GEOCHEMISTRY AND MINERALOGY OF THE RARE EARTHS 43 o b s e r v e d in most eclogites do not match those f o u n d in most basalts, so a simple, closed system recrystallization at high pressure and t e m p e r a t u r e is an inadequate explanation. It is difficult to melt eclogite partially and l e a v e - o n e phase preferentially behind, b e c a u s e under mantle conditions the t e m p e r a t u r e interval b e t w e e n beginning and completion of melting is very small.
44 L.A. H A S K I N A N D T.P. P A S T E R
7.1. Mathematical models
During rock-forming or magma-forming processes in nature, with several major and minor minerals involved, what do the lanthanides do? In most systems they do not form their own minerals, but disperse themselves as trace constituents of phases in which they are not essential components.
Numerous attempts have been made by geochemists to account for the ways in w h i c h t r a c e elements distribute themselves. It has not been possible to define the energies of different possible locations of trace elements so that their equilibrium distributions can be accurately predicted. There are primitive notions, e.g., that the lanthanides merely follow calcium, that are too crude to be of much use, although there is a grain of truth in them. The most useful rule is that lanthanides (and other trace cations) readily enter those sites in crystals normally occupied by more common cations of approximately equal ionic radius (e.g., Neumann et al., 1966). Effects on lanthanide behavior of balancing charge in the event that the cation substituted for was not 3+ seem to be second order for chemically complex natural systems as long as lanthanide concentrations are low (tens of ppm or less).
Morris (1975) synthesized several silicate and aluminate compounds some of which are end members of important naturallyoccurring solid solution minerals.
The compounds were doped with Gd(III) or Eu(II). From electron paramagnetic resonance studies Morris showed that in Ca and Ca-Mg ortho- and metasilicates, both Gd(III) and Eu(II) entered mineral crystal sites normally occupied by Ca(II). Charge compensation was apparently remote from the lanthanide-oc- cupied sites since it had no effect on the spectra. In triclinic anorthite (CaAI2Os), the end-member of the major natural feldspar group, Eu(II) entered the Ca(II) site but Gd(III) showed the disordered spectrum characteristic of a glass, either because the Gd(III) was small relative to the size of the site or because the 3+
ion so strained the site that its symmetry was disrupted. (The affinity of feldspars for Eu relative to other lanthanides has been mentioned previously.) Both Gd(III) and Eu(II) tended to form clusters of several ions in pure Mg ortho- and metasilicates, indicating that those ions did not ,ubstitute for Mg(II) in regular structural sites. (Nevertheless, the ferromagnesian ortho- and metasilicates in nature each show strong and characteristic selectivity favoring the heavy lanthanides.)
Common, rock-forming silicate minerals are regarded as excluding lanthanides from their structures. Cullers et al. (1970) indicate, however, that substantial concentrations (hundreds of ppm) of lanthanides can be readily incorporated into such minerals; apparently the minerals do not compete successfully against silicate liquids to obtain them.
The currently most useful approximations o f trace element behavior for quantitative prediction are based on a simple Nernst distribution for equilibrium partitioning of a solute between two phases. This concept can be formalized to account for exchange between the trace element and the major ion it is deemed to replace, to account for effects of compensation of charge, etc. (e.g., Mclntire,
G E O C H E M I S T R Y A N D M I N E R A L O G Y O F T H E R A R E E A R T H S 45
1963). For purposes of this discussion (and for most practical purposes at the present level of sophistication of the field), a simple mass concentration ratio is adequate; i.e., a distribution coefficient, D, is defined as the ratio of the concentration of a trace element in ppm in a crystal to the concentration of that trace element in ppm in the liquid in equilibrium with that crystal. (A similar ratio between two solid phases is sometimes also useful.)
The limitations on this approach (e.g., Cullers et al., 1970) are that the trace ion in question be so dilute in the system that no possible means of distributing it among the phases will affect the thermodynamic activities of the major ions and required charge-compensating species, including defects. In such a case, the ratio of the activity of the trace ion in the solid phase to that of the liquid phase is a constant. The value of D is equal to that constant times the dilute solution activity coefficient of the trace ion in the liquid phase divided by that in the solid phase. It is presumed that those activity coefficients are insensitive to trace element concentration and that changes in temperature, pressure, or bulk composition of liquid and solid phases do not affect the properties of those phases as solvents, which would change the value of D. Of course, all those conditions are violated in nature, and in ways that could ultimately prove useful in providing information about the history and origin of a rock. These violations appear to have only second order effects on values of distribution coefficients for lanthanides between common silicate liquids and minerals, especially on values of individual lanthanide distribution coefficients relative to values for the other lanthanides in the same system. The values of D reflect the selectivity of individual phases for different members of the lanthanide series.
As mentioned above, the concept of the distribution coefficient implies equili- brium among phases. Rocks seldom crystallize under conditions approaching true equilibrium. For example, minerals are often zoned, i.e., their major and trace element compositions change systematically from their interiors to their rims, reflecting substantial changes in composition of the liquid as fractional crystallization,proceeded. Equilibrium in such cases can only obtain between the newly forming surfaces of crystals and their parent liquids. Interior portions of crystals are effectively isolated from the system and do not further react to main- tain liquid-solid equilibrium in the general sense. The Nernst distribution can still be considered to hold between the bulk residual liquid and the forming crystal surface. The mathematical description of this process was given by Doerner and Hoskins (1925). Many variants tailored to geochemical uses are given in the literature (e.g., Mclntire, 1963; Schilling and Winchester, 1967; Gast, 1968;
Haskin et al., 1970; Shaw, 1970; Greenland, 1970; B a n n o and Matsui, 1973;
Hertogen and Gijbels, 1976). In this approach, the concentration CL.m of trace element m in the residual liquid from a parent liquid with initial concentration CA.m when fraction X of the original liquid has solidified is given by eq. (21.1)
CL,m = CA,m(1 -- X ) Dw,m-1. (21.1)
The parameter
Dw,m
is the solid/liquid distribution coefficient for element m.From a mass balance, the average concentration CS.m in the solid at the same
46 L.A. H A S K I N A N D T.P. P A S T E R
extent (X) of crystallization is given by eq. (21.2)
C S , m = CA,m[1 -- (1 -
X)Dw.m]/X,
(21.2)If the value of Dw,m exceeds unity, the liquid becomes depleted exponentially in element m as crystallization proceeds (fig. 21.18), and the average concen- tration of m in the solid decreases accordingly. If the value of Dw.m is less than unity, the concentration of m increases exponentially in the liquid and, therefore, in the solid. At X = 1, the average concentration of the solid equals CA,m, the concentration of the starting material•
To put this simple model into practice for describing a natural system is
1000
100
..z.
¢ J
§
<>
m (2e ...a z
z
•1
• 01
0 .9 .99 .999
FRACTION SOLIDIFIED (X)
Fig. 21.18. Theoretical behavior of a trace e l e m e n t during fractional crystallization of a liquid, according to eq. (21.1). Concentration of the trace element relative to that of the starting composition for the liquid (or average composition for the entire system) is given as a function of the fraction of the original liquid solidified, for various values of the distribution coefficient (i.e., concentration of the trace element in the crystallizing solid/concentration in equilibrium liquid).
GEOCHEMISTRY AND MINERALOGY OF THE RARE EARTHS 47 substantially more complicated than equations (21.1) and (21.2) might imply, but fairly straightforward (e.g., H e l m k e et al., 1972; Paster et al., 1974). Values of Dw,m are sums of individual values for each crystallizing mineral, weighted according to the fraction of the precipitating solid that c o r r e s p o n d s to each mineral. Compositional fractionation is so severe that some minerals m a y cease to form and new minerals begin, so the calculation must be done in stages, with appropriate changes in mineral proportions and values of distribution coefficients.
T h e n there remains the question of whether the process by which a natural rock f o r m e d has been a d e q u a t e l y described. If the minerals f o r m too rapidly, the liquid will not be well stirred and at the surface of the crystal will be depleted relative to its bulk concentration in those elements with values of D exceeding unity and enriched in those elements with values of D less than unity. T h e effect is f o r the minerals to crystallize according to apparent values of D that are closer to unity than are the equilibrium values. This and other complications have been considered by Albarede and Bottinga (1972).
Similar considerations arise as solids partially melt to form liquids. The entire liquid f o r m e d can r e a c h equilibrium with the residual solid (e.g., Shaw, 1970), can reach equilibrium with the solid in increments only, which are separated away as melting continues (an e x t r e m e case which can be described by equa- tions similar to (1) and (2); e.g., Haskin et al., 1970), can derive from lowest melting fractions too rapidly for any significant equilibrium to be established, or can establish surface equilibrium with the residual solid (e.g., Shaw, 1970).
Because of the combination of thermal and pressure gradients with depth, a zone of liquid might rise f r o m the mantle, extracting trace elements along its path (e.g., Schilling and Winchester, 1967).
In order to use any of the a b o v e models f o r fractional crystallization or partial melting, values for distribution coefficients must be obtained. Many are given in t h e literature. T h e e x t e n t to which they are applicable to quantitative modelling
of natural systems has yet to be defined.
7.2. Distribution coe~icients
Values for lanthanide distribution coefficients have been estimated from o b s e r v e d partitioning of lanthanides in natural systems and from laboratory measurements. The advantage of using natural materials is that the lanthanides in those have distributed themselves under fully natural conditions in real rock systems. The disadvantage is that the materials o n which m e a s u r e m e n t s are made may not actually r e p r e s e n t the sort of natural, equilibrium systems that we imagine them to be. The advantage of m e a s u r e m e n t s in the laboratory is our ability to control compositions, temperatures, pressures, and other conditions of the experiments. The disadvantage is that we may improperly simulate natural situations and only imagine that we are measuring parameters of useful predic- tive value.
E v e n if we do measure valid distribution coefficients for a given natural or
48 L.A. H A S K I N A N D T.P. P A S T E R
experimental system, those values may not be applicable to the next natural system of interest. Values of D depend, as pointed out earlier, on temperature, pressure, and bulk compositions of liquid and solid phases; these parameters vary widely in nature. Then there remains t h e question of which form, if any, of the mathematical models using distribution coefficients properly describes the system in question.
Enough studies have been done to demonstrate unequivocally the value of mathematical models to estimate lanthanide behavior during rock and magma- forming processes. Most results are rather semiquantitative, subject to substan- tial uncertainties. The extent to which lanthanide distributions can yield more quantitatively accurate information such as lanthanide concentrations in source regions, or fractions of liquids solidified or solids melted, or to enable estab- lishment of clear-cut genetic relationships is still being evaluated. One of the crucial steps in this evaluation is accurate determination of values for lanthanide distribution coefficients and the extent to which those values depend on pres- sure, temperature, and composition. Several studies have been done, but few systematic measurements to show effects of temperature, pressure, and composition have been done.
Schnetzler and Philpotts (1968) obtained values for lanthanide distribution coefficients by the "phenocryst-host matrix" method. Phenocrysts are large mineral crystals found in otherwise fine-grained frozen lavas. The phenocrysts are presumed to have grown in equilibrium with the lava at depth prior to eruption of the lava. On eruption, the phenocrysts are swept along with the host lava, which quickly chills around them. By measuring the lanthanide concen- trations in the phenocrysts (solid phase) and the host matrix (liquid phase) and taking the ratio, values for distribution coefficients are obtained. Accuracy depends on whether the ~henocrysts truly grew in equilibrium with the liquid, whether the host matrix really represents that liquid, and whether both phases can be sampled without contamination by each other and accurately analyzed.
The principal result of Schnetzler's and Philpotts' work was demonstration that values obtained for a given mineral (e.g., high-Ca pyroxene) from different phenocryst-matrix pairs are very similar in a r e l a t i v e s e n s e . This success has prompted many more measurements on phenocryst-matrix pairs (some of which it is a strain to imagine to represent actual equilibrium pairs; some phenocrysts are severely zoned, some may not be phenocrysts, and some matrices are not obviously chilled parent liquids). Values have been presented by Schnetzler and Philpotts (1968, 1970), Onuma et al. (1968), Higuchi and Nagasawa (1969), Nagasawa and Schnetzler (1971), and Dudas et al. (1971). Typical values are shown in fig. 21.19 for phenocryst-matrix distribution coefficients for the minerals feldspar, high-Ca pyroxene, low-Ca pyroxene, olivine, garnet, and apatite.
Estimates of values have been obtained in other ways on natural systems by Frey (1969), Balashov (1972), Paster et al. (1974), and Haskin and Korotev (1977), some to be discussed later.
values on synthetic systems have been made in several laboratories by
G E O C H E M I S T R Y A N D M I N E R A L O G Y O F T H E R A R E E A R T H S 49
50 I I I I I I I I I I I I I I I--
~- .5 j z _ ~ . ~ % ~ f -
_ o f - _
N .1 .05
× x
i
ka Ce Pr Nd Pm Sm Eu gd Tb Dy Ho Er Tm Yb ku LANTHANI DE ATOMIC NO.
Fig. 21.19. Typical v a l u e s of lanthanide distribution coefficients for c o m m o n , r o c k - f o r m i n g minerals.
Values f o r apatite inferred f r o m a n a l y s i s of S k a e r g a a r d i n t r u s i o n ( P a s t e r et al., 1974); r e s t f r o m p h e n o c r y s t - m a t r i k a n a l y s e s ( S c h n e t z l e r and P h i l p o t t s , 1970).
i
different techniques. Cullers et al. (1970, 1973) synthesized pure silicate minerals in the p r e s e n c e of w a t e r and trace a m o u n t s of lanthanides. T h e y m e a s u r e d their distributions against w a t e r and also the distribution b e t w e e n w a t e r and a silicate liquid. T h e y derived e s t i m a t e s f o r lanthanide distribution coefficients that were consistent with those f r o m p h e n o c r y s t - m a t r i x m e a s u r e m e n t s but in the lowest part of the range. T h e y f o u n d that both relative and, to a greater degree, absolute values were sensitive to the t e m p e r a t u r e of equilibration. Zielinski and F r e y (1974) did nearly identical e x p e r i m e n t s on Ca-rich p y r o x e n e and w a t e r but found nonequilibrium effects not o b s e r v e d by Cullers et al. (1973) that they felt might be responsible for s o m e of the a p p a r e n t effects of t e m p e r a t u r e . The r e a s o n s for the differences in results b e t w e e n the two laboratories are still not known. The technique of equilibrating minerals with water instead of with their p a r e n t silicate liquids m a k e s possible a clean separation b e t w e e n the p h a s e s after the
50 L.A. HASKIN AND T.P. PASTER
f u r n a c e charge is quenched. T h e potential of the m e t h o d as a m e a n s of providing accurate values of distribution coefficients and t h e effects of t e m p e r a t u r e , pressure, and c o m p o s i t i o n on them has not been fully evaluated. It seems inherently less promising than m e t h o d s that rely on the direct synthesis of crystals f r o m silicate melts.
Balashov et al. (1970b) did zone melting e x p e r i m e n t s on synthetic metasili- cates that indicated that the light lanthanides w e r e m o r e easily e x t r a c t e d into the liquid phase than w e r e the h e a v y ones, especially w h e n Mg(II) was the m a j o r cation (as o p p o s e d to Ca[II]). Equilibrium b e t w e e n melt and crystals was thought not to h a v e b e e n attained, so no values of distribution coefficient were given.
M a s u d a and K u s h i r o (1970) s y n t h e s i z e d lanthanide-doped, Ca-rich p y r o x e n e f r o m silicate liquid and m e a s u r e d values of distribution coefficients. Values fell near the middle and u p p e r part of the range of those obtained b y the p h e n o c r y s t matrix m e t h o d . Such m e a s u r e m e n t s are v e r y difficult b e c a u s e e v e n small inclusions of glass (quenched silicate liquid) with the crystals a n a l y z e d can contribute significantly to the m e a s u r e d concentration of a lanthanide if its true value for D is ~0.2. This p r o b l e m w a s recognized but not fully solved by M a s u d a and Kushiro.
G r u t z e c k et al. (1974) m e a s u r e d values for lanthanide distribution coefficients in Ca-rich p y r o x e n e grown f r o m a silicate liquid. T h e y o b t a i n e d values near the middle of the range f o u n d in p h e n o c r y s t - m a t r i x studies. T h e y s h o w e d that, at low partial p r e s s u r e s of o x y g e n , Eu w a s preferentially excluded relative to other lanthanides f r o m the p y r o x e n e b e c a u s e m u c h of it was in the 2+ oxidation state, which has a lower value of D than Eu in the 3+ state. To avoid the necessity of separating crystals f r o m glass, sufficiently high lanthanide c o n c e n t r a t i o n s were used (~> 1%) so that they could be d e t e r m i n e d b y electron m i c r o p r o b e . Such high c o n c e n t r a t i o n s are in risk of violating the dilute solution b e h a v i o r of the lanthanides, t h e r e b y yielding e r r o n e o u s values for D ; such effects were not f o u n d by G r u t z e c k et al. Weill et al. (1974) used the results of G r u t z e c k et al. to model the b e h a v i o r of Sm and Eu during lunar igneous differentiation.
D r a k e and Weill (1975) used the s a m e e x p e r i m e n t a l technique as G r u t z e c k et al. to d e t e r m i n e lanthanide distribution coefficients for f e l d s p a r crystallized f r o m a silicate liquid. Values fell a b o u t the middle of the range o b s e r v e d f o r pheno- c r y s t - m a t r i x pairs. T h e y found, as e x p e c t e d , that D for Eu increased as the partial p r e s s u r e of o x y g e n was d e c r e a s e d (fo2 ranged f r o m 10 -07 to 10-125).
L a n t h a n i d e distribution coefficients w e r e determined by Shimizu a n d Kushiro (1975) on a synthetic garnet. T h e garnet was separated f r o m the q u e n c h e d p a r e n t liquid for m e a s u r e m e n t of lanthanide c o n c e n t r a t i o n s b y isotope dilution m a s s s p e c t r o m e t r y . Values f o r the heavier lanthanides (fig. 21.20) are not a p p r e c i a b l y disturbed b y possible c o n t a m i n a t i o n by q u e n c h e d liquid, but those for the lightest lanthanides m a y be substantially too high.
M y s e n (1976) m e a s u r e d distribution coefficients for Sm in olivine and C a - p o o r p y r o x e n e by b e t a track m a p p i n g of synthetic silicates. T h a t technique allows use of v e r y low c o n c e n t r a t i o n s of a trace element. M y s e n r e p o r t e d a strong concert-
GEOCHEMISTRY AND MINERALOGY OF THE RARE EARTHS 51
I I J I I I I I I I I I s( ~
t 1 1 i f
PPM PYROXENE PPM GARNET -
-
o y - - ; y
PYROXENEj ~' 11
1 1
f , s f l "
.1 A " tO.
I I
• 05 l
- / D GARNET I I
] d'
I [ [ I I [ l I I I [ I I
La Ce Pr Nd Pm Sm Eu Gd Tb Dy Ho Er Tm Yb Lu LANTHANI DE ATOMI C NO.
Fig. 21.20. Comparison diagram for a hypothetical eclogite consisting of equal parts of garnet and Ca-rich pyroxene, grown under closed system conditions from a liquid with 1 ppm of each lanthanide and with distribution coefficients as shown (that for garnet from Shimizu and Kushiro, 1975, and that for pyroxene from Schnetzler and Philpotts, 1970). Note that the pyroxene is strongly enriched in light lanthanides despite its preference for heavy ones.
tration d e p e n d e n c e of D for Sm in olivine and o r t h o p y r o x e n e , a p h e n o m e n o n not o b s e r v e d by other investigators. T h e c o n c e n t r a t i o n range studied g a v e olivine ranging f r o m 0.04 to 1.6 p p m and that for low-Ca p y r o x e n e f r o m 0.08 to 2.6 ppm. T h e range of c o n c e n t r a t i o n s of Gd in the e x p e r i m e n t s by Cullers et al.
(1970) for oliv, ine was - 0 . 0 0 3 to > 1 0 0 0 p p m and for low-Ca p y r o x e n e was - 0 . 0 0 5 to > 1 0 0 0 p p m , without m e a s u r a b l e deviation f r o m c o n s t a n t values.
T h o s e e x p e r i m e n t s were f o r partitioning b e t w e e n the minerals and an a q u e o u s phase, rather than b e t w e e n minerals and silicate liquid. H o w e v e r , the lowest values of D r e p o r t e d by M y s e n were for the liquids m o s t c o n c e n t r a t e d in Sm, so failure of the liquid to retain increasing quantities of Sm was not important. The r e a s o n s for the discrepancies b e t w e e n the t w o e x p e r i m e n t s are not known.
F r o m the data in fig. 21.19 it can be seen that m o s t minerals do not r e m o v e lanthanides effectively f r o m silicate liquids (i.e., lanthanide D values are m o s t l y less than unity). E x c e p t i o n s are apatite (a lanthanide-concentrating mineral that is a c o m m o n late-stage p r o d u c t of crystallization of mafic liquids) and garnet (which c o n c e n t r a t e s h e a v y lanthanides but not light lanthanides). In acid mag- mas, lanthanide minerals m a y precipitate; these are discussed in later sections.
Ca-rich c l i n o p y r o x e n e s and h o r n b l e n d e s (related to p y r o x e n e s , but m o r e compositionally c o m p l e x ) c o m p e t e for lanthanides successfully against s o m e - what acidic lavas ( N a g a s a w a and Schnetzler, 1971).
52 L.A. HASK1N AND T.P. PASTER
Crystallization of garnet from a liquid will rapidly change the lanthanide distribution of that liquid. Derivation of a liquid from a solid containing several p e r c e n t or more of garnet will p r o d u c e a liquid with a lanthanide distribution very different from that of the initial solid. Crystallization of apatite from a liquid can l o w e r the lanthanide concentrations, but is more likely to decrease the rate at which lanthanide concentrations of the liquid increase, since the fraction of the total solid precipitating at any given time that is apatite is likely to be small. Partial melting of most sources containing apatite would be e x p e c t e d to melt all the apatite during the first v e r y few p e r c e n t of melting, so only in a special case would there by partitioning of lanthanides b e t w e e n residual apatite and a derived liquid. Crystallization of plagioclase can deplete Eu significantly relative to neighboring lanthanides. Partial melting against plagioclase can p r o d u c e liquids relatively deficient in Eu. Crystallization of Ca-rich p y r o x e n e s slowly changes the lanthanide distribution of the parent liquid. Crystallization of Ca-poor p y r o x e n e or olivine or of plagioclase (except for Eu) scarcely affects the lanthanide distributions of the parent liquids; although those minerals are strongly selective in which lanthanides they r e m o v e , they tend to r e m o v e such small amounts that the effect on the distribution of the residual liquid is nearly negligible. Precipitation of lanthanide minerals, which appears to o c c u r in granitic magmas, can deplete the residual liquid, sometimes selectively, in lanthanides.
Observed lanthanide distributions in analyzed minerals do not always reflect in an obvious way their distribution coefficients. The coefficients act on some parent distribution, which may differ substantially from the chondritic values c o m m o n l y used as a standard for producing comparison diagrams. Also, when a liquid crystallizes as a closed system (Helmke et al., 1972), essentially all the lanthanides m a y enter major minerals, each mineral becoming most enriched in the trace elements against which it discriminates least. For example, fig. 21.20 shows the lanthanide distributions foi" garnet and high-Ca p y r o x e n e from a hypothetical eclogite. The heavy-lanthanide selective p y r o x e n e is rich in the light lanthanides and depleted in the h e a v y ones because it has equilibrated with garnet, which has a stronger affinity than the p y r o x e n e for the h e a v y lanthanides and a weaker affinity for the light ones.
7.3. Eu anomaly
The anomalous behavior of Eu that results because Eu can be readily reduced to a 2+ ion with chemical properties unlike those of the 3+ lanthanides has been evident in m a n y of the lanthanide distributions shown above. The existence of this anomaly has been k n o w n for a long time (e.g., Goldschmidt, 1938). There has been discussion about whether the anomaly o b s e r v e d in igneous rocks is a result of c o n v e r s i o n of Eu(III) to Eu(II) in the liquid phase or of crystallo- chemical effects (i.e., r e a d y a c c e p t a n c e of Eu(II) but rejection of Eu(III) by some mineral, e.g., feldspar) (e.g., Philpotts and Schnetzler, 1968, 1969; Toweil et al., 1969). Both effects are important, of course. No mineral can take up Eu(II)