644 p.

VAN ACHTERBERGH, E., GRIFFIN, W.L., RYAN, C.G., O’REILLY, S.Y., PEARSON, N.J., KIVI, K. & DOYLE, B.J. (2004): Melt inclusions from the deep Slave lithosphere: implications for the origin and evolution of mantle-derived carbonatite and kimberlite. Lithos 76, 461–474.

ROMANCHEV, B.P. & SOKOLOV, S.V. (1980): Liquation in the production and geochemistry of rocks in carbonatite complexes. Geochem. Intern. 16, 125–135.

VEKSLER, I.V. (2004): Liquid immiscibility and its role at the magmatic–hydrothermal transition: a summary of experimental studies. Chem. Geol. 210, 7–31.

SCHAIRER, J.F., YAGI, K. & YODER, H.S., JR. (1962): The system nepheline – diopside. Carnegie Inst. Wash. Yearbook 61, 96–98.

VEKSLER, I.V. & KEPPLER, H. (2000): Partitioning of Mg, Ca, and Na between carbonatite melt and hydrous fluid at 0.1–0.2 GPa. Contrib. Mineral. Petrology 138, 27–34.

SEIFERT, W. & THOMAS, R. (1995): Silicate– carbonate liquid immiscibility: a melt inclusion study of olivine melilitite and wehrlite xenoliths in tephrite from the Elbe Zone, Germany. Chem. Erde 55, 263–279. 148


WYLLIE, P.J. (1974): Limestone assimilation. In The Alkaline Rocks (H. Sørensen, ed.). John Wiley & Sons, New York, 459–474.

VEKSLER, I.V., NIELSEN, T.F. & SOKOLOV, S.V. (1998a) Mineralogy of crystallized melt inclusions from Gardiner and Kovdor ultramafic alkaline complexes: implications for carbonatite genesis. J. Petrol. 39, 2015–2031.

WYLLIE, P.J. (1989): Origin of carbonatites: Evidence from phase equilibrium studies. In Carbonatites: Genesis and Evolution (K. Bell, ed.). Unwin Hyman, Boston, 500–545.

VEKSLER, I.V., PETIBON, C., JENNER, G.A., DORFMAN, A.M. & DINGWELL, D.B. (1998b): Trace element partitioning in immiscible silicate– carbonate liquid systems: an initial experimental study using a centrifuge autoclave. J. Petrol. 39, 2095–2104.

WYLLIE, P.J. & HUANG, W.-L. (1976): Carbonation and melting reactions in the system CaO–MgO– SiO2–CO2 at mantle pressures with geophysical and petrological applications. Contrib. Mineral. Petrol. 54, 79–107.

VON ECKERMANN, H. (1948): The genesis of the Alnö alkaline rocks. Int. Geol. Congr., 18th Session, Great Britain 3, 94–101.

WYLLIE, P.J. & LEE, W-J. (1998): Model system controls on conditions for formation of magnesiocarbonatite and calciocarbonatite magmas from the mantle. J. Petrol. 39, 1885– 1993.

WALLACE, M.E. & GREEN, D.H. (1988): Experimental determination of primary carbonatite magma composition. Nature 335, 343–346.

WYLLIE, P.J. & TUTTLE, O.F. (1959): Melting calcite in the presence of water. Am. Mineral. 44, 453–459.

WATKINSON, D.H. & WYLLIE, P.J. (1969): Phase equilibrium studies bearing on the limestone assimilation hypothesis. Geol. Soc. Am. Bull. 80, 1565–1576.

WYLLIE, P.J., TUTTLE, O.F. (1960): The system CaO–CO2–H O and the origin of carbonatites. J. Petrol. 1, 1–46.

WATKINSON, D.H. & WYLLIE, P.J. (1971): Experimental study of the join NaAlSO4–CaCO3– H2O and the genesis of alkalic rock–carbonatite complexes. J. Petrol. 12, 357–378.

YEVZIKOVA, N.Z. & MOSKALYUK, A.A. (1964): Gas–liquid inclusions in carbonatite carbonates. Acad. Sci. U.S.S.R. Doklady Earth Sci. 159, 108– 111.

WEBSTER, J.D. & DE VIVO, B. (2002): Experimental and modeled solubilities of chlorine in aluminosilicate melts, consequences for magma evolution, and implications for exsolution of hydrous chlorine melt at Mt. Somma–Vesuvius. Am. Mineral. 87, 1046–1061. WOOLLEY, A.R. (1987): Alkaline Rocks and Carbonatites of the World. Part 1: North and South America. British Museum Natural History, London. WOOLLEY, A.R. (2001): Alkaline Rocks and Carbonatites of the World. Part 3: Africa. The Geological Society, London. WOOLLEY, A.R. (2003): Igneous silicate rocks associated with carbonatites: their diversity, relative abundances, and implications for carbonatite genesis. Per. Mineral. 72, 9–17. WOOLLEY, A.R. & KEMPE, P.R.C. (1989): Carbonatites: nomenclature, average chemical composition and element distribution. In Carbonatites: Genesis and Evolution (Bell, K., ed.), London, Unwin Hyman, p. 1–14. 149


150

CHAPTER 7: MAGMATIC PROCESSES AND VOLATILE PHASE GENERATION IN PORPHYRYTYPE ENVIRONMENTS: A LASER ABLATION–ICP–MS STUDY OF SILICATE AND SULFIDE MELT INCLUSIONS

Werner E. Halter and Christoph A. Heinrich Isotope Geochemistry and Mineral Resources Department of Earth Sciences, ETH Zürich 8092 Zürich, Switzerland E-mail: [email protected] of the magma chamber, and to find out which processes lead to the formation of silica-rich magmas and volatile saturation. Our results are published in a series of six papers that report the discovery of sulfide melt inclusions (Halter et al. 2002a) and successively established the micro-analytical technique (Halter et al. 2004c, Halter et al. 2002b), the evolution of the magmatic complex based on field mapping, wholerock chemical and isotopic analysis and 40Ar/39Ar geochronology by stepwise laser heating (Halter et al. 2004a), the origin of andesitic magmas through mixing of basalt and rhyodacite (Halter et al. 2004b), and the origin of the ore fluid (Halter et al. 2005). The combination of geological constraints on the magmatic evolution of the complex with the mass balance and fluid–chemical requirements imposed by the formation of the 600 Mt ore deposit permits a detailed reconstruction of the timing of volatile saturation, the consequences of sulfide melt saturation, the source for the hydrothermal fluid and the chronology of the ore-forming events. Here, we use this geological and geochemical framework to determine the source magmas for fluids, ore metals, and sulfur, and to trace the behavior of Cu and Au during magma evolution. Our results contrast with current views of the processes responsible for orefluid formation, which generally assume that porphyry-mineralizing magmas have to be oxidized (see Rowins 2000) to prevent sulfide saturation and permit metal enrichment. The resulting model has the potential of practical application to regional mineral exploration.

INTRODUCTION Large porphyry-type hydrothermal systems are generally associated with silica-rich subvolcanic intrusions that form in several cycles of intrusion and alteration or mineralization (Gustafson & Hunt 1975, Proffett 2003). Due to this close temporal association, the link between the magmatic and the hydrothermal event is commonly accepted and intrusions hosting the mineralization have been investigated intensively. In particular, several studies have focused on reconstructing the evolution of the magmatic system that generated the porphyritic intrusions (e.g., Dilles & Proffett 1995, Halter et al. 2005, Hattori & Keith 2001, Keith et al. 1997, Kesler 1997, Landtwing et al. 2002, Sillitoe 1973). A powerful approach to do this involves the study of melt inclusions, trapped at depth during crystal growth. However, only few such studies have been conducted, mostly because these melts are rich in volatiles and, thus, difficult to homogenize (Student & Bodnar 2004). Homogenization under pressure is possible in quartz-hosted inclusions (Schmitt et al. 2002), but in minerals depicting a strong cleavage, successful homogenization can be quite difficult. The only analytical technique that does not require previous homogenization of the inclusions is laser ablation–ICP–MS. Analyses are done by ablation of entire inclusions below the sample surface. As part of the host mineral is ablated with the inclusion, quantification of the melt composition requires that the appropriate amount of material from the host mineral is subtracted from the analysis. This is done using the composition of the host phase, obtained within the same analysis, and an internal standard. We used this approach on samples from a large andesitic complex that hosts a world-class porphyry Cu–Au deposit to reconstruct the history of the complex, to identify the evolution

Melt inclusions in porphyry systems Studies that investigated melt inclusions in porphyry systems mainly focused on inclusions in quartz as they are the best candidates for a successful homogenization during microthermo-

Mineralogical Association of Canada Short Course 36, Montreal, Quebec, p. 151-164.

151

WERNER E. HALTER & CHRISTOPH A. HEINRICH

metric experiments. Due to their high water contents, inclusions in other phases generally decrepitate before the formation temperature is reached (Student & Bodnar 2004). Audétat & Pettke (2003) used quartz-hosted melt inclusions to quantify the element enrichment during crystallization of granitic systems, and their results indicate that very high degrees of fractionation and late volatile saturation need to be reached to increase significantly the Cu content in the residual melt. High Cu concentrations in silicate melt inclusions have been measured in extremely fractionated melts (Harris et al. 2003), but there is no evidence that sufficiently large quantities of such melts are associated with porphyry-type mineralization to explain the formation of large deposits. Recently, Student & Bodnar (2004) measured Cu concentrations of a few hundreds of ppm in melt inclusions of intermediate silica content in quartz phenocrysts from the Tyrone and Red Mountain porphyries. Entrapped melts with higher degrees of fractionation have lower Cu concentrations, and the authors suggested that this is consistent with the exsolution of a Cu-rich fluid from the melt. Thus, the highest Cu contents in the melt are not due to a progressive increase during fractional crystallization. Quartz is formed only in silica-rich magmas, and melt inclusions therein might not provide a record that spans the entire evolution of the system and a potential mechanism that enriches Cu in the melt prior to volatile saturation. Therefore, we extended the investigations to melt inclusions trapped in other phases formed earlier in the magma chamber. Below we present the results from the ore-forming Farallón Negro andesitic complex, Argentina, in which silicate and sulfide melt inclusions were analyzed in amphibole, pyroxene, plagioclase and quartz using laser ablation–ICP–MS.

vertical topography of the edifice (Ulrich 1999) and exposed the deep, internal parts of the volcano. The present study focused on the largest and bestexposed part of the complex, which records most of the intrusive rocks and hydrothermal events. Most of the volcanic pile could be sampled in sequence from the interior to the rim of the volcanic complex. 40 Ar/39Ar stepwise laser heating geochronology is consistent with this interpretation and allowed absolute timing of subvolcanic intrusions within this sequence (Halter et al. 2004a). Note that hereafter, volcanic rocks are called "extrusive" and subvolcanic stocks are called "intrusive" to distinguish their mode of emplacement. However, as no "intrusion" is really plutonic and the texture of extrusive rocks and subvolcanic stocks are very similar, we used the same terminology (from volcanic rocks) to describe their compositions. Emplacement of intrusive and extrusive rocks Volcanic rocks are mainly high-potassium calcalkaline basalt, basaltic andesite and andesite, containing 45 to 66 wt.% SiO2 and phenocryst assemblages of amphibole (hornblende) + plagioclase + magnetite ± pyroxene. Andesite is by far the most abundant extrusive rock type. Biotite-bearing volcanic rocks appear only after 8.0 Ma. The main volcanic activity ceased at 7.5 Ma with the crystallization of the equigranular Alto de la Blenda stock, interpreted to be the central volcanic conduit (Llambías 1972). The only younger extrusive rocks are quartz-bearing dacitic ignimbrite extruded in a flank eruption at 7.35 Ma. The coeval Agua Tapada stock is the likely intrusive feeder, as indicated by dykes connecting it to the ignimbrite. Subvolcanic stocks were first emplaced at 9.0 Ma and become abundant after 8.5 Ma. Individual stocks consist of up to eight intrusions, emplaced within a restricted time frame, but comprising a wide range in compositions. Stocks are mostly dacitic but range from andesitic to rhyolitic and contain amphibole + plagioclase + magnetite ± pyroxene or plagioclase + quartz + magnetite ± amphibole ± biotite as the main phenocryst phases. Pyroxene is restricted to stocks prior to 7.5 Ma and represents at most 10 vol. % of phenocrysts. Biotite appears as a major phenocryst phase in stocks after 7.5 Ma. Amphibole is ubiquitous in all fresh intrusive stocks, except in the latest and most silica-rich ones dominated by biotite.

Farallón Negro Volcanic complex Geological setting The Farallón Negro Volcanic Complex (FNVC) is located in northwestern Argentina, ~200 km east of the present Andean volcanic arc and 500 km from the trench. Previous mapping and analysis of this system (Llambías 1970, Sasso 1997, Proffett 2003, Halter et al. 2004a) have shown that the complex represents the remnants of a 20 km wide and 4.5 km high stratovolcano (Sillitoe 1973, Halter et al. 2004a) intruded by numerous subvolcanic stocks (Fig. 7-1). Extensive erosion removed over 3.5 km of the

152


Tertiary Igneous Rocks

27° 10'

A

Volcanic Rocks Basalt to dacite

Campo de Arenal

Porphyries Rhyolite Rhyodacite and dacite Andesite Basaltic andesite

Cerro Durazno

Tertiary Sediments Continental and marine

27° 20'

Basement Farallón Negro lita

s

Paleozoic Granite Metagreywacke Migmatite and gneiss

pil Ca 10°

27° 30'

Overburden Known fault Inferred fault Reverse fault

Agua Rica

Neogene foreland thrust belt

Subduction angle 5-10°

modern magmatic arc

66° 40'

66° 30'

20

Loma Morada

23

25

Inf ere d

fa

Ca ld eFaurtl

ul

t

Agua Tapada

na Pu

Farallon Negro Volcanic Complex volcanoes

Porphyries Rhyolite Macho Muerto rhyodacite Agua Tapada dacite Alto de la Blenda monzonite Chilca andesite Durazno basaltic andesite Tampa Tampa andesite Basaltic andesite dyke Andesitic dyke

Inf er re d

ca

l de

ra

6 985 000

hile Tre nch

500 km

Volcanic rocks Gravel Quartz dacite Basement Biotite andesite Veins Hornblende andesite Bedding Andesite (undiff.) Pyroxenehornblende andesite 2 km

B

Peru-C

27° 40'66° 50'

5 km

Tucuman Transform Zone

Sierras nas Pampea

0

Subduction angle 5-10°

no

30°

NAZCA PLATE

Subduction angle ~30°

la Altip

20°

Macho Muerto Tampa Tampa

30 21

Alto de la Blenda

Leo

6 980 000

Bajo el Durazno

nes

Las Pampitas

30 fa

ul

t

Chilca

Bajo de la Alumbrera Casitas

27 25

3 440 000

3 435 000

3 430 000

6 975 000

FIG. 7-1. A. Geological setting of the Farallón Negro Volcanic Complex. The volcanic sequence covered some 700 km2 and was variably uplifted through reverse faults during late compression (modified from Martinez et al. 1995 and Sasso 1997). B. Detailed geological map of the northwestern part of the complex. Erosion of some 3.5 km of the complex exposed the various intrusions and volcanic rocks erupted over most of the magmatic activity. Modified from Llambías (1970).

153


approximately 15 km3 of mafic magma (Halter et al. 2005).

Some 1.2 m.y. after initiation of extrusive activity, a continuous compositional evolution of subvolcanic intrusions with time is taken as evidence for the gradual build-up and growth of a subvolcanic magma chamber. Heat is provided by continued advection of magma through the center of the volcano (Halter et al. 2004b). Cessation of the main volcanic activity at 7.5 Ma was probably caused by the end of magma supply to the chamber, concurrently ending the heat supply and forcing the system to crystallize progressively. Following this event, only dacitic and more silica-rich intrusive rocks were emplaced. A change in the tectonic regime from locally extensive (Sasso & Clark 1998) to dominantly transpressive probably contributed to the termination of extrusive activity, and constrained the magma chamber to evolve internally towards bulk fluid saturation (Halter et al. 2005). The NE–SW alignment of intrusions emplaced after 7.5 Ma corresponds to a fracture zone that controlled the ascent and spatial emplacement of volatile-saturated magmas.

Analytical approach for melt inclusions Melt inclusions in andesitic systems have been analyzed with various techniques, most of which require preliminary homogenization and exposure to the sample surface (see Pettke 2006). In this contribution, we focus on analyses of heterogeneous inclusions by laser ablation– inductively coupled plasma–mass spectrometry (LA–ICP–MS). Melt compositions and their uncertainties were quantified following the procedure described in Halter et al. (2002b). This procedure is reviewed below. Since multiphase inclusions are not homogenized to a glass, critical data evaluation has to be applied to identify inclusions that are not representative of a melt because of heterogeneous entrapment, because of back reactions between the host mineral and the melt or because they behaved as open systems after entrapment. This is done by analyzing several inclusions in a given assemblage of simultaneously trapped inclusions (e.g., on a same growth zone, Fig. 7-2). Analyses are considered representative of the melt only if they yield consistent results within an inclusion assemblage. Individual inclusions were not analyzed and inclusions which differed significantly from other inclusions of the same assemblage were not retained.

The Alumbrera porphyry The Bajo de la Alumbrera porphyry Cu–Au deposit formed between 7.1 and 6.8 Ma, toward the end of the magmatic activity. Most of the stocks formed prior to the Alumbrera deposit show some hydrothermal alteration and Cu and Au enrichment to various extents, but only Alumbrera shows economic metal concentrations. It formed as a series of eight intrusions (Proffett 2003) showing a general evolution toward more mafic rocks, from early dacite to late andesite (Ulrich & Heinrich 2001, Proffett 2003). This evolution of intrusion composition within the Alumbrera stock is taken as evidence for a chemically structured magma chamber, with initial extraction of silica-rich melts from the top and progressively more mafic melt from deeper parts of the chamber (Halter et al. 2004b). Emplacement of intrusions was probably caused by the buoyancy of silica- and volatile-rich magma in the roof zone of the magma chamber, together with structural focusing. Following a barren intrusion that predated the porphyry stock, Cu–Au mineralization is associated with each magma pulse up to the fourth intrusion, with a general trend toward less intense alteration and mineralization with each pulse. Post-mineralization andesitic dykes that cut the stock are unaltered and barren, but they carry sulfide melt inclusions. The total amount of Cu deposited in the porphyry stock requires a minimum source melt volume of

Analytical procedure Only entire, unexposed melt inclusions 5 to 30 µm beneath the sample surface were considered for analysis. Analysis of only parts of heterogeneous inclusions would yield a nonrepresentative composition of the melt. Inclusions were ablated using an optically homogenized UV beam from a 193 nm ArF Eximer laser (Gunther et al. 1997). Details of run conditions are reported in Halter et al. (2002b). The size of the laser spot can be adjusted between 8 and 80 µm using pinhole apertures. For each inclusion, the analytical pit size was selected to be slightly larger than the inclusion diameter. The ablated material was transported by He carrier gas to an Elan 6100 quadrupole mass spectrometer, which sequentially recorded signals for all the elements of interest. Ablation was monitored through an optical microscope and the recorded signal was displayed in real time on a computer monitor to assess the control of the ablation procedure (Fig. 7-3). Each transient signal is composed of 20 to 30 s of instrument background

154


FIG. 7-2. A to C Partially crystallized silicate melt inclusions along growth zones in a plagioclase crystal. Due to partial or complete crystallization of inclusions, heterogeneous entrapment cannot be recognized petrographically if inclusions are not rehomogenized. Thus analyses were considered representative of a melt only if several inclusions in the same assemblage (e.g., along a same growth zone) yielded the same compositions. Sulfide melt inclusion in amphibole in transmitted (D) and combined transmitted and reflected light (E).

measurement, followed by the analytical signal from the ablation of the host mineral and the inclusion. The analytical signal was composed of a steady response from the ablation of the host mineral before and after the inclusion, and an intermediate transient signal from the combined ablation of host and inclusion in varying proportions. Details of the analytical set-up are given in Gunther et al. (1997, 1998) and Heinrich et al. (2003), and further details are presented in Pettke (2006). Sulfide melts were analyzed using the same approach as described for silicate melt inclusions (Halter et al. 2004c). As for silicate inclusions, the signal consists of a background measurement, monitored prior to ablation, a signal from the host before and after the ablation of the inclusion and an evolving signal reflecting the mixture of host plus inclusion.

Quantification of melt inclusion compositions Quantification of the chemical composition of silicate melt inclusions is obtained through a threestep procedure. First, analytical signals of the host and the host plus inclusion mixture are converted into element ratios using element sensitivity factors determined through external standards analyzed before and after the unknown (SRM–610 from NIST, hereafter NIST 610). Second, element concentrations in the host and the mixed signal are calculated by normalizing them to a total concentration of 100 wt.% major elements on an anhydrous basis (or less if a certain water concentration is expected). Finally, the relative contributions from the host and the inclusion to each element concentration in the mixed host + inclusion signal needs to be assessed, to determine element concentrations in the originally trapped melt. This is done for each inclusion using an internal standard (Halter et al. 2002b). An internal

155

Concentration

WERNER E. HALTER & CHRISTOPH A. HEINRICH

C iHOST C iMIX

Extr

apo

latio

n

CiINCL

0

0.2 X 0.4

0.6

0.8

1

Mass inclusion

(mass inclusion + host)

Host 7

Laser on

Host

Host + inclusion

Laser off 27 23

Log (counts/second)

6

Al Na

39

K Si 88 Sr 29

5 Background 57

4 86 232

3

Th

Fe

Rb

2 0 20

40

60 Time (seconds)

80

100

FIG. 7-3. Typical laser ablation–ICP–MS signal for various elements during ablation of inclusions in a plagioclase host and schematic representation of the procedure to quantify melt inclusion compositions. CiHOST, CiMIX and CiINCL are the concentrations of an element, i, in the host, the host + inclusion mixture and the inclusion, respectively. The signal consists of an initial signal from the host mineral, followed by an evolving signal of host + inclusion mixture and then again from the host only, after the inclusion is completely ablated. Quantification is done by using the relative proportion, X, of [inclusion/(inclusion + host)] in the mixed signal determined using an internal standard (e.g., an element for which the concentration CiINCL in the inclusion is know). Element concentrations in the inclusion, CiINCL, are calculated by removing the host contribution from the mixture, i.e., by extrapolating from CiHOST and CiMIX to an inclusion/(inclusion + host) ratio of 1. standard is a known element concentration in the melt, by which the mass ratio between the host and the inclusion can be uniquely determined (Halter et al. 2002b). In this study, we took advantage of the essentially constant concentration of Al2O3 in the bulk rock and used the well-defined correlation between Al2O3 and FeO in bulk rocks to constrain the Al2O3 content of the melt inclusions. Once the inclusion/host ratio, X, is known, the concentration of all the elements in the melt can be quantified through a simple mathematical procedure, graphically presented in figure 7-3. This approach also corrects for any quantity of host mineral crystallized from the melt onto the inclusion wall.

Thus, remelting of the inclusion wall is not necessary with this approach. The composition of the various host phases obviously varies significantly during magma evolution, but since we always use the host analyzed just above or below the inclusion, this has no effect on the quantification of the inclusion composition. There is no evidence for significant compositional changes due to small scale zonation, which could generate differences in the host composition at the scale of a few tens of micrometres (i.e., the vertical distance between the analysis of the inclusion and the host). Important to note is that uncertainties of individual element analysis can be calculated from 156


most elements (Fig. 7-4), except for small but diagnostic deviations observed for some elements at particular stages of the evolution. Inclusions depict a dominantly bimodal distribution with either basaltic or rhyodacitic compositions, in strong contrast to the predominantly andesitic bulk rock compositions. Moreover, various phenocrysts contain melt inclusions with contrasting compositions. The mafic component is preserved as melt inclusions with 42 to 45 wt.% SiO2 in amphibole. Plagioclase dominantly crystallized from the felsic magma prior to mixing, as rhyodacitic melt inclusions are present almost exclusively in this phase. This shows that phenocryst minerals observed in a given rock crystallized from different melts, and are in fact xenocrysts in the bulk rock. These observations and linear compositional variations of intrusive and extrusive rocks are best explained by a process of mixing varying proportions of a mafic and a felsic (rhyodacitic) magma that already contained phenocrysts. Both magmas may have been at similar temperatures and viscosities, thanks to the high H2O content of the mafic magma, thus favoring complete hybridization of the melts. Intrusions and extrusions have characteristic differences, as evidenced by a systematic distinction in the compositions of intermediate melt inclusions and in their occurrence in the various phases. In volcanic rocks, intermediate melts generated upon mixing are mainly recorded in plagioclase and pyroxene. Amphibole is partially resorbed, indicating a shift from an amphibole (±plagioclase) stable assemblage to the plagioclase + pyroxene stability field. This is explained by decompression during the ascent of the magma towards the surface. Sulfide melt inclusions are generally absent from any phenocrysts in the extrusive andesite, suggesting that these magmas never exsolved a sulfide melt on a broad scale. Rapid degassing upon magma mixing likely suppressed the formation of a sulfide melt (Keith et al. 1997) and induced eruption selectively at times when volatile exsolution occurred or slightly thereafter. In intrusive rocks, melts of intermediate compositions are trapped preferentially by amphibole. Pyroxene with melt inclusions of intermediate composition is only present in the Alto de la Blenda stock, which is the inferred volcanic conduit. The coexistence of amphibole and plagioclase and the absence of pyroxene indicate high water contents and suggest that the source

the LA–ICP–MS signal. These uncertainties increase with decreasing signal intensity and duration and, more significantly, with increasing concentration of the element in the host phase. Typical values are between a few weight percent for incompatible elements and a few tenths of a weight percent for elements abundant in the host mineral. Details of the uncertainty calculation are given in Halter et al. (2002b), and uncertainties from the data used in this study are in Halter et al. (2004b). Sulfide melt inclusions were quantified using the same approach and the same external standard as for silicate melt inclusions, i.e., a silicate glass NIST 610. This is possible because ablation with the Eximer 193 nm laser and the homogenized beam profile induces no matrixdependent fractionation (Halter et al. 2004c). Sulfur was not quantified and element concentrations were obtained by assuming that the inclusion was stoichiometric (Fe,Cu)S. Copper concentrations were sufficiently small (max. 3 wt.%) that the valence and stoichiometric states of Cu are irrelevant within the analytical uncertainty (5–10 wt.%). The mass ratio between the inclusion and the host was calculated by assuming that the inclusion contained no silica, i.e., element contributions from the host mineral (mostly Fe from amphibole in this case) were subtracted in proportion of the SiO2 content of the host mineral. To increase the number of determinations and the counting time on Au, the number of elements analyzed was reduced to four (Si, Fe, Cu, Au). This significantly decreased the uncertainty and the limit of detection for Au. Moreover, it granted a more representative sampling (Pettke et al. 2000) of the Au signal by increasing the number of measurements over the short time interval over which tiny nuggets are ablated. Silicate and sulfide melt inclusion analyses Results for approximately 200 individual melt inclusions in 19 samples are presented graphically for selected major and trace elements in Figure 7-4 in comparison with bulk-rock compositions. The full data set is provided in Halter et al. (2004b). Melt inclusions are identified with respect to the bulk rock composition and the host mineral phase in which they occur. Silicate melt inclusions have SiO2 contents between 42 and 75wt. %, covering a similar compositional range as that of the bulk rocks. Changes in the composition of melt inclusions with increasing SiO2 content and other inter-element correlation trends follow those of bulk rocks for

157


4 3.5

Volcanic rocks

Intrusive rocks

4 3.5

TiO2 (wt.%)

3

3

2.5

2.5

2

2

1.5

1.5

1

1

0.5

0.5

0 10 9 8 7 6 5 4 3 2 1 0 18

0 10 9 8 7 6 5 4 3 2 1 0 18

K2O (wt.%)

16

16

CaO (wt.%)

14

14

12

12

10

10

8

8

6

6

4

4

2

2

0 2000 1800 1600 1400 1200 1000 800 600 400 200 0 10000

0 2000 1800 1600 1400 1200 1000 800 600 400 200 0 10000

Ba (ppm)

Cu (ppm) 1000

1000

100

100

10

10

1 40

45

50

55

60

65

70

75

80

1 40

45

50

Host Mineral amphibole

60

65

70

80

Bulk rock

Host rock plagioclase

75

SiO2 (wt.%)

SiO2 (wt.%) pyroxene

55

quartz

basaltic andesite

andesite

dacite

rhyolite

FIG. 7-4. Variation in major- and trace-element concentrations in melt inclusions and bulk rocks as a function of the SiO2 content. Symbols identify host minerals; grey shades represent the host rock. Linear changes, a bimodal distribution of mostly mafic and silica-rich inclusions as well as contrasting melt inclusions in various phases of the same host rock suggest that most rock types result from mixing between a very mafic and a rhyodacitic magma. Most element variations in melt inclusions follow the trend determined by bulk rocks. Only Cu shows deviation in melt inclusions from bulk rocks with a 2- to 5-fold enrichment in volcanic rocks and a strong depletion in intrusive rocks. Cu-poor inclusions of intermediate compositions in intrusive rocks are associated with sulfide melt inclusions.

158


magmas for intrusions were not actively degassed during mixing. This is consistent with the common association of sulfide melt inclusions with silicate melt inclusions of intermediate composition in amphibole phenocrysts of subvolcanic intrusions. Degassing would destabilize sulfides through the loss of sulfur to the volatile phase and suppress sulfide saturation.

could be measured in the largest inclusions, where limits of detection are as low as 0.1 ppm (Fig. 7-5), and the concentrations range from 0.1 to 4.6 ppm. Early high-temperature fluid inclusions, characterized texturally and associated with the oreforming event in the Alumbrera deposit, are hightemperature brines (salinity ~50% equivalent NaCl, entrapped in pre-ore vein quartz at T >700°C, Ulrich et al. 2001) containing Cu and Au in the same concentration ratio as the bulk ore (fig. 5). The same Cu/Au ratio is also recorded by the sulfide melt inclusions in barren porphyry intrusions of the Farallón Negro Volcanic Complex (Halter et al. 2002a) and, in particular, in the late, unaltered andesitic intrusions of the Alumbrera stock. The combined textural and metal ratio data, therefore, indicate that the magmatic sulfides and the ore fluids share a common magmatic source.

Copper and gold contents of silicate and sulfide melt inclusions In volcanic rocks, the abundance of Cu in silicate melt inclusions displays a very large variation, sharply contrasting with the simple mixing trend of all other elements that closely mimic the composition of bulk rocks (Halter et al. 2004b, Fig. 7-4). In further contrast with other elements, Cu is almost systematically more abundant in silicate melt inclusions than in bulk rocks with the same silica content. In particular, Cu contents of the most primitive melts are approximately 100 to 200 ppm, corresponding to 2–5 times the bulk Cu concentration in mafic rocks. In silica-rich melts, Cu concentrations of 20 to 50 ppm in inclusions represent a similar enrichment factor over bulk rocks. In intrusive rocks, melt inclusions of intermediate composition (50 to 65 wt.% SiO2) are mostly trapped in amphibole and have very low Cu contents of a few ppm. Most Cu concentrations are indeed below the limit of detection of 2 to 3 ppm. Such inclusions are always associated with small (

Mineralogical Association of Canada Short Course, 36 MELT INCLUSIONS IN PLUTONIC ROCKS

Mineralogical Association of Canada Short Course, 40 Laser Ablation–ICP–MS in the Earth Sciences

WORKING WITH MIGMATITES Mineralogical Association of Canada Short Course Series Volume 38

Melt inclusions in volcanic systems: methods, applications and problems

Atlas of plutonic rocks and orthogneisses in the Bohemian Massif

Melt Inclusions in Volcanic Systems: Methods, Applications and Problems

Melt

Melt

Krakow Melt

Melt Me

Krakow Melt

Krakow Melt

Biochemistry: A Short Course

Frosty Melt

Melt Me

Melt Me

A Short Course in Discrete Mathematics

A Short Course in Discrete Mathematics

A Short Course in General Relativity

A Short Course in Quantum Information Theory

A short course in intellectual self-defense

A Short Course in Technical Trading

Rocks

Rocks

A Short Course In Discrete Mathematics

A Short Course in Digital Photography

A Short Course in General Relativity

A short course in general relativity

A short course in general relativity

A Short Course in Intellectual Self-Defense

A Short Course in Nikon D70 Photography

Mineralogical Association of Canada Short Course, 36 MELT INCLUSIONS IN PLUTONIC ROCKS

Mineralogical Association of Canada Short Course, 40 Laser Ablation–ICP–MS in the Earth Sciences

WORKING WITH MIGMATITES Mineralogical Association of Canada Short Course Series Volume 38

Melt inclusions in volcanic systems: methods, applications and problems

Atlas of plutonic rocks and orthogneisses in the Bohemian Massif

Melt Inclusions in Volcanic Systems: Methods, Applications and Problems

Melt

Melt

Krakow Melt

Melt Me

Krakow Melt

Krakow Melt

Biochemistry: A Short Course

Frosty Melt

Melt Me

Melt Me

A Short Course in Discrete Mathematics

A Short Course in Discrete Mathematics

A Short Course in General Relativity

A Short Course in Quantum Information Theory

A short course in intellectual self-defense

A Short Course in Technical Trading

Rocks

Rocks

A Short Course In Discrete Mathematics

A Short Course in Digital Photography

A Short Course in General Relativity

A short course in general relativity

A short course in general relativity

A Short Course in Intellectual Self-Defense

A Short Course in Nikon D70 Photography

Recommend Documents