UThRa radioactive disequilibria and magmatic processes

Earth and Planetary Science Letters, 90 (1988) 243-262

243

Elsevier Science Publishers B.V., Amsterdam - Printed in The Netherlands [11

U-Th-Ra radioactive disequilibria and magmatic processes M . C o n d o m i n e s a, C h . H e m o n d

2,. a n d C.J. A l l 6 g r e 2

I Centre de Recherches Volcanologiques (CNRS UA 10), O.P.G.C., Universit$ Blaise Pascal, 5 rue Kessler, F-63038 Clermont-Ferrand Cedex (France) 2 Laboratoire de G$ochimie et Cosmochimie, lnstitut de Physique du Globe et U.E.R. des Sciences de la Terre de l'Universit$ Paris VII, 4 Place Jussieu, F-75252 Paris Cedex 05 (France)

Revised version accepted February 3, 1988 This paper presents the available data on U-Th-Ra disequilibria in recent volcanic rocks, and their significanceto the understanding of magmatic processes. It is shown that 238U-23oTh and 226Ra-23oTh disequilibria are particularly useful in the study of mantle sources, their recent modifications, the transfer time of the magmas towards the surface and magma dynamics in active volcanoes. The Th-Sr isotopic correlation is confirmed by new results on MORB and OIB. These magmas must have short (< 10 4 years) transfer times. In contrast with MORB and OIB data, which have activity ratios (238U/23°Th)< 1, many subduction zone volcanoes display activity ratios (238U/23°Th) > 1, together with 226Ra enrichments over 23°Th. The latter are largest in continental ultrapotassic and related rocks. This suggests that fluids play an important role in magma genesis processes in subduction zones, before (metasomatism) and/or during partial melting. In the absence of fluids, the enrichment pattern is Ra = TH > U, and in a "wet" environment Ra> U >Th. The study of (23°Th/232Th)o initial ratios through the volcanic history of a single volcano confirms that basaltic volcanoes on oceanic islands, such as Piton de la Fournaise (Reunion), Mauna Kea (Hawaii), Marion (Prince Edward hot spot) have short magma transfer times (< 10 4 years). Other volcanoes, like Etna, located in continental crust areas, appear to have longer transfer times (> 105 years). From the few available results on Ra-TH disequilibria, it seems that the time of magma transfer in the volcanoes studied so far is always longer than 100 years. Work in progress on Etna shows that Ra enrichment is a characteristic of the whole historic period and a deep-seated phenomenon.

1. Introduction The discovery of radioactive disequilibria in volcanic rocks was m a d e very early in the 20th century. Joly in 1909 [1] f o u n d r a d i u m excesses in volcanic rocks from Vesuvius. But it was n o t u n t i l half a c e n t u r y later that a few other works appeared on this subject. F o r example, B e g e m a n n et al. [2] a n d later H o u t e r m a n s et al. [3] m e a s u r e d 21°pb in volcanic sublimates from Vesuvius. Systematic m e a s u r e m e n t s in a variety of volcanic rocks were then u n d e r t a k e n i n several countries (e.g. [4-7]) m a i n l y i n Russia, Japan, Italy, a n d the U n i t e d States of America. O n e of the goals of these studies was to use radioactive disequilibria in rocks a n d minerals to date y o u n g volcanic * Present address: M.P.I. fi~r Chemie, Abteilung Geochemie, Postfach 3060, D-6500 Mainz, F.R.G. 0012-821X/88/$03.50

© 1988 Elsevier Science Publishers B.V.

rocks. But the i m p o r t a n c e of radioactive disequilibria for the study of m a g m a t i c processes was soon realized, as d e m o n s t r a t e d b y the paper b y Oversby a n d G a s t [8]. F o r the last twenty years this d o m a i n has b e e n in c o n s t a n t development. T h e interest has b e e n m a i n l y c o n c e n t r a t e d o n 238U-23°Th disequilibria. These are used as a dating m e t h o d for volcanic rocks y o u n g e r than 300,000 years, a n d also as tracers of m a g m a t i c processes. I n recent years, shorter-lived isotopes of the radioactive series like 226Ra, 228Ra, 228Th, 21°pb, 2a°po, etc., have also b e e n m e a s u r e d in rocks from active volcanoes. The p u r p o s e of this p a p e r is to show the m a i n results o b t a i n e d b y the use of radioactive disequilibria ( m a i n l y b e t w e e n 238U, 2 3 ° ~ 1 , 226Ra), a n d their b e a r i n g o n p r o b l e m s such as m a n t l e sources, m e t a s o m a t i s m , m e c h a n i s m s of m a g m a genesis, transfer of m a g m a s towards the surface a n d

244

magma dynamics. We will first recall some basic principles of radioactive disequilibria, as they are still less familiar to most geochemists than the classical isotopic tracers such as Sr, Nd or Pb.

2. Basic principles of radioactive disequilibria 2.1. General considerations In a natural radioactive decay chain, where the parent has a much longer half-life than the daughter nuclides, a state of radioactive equilibrium in the whole chain (secular equilibrium) is reached within about 4 or 5 half-lives of the longest-lived intermediate nuclide. The activities of all the nuclides are then equal: ~klN 1 = ~k2N 2 = • -- = ~ N k (~i is the radioactive decay constant, and ~ the number of atoms of nuclide i). In the 238U chain for example (Fig. 1), the longest-lived intermediate nuclide is 23°Th with a half-life of 75,200 years, and secular equilibrium is reached after about 300,000 years. Equilibrium between two successive nuclides I. and I.+1 having very different half-lives, T. being much greater than T. +1, follows the same rule and is established at a rate depending on the half-life of the daughter nuclide T.+ t- A group of nuclides in radioactive equilibrium behaves as a single isotope having the half-life of the parent of the group. For example, the very short-lived nuclides 234Th and 234pa very quickly reach equilibrium with 238U (within 3 - 4

months), and 234U can be considered as the direct daughter of 238U. Disequilibrium is created when fractionation occurs between two nuclides belonging to different chemical elements (U and Th for example). The general equations giving the activity of a nuclide in a decay series have been established by Bateman [9]. For a more complete discussion on radioactive disequilibria, the reader is referred to the book by Ivanovich and Harmon [10]. 2.2. U-Th-Ra disequilibria in volcanic rocks The problem of radioactive disequilibria in the 238U series in volcanic rocks is further simplified by the fact that 234U and 238U are always in equilibrium in non-altered rocks ((234U/238U) activity ratio = 1). So 23°Th can be considered as the direct product of the decay of 238U: 238 U

~

T=4.47 × 10 9 y

230Th

~

T= 75,200 y

226 Ra

T= 1600 y

The equation giving the activity ratio (Z3°Th/ 232Th) is then: 230Th

238U ]

( 23°Th

X being the 23°Th decay constant. Throughout this paper, parentheses denote activities. Following a suggestion by All~gre [11], 232Th is taken as a stable isotope, which allows the use of an isochron diagram ( 2 3 ° T h / 2 3 2 T h ) - (238U/

234T h 4 47~109y

j

234pj 222Rn 75,200y

~'~

1,600y

224R a

218 Po

214Pb

210TI 214P0j ~op-~" 210B i ~22y J 21OPo 2O6pb 214Bi --~-....

220Rn

216p0

1.91y

212Pb 212Bi . . ~ _ . . . 208TI

212P0 j

2O8pb""

Fig. 1. 238U and 23ZTh decay series with the main nuclides of interest for this study and their half-lives (in the boxes).

245

Fig. 2. The (230Th/232Th)-( 238U/232Th) isochron diagram. A similar diagram can be used in the interval O-30 years: ( “’ Ra/226 Ra) versus ( 232Th/226 Ra).

232Th) (Fig. 2). A classical application of this diagram is the dating of young volcanic rocks by internal isochrons (e.g. [12-141). The slope gives the age of the U-Th fractionation, and the intercept with the equiline gives the (230Th/232Th)0 initial ratio. The equiline is the line where (230Th) = (238U) and is the ultimate position of the iso&on when t tends towards infinity. In such a diagram, a U-Th chemical fractionation is represented by horizontal vectors, and radioactive decay by vertical vectors. The use of 230Th-238U disequilibria is limited to the approximate range O-300,000 years. Generally speaking, radioactive disequilibria are useful on timescales of the order of the half-life of the shorter-lived radionuclide involved. Hence 226Ra- 230Th disequilibria can be used in the range O-8000 years, and 228Ra-232Th disequilibria in the interval O-30 years ( 228Ra has a half-life of 5.77 years). It is worth noting that in the range O-30 years, 226Ra can be considered as a stable Ra isotope, owing to its much longer half-life of 1600 years. It is thus possible to write an equation similar to that describing the evolution of the (230Th/232Th) ratio:

An isochron X’ being the 228Ra decay constant. diagram ( 228Ra/**‘j Ra)-( 232Th/226Ra) can then be used (see Fig. 2) to date Ra-Th fractionations in

the range O-30 years. We will further discuss the interest of the two Ra-Th pairs in section 5. The main interests of the study of radioactive disequilibria in volcanic rocks, besides the possibility of dating young volcanics, are the following: firstly, they allow to put time constraints on the chemical fractionations occurring during magmatic processes and, secondly, they give a unique way of estimating the extent of these fractionations. In a number of cases indeed radioactive equilibrium can be assumed before the fractionations take place, and thus the initial ratio of two elements is known. This is best illustrated by a simple model concerning the partial melting of mantle sources (Fig. 3). The mantle source is in secular radioactive equilibrium if no recent (< 300,000 years) event has affected it. Partial melting induces fractionation between U and Th, and thus disequilibrium between 238U and 230Th. But the ratio (230Th/ 232Th) remains the same in the magma as in the mantle source. If the time of transfer of this magma towards the surface is short compared to the 230Th half-life (less than a few thousands of years), then the (230Th/ 232Th) ratio measured in the lavas represents the (238U/232Th) ratio of the mantle source [8,15]. The disequilibrium ratio k = (238U/230Th) in the lavas is a measure of fractionation during partial melting. For a mantle source with constant U and Th partition coefficients, smaller degrees of partial melting produce magmas with lower k ratios [16]. MacKenzie [17] used published values for mid-ocean ridge basalts

Fig. 3. Simple model of partial melting of P mantle source in radioactive equilibrium. k is a measure of the Th/U fractionation due to partial melting: k = (U/Th),/(U/Th)s where subscripts M and S refer to the magma and mantle source respectively.

246 and an experimentally determined partition coefficient for U of 0.0015 to propose a very small degree of melting (--1%) for the generation of MORB. This calculation, however, is critically dependent on the U and Th partition coefficients in the mantle sources, and very little is known about these parameters. Although a truly quantitative interpretation of these disequilibria seems premature for this reason, they can give some fundamental information about processes occurring in the mantle during partial melting as will be discussed in the following sections.

for the other classical isotopic ratios (Sr, Nd, Pb), we now have a fair amount of data on rocks from different geodynamic environments, especially basaltic rocks from mid-ocean ridges and oceanic islands, continental alkali basalts and calc-alkaline lavas from subduction zones. The range of (23°Th/232Th) ratios for most volcanic rocks is between 0.65 and 1.35, which corresponds to T h / U ratios of the mantle sources between 4.67 and 2.25 (note that T h / U weight ratios = 3 . 0 3 4 / ( 2 3 8 U / 2 3 2 T h ) ) . T h / U ratios measured in the same volcanic rocks are in the interval 1.44-4.67, thus a somewhat larger range than the above range for the mantle sources. The disequilibrium ratios k = ( 2 3 8 U / 2 3 ° T h ) vary between about 0.7 and 1.8. It is worth mentioning that some volcanic rocks fall outside these intervals: such is the case for the ultrapotassic lavas from Vico (Latium, Italy) [18], which may have T h / U ratio as high as 11, for ultrapotassic rocks from the East African rift (Kenya and Tanganyika) [19] or for the carbonatites from Oldoinyo Lengai [20] with T h / U ratios as low as 0.18. If we except these unusual cases, where additional U-Th fractionation mechanisms, like fluid transfer, are probably involved, T h / U ratios of most nonaltered volcanic rocks must be in the approximate range 1.5-5.0. Uranium is a relatively mobile ele-

3. Th isotopic geochemistry The (23°Th/232Th) ratios in present volcanic rocks, or initial (23°Th/=32Th)o ratios in recent ones ( < 300,000 years), can be used as isotopic tracers in the same way as 87Sr/86Sr ratios for example. But whereas a 87Sr/86Sr ratio is a function of the R b / S r ratio of the source, integrated over the whole history of the mantle, the (23°Th/ 232Th) ratio gives direct information on the present T h / U ratio of the source, as explained above.

3.1. Th isotopic ratios and T h / U ratios in volcanic rocks Although many fewer measurements are available for Th isotopic ratios in volcanic rocks than

Th/ 1.3 1.2 1.1

-0

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0 Famous 0

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~Karthala Mauna Loa

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0.7

~.5 5.0

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I

I

0.7030

0.7035

0.70 L,O

I

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0.70/+5 0.7050 a7Sr/86Sr

L

O.7055

Fig. 4. Th-Sr isotopic correlation for MORB and OIB, which define the "mantle array". References of the data: Atlantic 30 ° N. FAMOUS, Ardoukoba 1978 lava (Asal rift) [15]; Icelandic tholeiites [23,24]; Surtsey, Heimaey[23,53]; Faial 1672 laval flow (Azores), Karthala 1962 laval flow (Comores), Erta Ale 1971 lava flow (Afar) [24]; Marion [21]; Tristan da Cunha [8,24]; Savai'i (Samoa) [21]; Cha]ne des Puys [14]; Kilauea, Manna Loa [5,21,24,25]; Rrunion ([15 and M. Condomines, unpublished results on Piton de la Fournaise old and present lavas). "B.E." is the "bulk earth" value according to Allrgre et al. [22]. [Th/U]s is the Th/U ratio (weight ratio) of the mantle source.

247 cover a wide range of values extending from the field of MORB to a value of 0.77 for Savai'i, Samoa [21]. They are obviously derived from less depleted sources than MORBs, having higher T h / U and R b / S r ratios. A "bulk earth" value could be deduced from this correlation, assuming a 87Sr/86Sr ratio for the "bulk earth" from the Nd-Sr correlation. A T h / U value between 3.5 and 3.7 is obtained in this way. Recently, however, All6gre et al. [22] have shown that it is safer to calculate the " b u l k earth" T h / U value from the "radiogenic lead ratio":

ment during alteration and weathered volcanic rocks could show high T h / U ratios due to U leaching. On the other hand, U addition from seawater could explain some very low T h / U ratios in oceanic tholeiites. In both c a s e s (234U/238U) ratios different from unity could be an indicator of secondary alteration processes (the seawater ( 2 3 4 U / / 2 3 8 U ) ratio is 1.14 [10]). 3.2. T h - S r isotopic correlation

One of the main results of Th isotopic geochemistry was the discovery of a negative correlation between (23°Th/232Th) and 87Sr/86Sr ratios in MORB and OIB [15]. This correlation has been confirmed by many other results, which define a "mantle array", similar to the well known Sr-Nd "mantle array" (Fig. 4). This Sr-Th isotopic correlation is interpreted in the following way: the (23°Th/232Th) ratio represents the (238U/232Th) ratio of the mantle source (see Fig. 3), and the correlation implies that the T h / U and R b / S r ratios are correlated in these mantle sources. Interpretation of the data in such a diagram is very similar to that in a Nd-Sr correlation diagram. It is clear, for example, that MORBs have the highest Th isotopic ratios, thus coming from depleted sources with low T h / U and low R b / S r ratios. The OIBs have lower Th isotopic ratios (and higher Sr isotopic ratios) and

2°8pb* t = (2°8pb/2°4pb)M -(2°8pb/a°4pb)° (~

- _ //

1.2

1.1

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1.0

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/

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Ardoukob~Y - - r ~ - ÷ . . . . . • / ~-L Iceland Karfh~ata

0.7

/

P(2°6pb/2°4 b ) ~

where subscripts M and 0 refer to the measured and initial ratios (Canon Diablo values). The "bulk earth" T h / U ratio calculated in this way is close to 4.2 and is in agreement with data on komatiites [22]. This value would lead in the Sr-Th isotopic correlation diagram (Fig. 4) to a "bulk earth" 87Sr/86Sr ratio around 0.7055. The data for Tristan da Cunha and Samoa are the closest to this bulk earth value. The complement of the depleted mantle is the continental crust, which should have a mean T h / U ratio higher than 4.2. The progressive depletion of the mantle must be due to successive partial melt-

1.3

O.8

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Erfa ,, ~ / ' I

'~C.des

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L o~ __~" L°"/~l "B E " /

/,

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, I

0.7

0.9

I

i

i~

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Fig. 5. MORB and OIB data in the isochron diagram. References: Hawaii (K = Kilauea, M L = Mauna Loa, M K = Mauna Kea, H = Hualalai) [21,25]; Samoa, Marion [21]; East Pacific Rise (E.P.R) [26]. Referencesfor the other data are given in the caption of Fig. 4.

248

ing events, the partial melts having higher T h / U ratios than the solid residues. This is in agreement with the position of MORB and OIB data to the left of the equiline in the isochron diagram (Fig. 5). In order to discuss 238u-z3°Th disequilibria in recent volcanic rocks, it is very useful to plot the data in both the isochron diagram and the Sr-Th isotopic correlation diagram. But first, it is worth considering the possible mechanisms able to explain results which fall well outside the " m a n t l e array" in the Sr-Th isotopic diagram. These processes can be divided into three main categories: (1) Recent chemical fractions affecting the mantle sources. Any recent modification in the R b / S r and T h / U ratios of the mantle source will obviously affect the (23°Th/232Th) ratio, but will have very little or not effect on the 878r//868r ratio, because of the very long half-life of 87Rb. Thus if the R b / S r and T h / U ratios are modified by a recent partial melting or a metasomatic event, the Sr isotopic ratio m a y well remain constant, while the Th isotopic ratio will change by radioactive decay. If the T h / U ratio is increased in the process, (23°Th/Z32Th) will be lowered and the data moved below the correlation. If the T h / U ratio is decreased, the data will plot above the mantle array. (2) Mixing or contamination processes. This category includes a wide range of phenomena, such as mixing of mantle sources, mixing of magmas, crustal assimilation, selective contamination, etc. Mixing of two components situated on the Sr-Th correlation array will generally give a curve whose shape depends on the S r / T h ratios of the two components. Only if the S r / T h ratio is the same for the two end-members, will this curve be a straight line. For example, assimilation of an upper crustal component (granite) by a basaltic m a g m a will give a concave mixing curve below the Sr-Th correlation (the T h / S r ratio in the granitic component being much higher than that of the basalt). In fact, since Th is a magmaphile element whose behaviour is somewhat similar to that of Nd, the mixing curves are not very different in the Th-Sr and Nd-Sr isotopic correlation diagrams. An example of selective contamination can be found in some M O R B from the F A M O U S area, where (23°Th/Z32Th) ratios have been increased

by Th addition from seawater, and the data plot well above the mantle array [15]. (3) Non-negligible times of transfer of the magmas. This is indeed a very important process for the Th isotopic ratios. In the simple model explained in section 2.2, we consider that the time of transfer of the m a g m a towards the surface is negligible compared to the 23°Th half-life (less than a few thousands of years). If this is not the case, the (23°Th/Z32Th) ratio will change during the transfer of the m a g m a simply by radioactive decay. Depending on the position of the initial m a g m a in the isochron diagram, the ratio will decrease (magma to the left of the equiline) or increase with time (magma to the right of the equiline). The timescale of such a transfer (10 4 to 10 6 years) is too short to affect the Sr isotopic ratios in any way. The data would then be displaced vertically in the Th-Sr diagram, below or above the mantle array. If we reverse the reasoning, the fact that most basaltic rocks from M O R B and OIB analysed so far do define a mantle array may be taken as an argument for short transfer times of these magmas. 4. U-Th-Ra systematics in recent volcanics The following discussions will be based on results obtained on historical or well-dated volcanic rocks, except for the M O R B submarine glasses.

4.1. Mid-ocean ridge and oceanic island basalts e3SU-e3°Th disequilibria. We present here the results obtained on MORB and OIB by several authors ([8,14,15,21,23-26], and unpublished results by M. Condomines et al.). A basalt from the Cha~ne des Puys (French Massif Central) has been included as an example of continental alkaline volcanism, very similar to oceanic island alkalibasaltic volcanism. As we have just seen, the Th isotopic ratios measured in M O R B and OIB are inversely correlated with the 87Sr/86Sr ratios and define a " m a n tle array" (Fig. 4). In the isochron diagram, almost all data plot to the left of the equiline, showing (23°Th/238U) higher than 1. The parameter k (=238U/23°Th) varies between 0.7 and 1. We attribute the enrichment in 23°Th relative to 238U to

249 a partial melting event, which gives a melt with a higher T h / U ratio than that of the solid residue. Partial melting is indeed the more likely process in this case to fractionate Th and U, which are both magrnaphile elements, as discussed by MacKenzie [17]. Fractional crystallization only introduces small changes in T h / U ratios, as demonstrated by numerous studies (e.g. [27,28]). In fact a slight increase of T h / U ratios with differentiation can be expected, due to the more magmaphile character of Th relative to U. A good example of the variations of T h / U ratios in an oceanic environment is given by the data on Iceland [23,24]. Although fractional crystallization is not the only process at the origin of silicic rocks in Iceland, the T h / U ratios vary only between 2.9 for the most primitive tholeiites to 3.6 for some rhyolites, whereas the Th content displays a variation by a factor of more than 700. We thus believe that fractional crystallization can produce maximum T h / U variations of about 20% in the more differentiated products. This is in agreement with the fact that the major minerals have much lower U and Th concentrations than the groundmass and do not fractionate significantly the T h / U ratio, with the possible exception of opaque minerals [13,14]. The latter minerals generally have lower T h / U ratios than the whole rocks and they may be responsible for a slight increase of T h / U ratios during differentiation. However, it is not clear whether U and Th measured in opaque minerals separates are in the lattice of these minerals, or if they come mostly from micro-inclusions of U- and Th-rich accessory phases, like apatite or zircon. In some instances, such microphases have been identified by a coupled fission-track and electron microprobe study [29,30]. Only in the case where such accessory minerals (zircon, apatite, monazite) would fractionate in appreciable amount could the T h / U ratios be drastically changed. Moreover, it must be emphasized that all data plotted in the isochron diagram of Fig. 5 were obtained on basaltic rocks, either "primitive" or only slightly differentiated. It is thus clear that fractional crystallization cannot explain the observed disequilibria. It has also been argued [31] that some disequilibria could be produced by secondary U leaching during post-eruptive weathering. This is very unlikely as all the rocks we have analysed are very

fresh; some fresh glasses have been analysed from submarine tholeiites. The subaerial rocks often come from very recent historical eruptions, and some were sampled during the eruption or only a few days after (Ardoukoba, Afar, 1978; Piton de la Fournaise, Rfunion, 1977, 1986). Moreover, there is no systematic difference in (23°Th//z38u)0 ratios when present and older rocks from the same volcano are analysed, as demonstrated by the R6union results, where initial (23°Th/Z32Th)0 ratios of volcanic rocks older than 105 years are reported in Fig. 5, together with Th isotopic ratios measured on present lavas from the Piton de la Fournaise. We thus consider that, in the case of MORB and OIB, partial melting is the main process fractionating the T h / U ratio and producing the 23°Th- 238U disequilibria. As mentioned above, the parameter k depends on the degree of partial melting and should be lower for a smaller degree of melting, for a unique mantle source, and constant U and Th partition coefficients. As suggested by Allrgre and Condomines [16], this may be roughly true for the mid-ocean ridge tholeiites (Atlantic: 3 0 ° N , FAMOUS, Afar), although the East Pacific Rise resuits [26] show a wide range in k values. In Hawaii the 1801 alkali-basalt from Hualalai has a lower k value than the tholeiites from Kilauea or Mauna Loa, as expected if alkali-basalt derives from a smaller degree of melting of the mantle source. But, on the whole, there is no systematic difference between MORB and OIB, except that the latter normally have lower (23°Th/232Th) ratios, and are derived from less depleted sources.

226Ra-23°Th disequilibria. We have very few data on 226Ra-23°Th disequilibria in MORB and OIB. Krishnaswami et al. [31] reported a few data on MORB volcanic glasses from the Galapagos spreading centre. Among the four samples analysed, three have 226Ra in equilibrium with 23°Th. The authors consider that these samples are very young and that the 226Ra-23°Th equilibrium is a characteristic of the magmas. However, MacDougall and Rubin [32] have recently found variable but systematic 226Ra excesses in MORB glasses from the East Pacific Rise ((226Ra/Z38u) up to 2.5). Obviously the lack of precise ages on sub-

250 marine volcanic glasses, in the range 0-8000 years, will often make doubtful the interpretation of 226Ra-Z3°Th disequilibria, Moreover, because of the short half-life of 226Ra (1600 years), the problem of the time of transfer of the m a g m a is much more critical in the case of the 226Ra-Z3°Th disequilibria. In a few hundreds of years, the (226R a / 23°Th) ratio can be significantly affected. Krishnaswami et al. [31] also report one analysis of the Krafla 1981 lava flow (Iceland), which again shows 226Ra and 23°Th in equilibrium. They also give disequilibria results for historical lava flows from Mauna Loa and Kilauea in Hawaii. Only two recent flows from Kilauea (1979, 1982) show a slight 226Ra excess ((226Ra/23°Th) = 1.18), all the other samples having (226Ra/Z3°Th)= 1 within analytical errors. These results seem in general agreement with the earlier data reported by Nishimura on historical lavas from the same volcanoes [7]. More recently, Turekian arid Reinitz [33] have found (226Ra/Z3°Th) ratios smaller than unity in 1983-1985 lavas from the Kilauea. It is worth mentioning also the early results obtained by Oversby and Gast [8]. Their data indicate 226Ra excesses in the 1961 lava flow from Tristan da Cunha ((226Ra/Z3°Th)= 1.23) and in the 1958 volcanic rocks from the Capelinhos eruption (Faial, Azores), where (226Ra/Z3°yh)= 1.74. From these very scarce data, it is not possible to draw any general conclusion. All we can say is that MORB may show 226Ra-Z3°Th equilibrium or 226Ra excesses. Oceanic islands like Hawaii also show 226Ra-23°Th equilibrium or only small ~excesses or deficiencies of 226Ra relative to 23°Th. Other oceanic islands m a y have greater 26Ra excesses (Azores).

23°Th-2~SU disequilibria. We have plotted the results on 238U-23°Th disequilibria in recent volcanic rocks from subduction zones in the isochron diagram (Fig. 7) and in the Sr-Th isotopic correlation diagram (Fig. 6) when 87Sr/86Sr ratios are available. It is obvious that the situation is much more complex than in the case of M O R B and OIB. In the Sr-Th isotopic correlation diagram, only a few data are located in the " m a n t l e array" previously defined by M O R B and OIB data. Many others plot above or below it (Fig. 6). Perhaps even more striking is the situation in the isochron diagram (Fig. 7). There are data points on both sides of the equiline, and some on the equiline itself. The parameter k = (238u/z3°Th) is in the range 0.8-1.4 for most rocks, but it can be as high as 1.8 in the Marianas volcanics (Fig. 7). Moreover the (23°Th/Z32Th) ratios are highly variable, from values higher than those of the M O R B (>~ 1.30) to very low values, less than 0.7. This variability clearly indicates a diversity of sources and genetic processes for the subduction zone magmas. When looking at the data in a more detailed way, however, some regional regularities do appear. The volcanic rocks having 23°Th in excess relative to 238U, and thus located to the left

sl ~eer~ 1.3

[Th/U]s

EoJmo Pclricutiq ~ C ~ r , ~ .~ '" Nyirogongc.

2.5

1.2 1.1 1.0

Grroy Arero[~ ~ ~ ~ \ , ~ El Cnic'~or' /i~\ ~ Etna(!~' \ ~ VI

I Eolion ore

0.9

D

....

4.2. Volcanism of subduction zones In this section we present the results obtained on recent volcanics from subduction zone environments, either oceanic or continental, from active or "fossil" subduction zones. The rocks belong to the calc-alkaline series and most are classified as andesites. The results on Etna are also discussed in this section, although its rocks belong to the alkaline series, because Etna is in a particular geodynamic context and it will be easier to compare the results with those of the neighbouring volcanoes of the Eolian arc or Vesuvius.

0.8

3.0Vesuvius 1

MtPetee

0.7 0.6

Merapi", I~

5.0

.

0.7020 o.7o3o o.764o o.765o 0.7050 o.7 7o 87Sr/86Sr Fig. 6. Data for subduction zone volcanoes, and ultrapotassic and associated volcanic rocks in the Sr-Th isotopic correlation diagram. All samples come from historic eruptions apart from Popocatepetl samples [34]. Data from [24], except for Nyiragongo 1977 lava (M. Condomines, unpublished results and [54]); Vesuvius [46,55]; Mt. St. Helens [35,56]; Eolian islands (VI = Vulcanello, V = Vulcano, L = Lipari, S = Stromboli) [39]; Aleutians (Bogoslof) [37,57]. The data for Etna [40] are also reported in this diagram.

251 230Th\ ~ | Th] 1.3

¢b/ . /®~/Mt. St. Helens ,~5~/C011ma , / . - " ~ v A / ALeutians , , ~ / . . . . ..,~!<~-~_.~-,j ~ - - --4;-, Paricufinz ~ - ~ / - - -~'~Cm~il~- --~ - "- . . . . . f o - / - - > 4 " , _~ ~ .~", / st.vi . . . . t i / _ . ~ - ~'_---~ o / / *

...~ -:\~/

12 //~,r (Opop . . . . 7 _ _ - ~ ' / . ' / / v I j", ../~./~-~÷-/~4,z L.-) \ - , k ~ l l l y _ . _ . , { u ~ ~

11 1.0

/,na

/ 0.9

/

2 ...-'~ B'

~ ChiliS;n- - " ~ ~_i@._ Eotian arc i

o

/ Ht'Pe[ee *~--~/ . ,~. . . .

&.,

/

/

/k=1,8

/ " .

.

.......

/

o8

~ voas ml)

.<.,v ....

/

/

/

0"Tt/ ,~,/ySakurajima I/

0.6 [I / v-

~'7~,.,.u

-

/

/

f,,ou )

/ , / 0.7

~2~-~Th} , 09

,

, 1.1

,

1.3

, 1.5

Fig. 7. Data for subduction zone volcanoes and ultrapotassic rocks (Vesuvius, Nyiragongo: N) in the isochron diagram. References as in Fig. 6, and additional samples from Poasand Arenal [24], Cascades [35], Aleutians and Marianas [37].

of the equiline, mainly come from the Mexican volcanoes (Colima, Paricutin, Popocatepetl) [24,34] or from the volcanoes of the Cascades range [35]. Both zones are characterized by a calc-alkaline volcanism emplaced on a continental crust. In each of these regions, a geographical variation is apparent, considering the (23°Th/Z3ZTh) ratios. This has been emphasized by N e w m a n et al. [35] for the southern Cascades, where the highest Th isotopic ratios are found in the northernmost volcanoes (e.g. Crater Lake), and the lowest in the southern part of the volcanic chain (e.g. Lassen peak). Newman et al. [35] propose that this variation may be related to the increasing thickness of the continental crust towards the south. For Mexican volcanoes, the Th isotopic ratios seem to decrease from Colima to Popocatepetl [24], that is from west to east, with increasing distance from the Central America trench and increasing thickness of the crust. There are however some significant variations of the (23°Th/Z32Th) ratios in the same volcano in the cases of Paricutin and Popocatepetl, which may be attributed to some crustal contamination [24,36]. It is worth noting that the data for E1 Chichon plot near the equiline. Other Central American volcanoes, like Arenal, Poas and Irazu in Costa Rica [12,24], have erupted lavas with a (238u/z3°Th) higher than 1, while

some lavas still have (238u/v3°Th) < 1 in the Irazu and Poas volcanoes. (238u/Z3°yh) ratios higher than 1 are a general feature of the volcanic rocks from oceanic island arcs like the Aleutians and the Marianas [37]. The results for two Japanese volcanoes (Asama, Sakurajima) and three Indonesian volcanoes (Krakatoa, Merapi, Galunggung), [24,38] reveal no significant disequilibrium between 238U and Z3°Th (or only a very small 23SU enrichment in the Merapi lavas). The recent volcanic rocks from Montagne Pel6e (Lesser Antilles) also are in equilibrium, whereas those from the St. Vincent Soufri6re in the same island arc are much enriched in Z38U relative to 23°Th [24]. Volcanic rocks from the Eolian arc studied by Capaldi et al. [39] display fairly constant Th isotopic ratios, but the data from Stromboli plot to the left of the equiline, and those from the other islands (Vulcano and Vulcanello, Lipari) to the right. The results on recent lava flows from Etna [40] are also reported in the isochron diagram (Fig. 7). They are very close to those of Stromboli with slightly higher Th isotopic ratios, and near the equiline, although the enrichment in 23°Th is real ((23°Th/Z38u) = 1.06-1.12). But the study of the

252 initial (23°Th/232Th)0 ratios through time in Etnean lavas (see section 5) indicates a long transfer time of the magmas, and thus the Th isotopic ratio is lower than that of the primitive magma, which probably explains the position of Etna's data below the Sr-Th correlation (Fig. 6). The interpretation is further complicated by a process of magma mixing occurring in recent years (see section 5.2). In fact the problem of the time of transfer of the magmas must always be kept in mind, when we try to interpret the 23°Th-238U (or 226Ra-23°Th) disequilibria in terms of mantle sources or melting processes. Ideally, a study of the initial ratio (23°Th/232Th)0 through time in a given volcano should be undertaken before any interpretation of the Th isotopic ratios measured in the present lavas. Although in many cases the influence of a long transfer time cannot be excluded, the consideration of both the isochron diagram and the Th-Sr correlation diagram shows that this process alone cannot explain the position of many data outside the "mantle array". As explained in section 3.2, this process can only be invoked to explain, either the position below the correlation band of volcanic rocks plotting to the left of the equiline in the isochron diagram, or the position above the correlation for volcanic rocks plotting to the right of the equiline. For the data on the equiline, the Th initial ratio can obviously have been lower or higher than the present one. Among the data reported in the Th-Sr diagram (Fig. 6) and located well outside the "mantle array", only the results from the St. Vincent Soufri6re, Lipari and Krakatoa could be explained in this way. It is worth noting that many data located on or near the equiline are also situated inside or near the "mantle array" (El Chichon, Montagne Pel6e, Asama, Sakurajima, Merapi), except for Krakatoa. It thus seems unlikely that the position on the equiline is the result of a transfer time long enough for the magma to be back to radioactive equilibrium. The apparent absence of U-Th fractionation in these magmas is not really understood at present. On the other hand, the presence of many data from subduction zone volcanoes to the right of the equiline has been interpreted [16,39,41] as the result of fluid interaction occurring before or during partial melting. U is indeed more mobile than

Th in the presence of fluids, and it is possible that " w e t " melting of a mantle source produces melts with higher U / T h ratios than the solid residue, just the reverse of what happens in "dry" melting without fluids, as in MORB and OIB. In explaining the data for subduction zone volcanic rocks, what is important also is not only the presence or absence of fluids, but when these fluids may have played a role. The important parameter is the time interval between their intervention and partial melting. If this interval is short compared to the 23°Th halfqife, and if the time of transfer of the magma is negligible, then the Th isotopic ratio will still be representative of that of the mantle source, assuming no Th with different isotopic composition is involved in the fluid transfer. In this case, the data may still be located in the "mantle array" in the Sr-Th correlation diagram (e.g. Arenal). But the time interval between the intervention of fluids and partial melting may be long enough to modify the Th isotopic composition of the mantle source. A classical model for the genesis of andesitic magmas is the melting of the "mantle wedge" above the subducted plate, this mantle wedge having been metasomatized by the fluids derived from the subducted plate. If a source in this mantle wedge has been enriched in U by fluids, its Th isotopic ratio will increase by radioactive decay, until the source eventually reaches radioactive equilibrium. Then the Th isotopic ratios of magmas formed by partial melting of such a source will be greater than that of the initial mantle source before the metasomatic event. And the data will plot above the Sr-Th correlation. Such a model is schematically represented in Fig. 8. It can well explain the high Th isotopic ratios observed in volcanoes from Mexico, the Cascades (e.g. Mt. St. Helens), or the Aleutians, which have (23°Th/232Th) as high or even higher than those of MORB. An alternative model could be the melting of an oceanic crust enriched in U by seawater alteration. If the fluids still play a role during partial melting, then the magmas will plot to the right of the equiline (Fig. 8); if not, they will be on the left-hand side of the equiline. So can be explained the difference between an island arc in oceanic environment like the Aleutians (and the Marianas) and a volcanic range on continental crust (Mexican volcanoes, Cascades). This difference was al-

253 /23°Th/ \ 232Th1 1.3

1.11.21."~9,. 0~ 0.9

~

J

"'~"'" eted Mantle

'Wet'MeLting Uenr(i~nher~ S~toofa~~semS ~....

0.8

2J. . /2,2,., crust

. ~ j -

0.6 /

I~)~23a\U I

I

I

I

I

0.7

0.9

1.1

1.3

1.5

Fig. 8. A schematic model of the effect of fluids in metasomatic processes in the mantle source or during partial melting, in subduction zone volcanism (see text for explanation).

ready noted by N e w m a n et al. [35,37]. It is possible that, under continental crust, the metasomatised mantle sources still may undergo partial melting, without any new fluid transfer from the subducted plate; because of the growth of the continent, the fluid transfer would presently only take place further towards the trench. In oceanic island arcs, the fluids from the subducted plate would still play a role during partial melting. It is also possible that the compositions of the fluids (CO2, H20, etc.) have some importance as suggested by Newman et al. [35]. It is worth noting that the proposed metasomatic processes can also affect the Sr isotopic composition, if Sr is present in the released fluids. We think that many data on volcanoes from subduction zones and, in particular, those situated above the "mantle array" in the Sr-Th correlation diagram, can be explained by more or less complex "metasomatic" processes involving the transfer of U enriched fluids in the source regions of the magmas. The data for Indonesian volcanoes (Merapi, Krakatoa, Galunggung) or Japanese volcanoes, like the Sakurajima require a different explanation. They all have very low Th isotopic ratios (~< 0.7), which seem more typical of continental crust than

of mantle values. Mixing with sediments or a crustal contamination are possible processes [24]. The Krakatoa results in the Sr-Th correlation diagram have been explained, in particular, by contamination of a mantle-derived m a g m a by subducted sediments [24]. Clearly every volcanic area in subduction zone environments has its particularity and needs a detailed discussion. This would be beyond the scope of this paper, but some of the data presented in our diagrams will indeed be discussed elsewhere (Ch. H e m o n d et al., in preparation).

226Ra-23°Th disequilibria. There are a few data on 226Ra-23°Th disequilibria on recent volcanic rocks from subduction zone environments. Capaldi et al. measured these disequilibria in historical lavas from Stromboli, Etna and Vesuvius [42,43], and from Vulcano, Vulcanello and Lipari, in the Eolian arc [39]. We also have data on Mt. St. Helens, E1 Chichon and Arenal volcanoes [31,44], and some indications on the (226Ra/23°Th) ratios in Merapi and Galunggung (Java) and Fuji and Osima (Japan) [38,45]. A study on recent lava flows from Etna [40] has provided precise (226Ra/23°Th) measurements in agreement with the previous results [42]. The values found in these volcanoes are given

254 TABLE 1 Compilation of results on 226Ra-23°Th and 23°Th-238U disequilibria in recent lavas from subduction zone volcanoes and Etna ( 226 Ra/230 Th)

Stromboli Vulcanello Vulcano Lipari Mt. St. Helens

2.50-2.90 1.18-1.39 0.95-1.32 1.05-1.09 1.27-1.37

E1 Chichon

1.31

Arenal Paricutin Merapi Galunggung, Fiji, Osima Etna

1.09 = 1.6 ---3 1.73-2.56

( 238U/230 Th ) 0.88-0.98 0.97-1.04 0.98-1.11 1.02-1.03 f 0.99-1.04 0.93-1.03 f 1.00

\ 1.04

Enrichment pattern

[42,43] [39] [39] [39] [31] [35] [31]

Ra > Th > U Ra > Th = U Ra>U>~Th Ra = U = Th Ra>Th>U Ra>Th=U

[241

f 1.00-1.08 1.11-1.23 0.79-0.89 1.06-1.08 0.89-0.94

in T a b l e 1, together with the (238U/23°Th) ratios a n d the relative e n r i c h m e n t p a t t e r n b e t w e e n U, Th a n d Ra. The (226Ra/23°Th) ratios in volcanic rocks f r o m Vulcanello, V u l c a n o a n d L i p a r i p r o b a b l y d o n o t r e p r e s e n t p r i m a r y values, but, a c c o r d i n g to C a p a l d i et al. [39], they have b e e n lowered b y r a d i o a c t i v e d e c a y while the m a g m a s were staying in crustal reservoirs. These a u t h o r s p r o p o s e that o n l y S t r o m b o l i d i s p l a y s p r i m a r y values. Thus, with the p o s s i b l e exceptions of A r e n a l a n d G a l u n g g u n g , all the volcanic rocks f r o m the o t h e r v o l c a n o e s are enriched in 226Ra c o m p a r e d to 23°Th, M e r a p i , S t r o m b o l i a n d E t n a having the highest (226Ra// 23°Th) ratios. N o d a t a show a 226Ra deficiency. It seems very unlikely that 226Ra-23°Th disequilibria can b e c r e a t e d b y f r a c t i o n a l c r y s t a l l i z a t i o n or even p a r t i a l melting. R a should then h a v e a p a r t i t i o n coefficient m u c h smaller t h a n Th, which is a very i n c o m p a t i b l e element. Its b e h a v i o u r is p r o b a b l y similar to that of Ba a n d the B a / T h r a t i o does n o t v a r y very m u c h in the first steps of fractional crystallization. But R a is a very m o b i l e e l e m e n t in the presence of fluids. T h e r e are thus two possibilities to e x p l a i n the R a excesses in m o s t s u b d u c t i o n zone volcanics: either R a is int r o d u c e d into the melt b y " m e t a s o m a t i c " fluids in the source region of the mantle. In that case, R a c o u l d be e x t r a c t e d b y the fluids f r o m a zone larger t h a n that actually u n d e r g o i n g p a r t i a l melting, as p r o p o s e d b y C a p a l d i et al. [46]. A l t e r n a t i v e l y ,

Reference

[31] [24] [24,36] [24,45] [38]

Ra = U > Th

[40]

Ra > Th > U

Ra > Th > U Ra > U > Th Ra >~U -- Th

R a - r i c h fluids f r o m the wall-rocks could cont a m i n a t e the m a g m a in the m a g m a t i c reservoir or d u r i n g its t r a n s f e r t o w a r d s the surface. T h e fact that n o v o l c a n i c r o c k showing a (226Ra/Z3°Th) r a t i o lower t h a n 1 has ever been f o u n d in s u b d u c tion zone v o l c a n o e s argues against the general v a l i d i t y of a m o d e l in which the R a w o u l d be e n r i c h e d at the t o p of the m a g m a c h a m b e r b y u p w a r d m i g r a t i o n of R a - r i c h fluids c o m i n g f r o m a d e e p e r p a r t of the m a g m a itself. It is w o r t h n o t i n g that, in a given v o l c a n o o r volcanic area, the less d i f f e r e n t i a t e d lavas seem to b e the m o s t R a enriched, as in E t n a [40] or S t r o m b o l i [39]. A s emp h a s i z e d b y C a p a l d i et al. [39], this is an a r g u m e n t for R a e n r i c h m e n t d u r i n g p a r t i a l m e l t i n g rather t h a n in a m a g m a t i c c h a m b e r . T h e e n r i c h m e n t p a t t e r n s b e t w e e n U, Th a n d R a in s u b d u c t i o n zone v o l c a n o e s are R a > Th > U o r R a > U > Th, or R a > U - - Th.

4.3. Continental ultrapotassic volcanism and associated rocks T h e d a t a o n 238u-z3°Th-226Ra d i s e q u i l i b r i a in such v o l c a n i c rocks are very scarce b u t they give very i m p o r t a n t i n f o r m a t i o n a b o u t the processes g o v e r n i n g U - T h - R a f r a c t i o n a t i o n s . Mt. Vesuvius in I t a l y was the first v o l c a n o in which disequil i b r i u m b e t w e e n R a a n d U was d i s c o v e r e d [1]. Several a u t h o r s have since r e p o r t e d m e a s u r e m e n t s of 238U, 23°Th a n d 226Ra in Vesuvian lavas. W e have p l o t t e d in the i s o c h r o n d i a g r a m (Fig. 7) the

255

most recent results obtained by Capaldi et al. [46] on historic lavas (leucite tephrites, leucitites): all the data are situated in the right part of the diagram (k = 1.03-1.41). In the same samples, the (226Ra/230Th) ratios vary between 8.2 and 10.1 (the ratios have been corrected for post-eruptive decay of 226Ra). The volcanic rocks from Vesuvius are thus clearly enriched in 238U and 226Ra compared to 230Th (Ra > U > Th). Until very recently the ( 226Ra/ 230Th) ratios in Vesuvian lavas were the highest measured in recent volcanic rocks. But in 1986 Williams et al. [20] published their measurements on the carbonatites from the 1960-1966 eruption of Oldoinyo Lengai volcano (Tanzania) in the East African Rift. They report huge 238Uthe largest 230Th and 226Ra- 230Th disequilibria, ever measured in any volcanic rock. For the true = 11.3-15.0 and carbonatites ( 238U/230Th) (226Ra/230Th) = 62.8-83.0. For the 1966 ash which looks like a mixture of nephelinite and carbonatite: ( 238U/230Th) = 1.6 and (226Ra/ 230Th) = 2.07. The (230Th/232Th) ratios of the three analysed samples show rather usual values (1.00-1.11). It is worth noting also that a disequilibrium exists in these lavas between 228Ra and 232Th (( 228Ra/232Th) = 27 at the time of eruption). The existence of disequilibria in both Ra-Th pairs allowed the authors to propose a timescale for the carbonatite magma formation (see section 5.2). Oldoinyo Lengai is the only volcano where un228Ra- 232Th disequilibria have been doubted found. There are also some recent data on the volcanic products from Nyiragongo volcano (East African rift) [47]. One group of recent nephelinites and leucitites from the main cone shows large Th-U fractionations (Th/U between 1 and 3) with data located on both sides of the equiline in the isochron diagram, while the (230Th/232Th) ratios remain fairly constant (1.31-1.43). The most recent lavas, however, do not show such a phenomenon and they all are close to 238U-230Th equilibrium [47]. This is confirmed by our result on the 1977 Nyiragongo lava flow, which is on the equiline with a ( 230Th/232Th) ratio of 1.23. In the Th-Sr isotopic correlation diagram (Fig. 6) all data for Vesuvius, Oldoinyo Lengai and Nyiragongo are above the “mantle array”. It has long been recognized that fluids have a very important role in the genesis of the ultra-

potassic continental volcanism (e.g. [48]). This is particularly evident in the case of the carbonatite magma, which is probably formed by exsolution of carbonate-rich fluids from a nephelinite magma [20]. In the three examples above, the authors agree that the fluids (in particular CO,-rich fluids) have produced the U-Th-Ra fractionations, i.e. an enrichment in Ra, and, to a lesser extent, in U relative to Th. These fluids are clearly present during the formation and evolution of the magmas; but fluids must also have played a role in the metasomatism which has affected the mantle sources [47,48], as indicated by the position of the data above the “mantle array”. In fact, the data on these ultrapotassic volcanic rocks, showing the most extreme Ra and U enrichments, provide the best argument for the importance of fluids in the fractionation processes of U, Th and Ra, and thus for the interpretation of disequilibrium results. It is an “a posteriori” justification of our statements about the possible influence of fluids in the other types of volcanism. As previously proposed [40], the role played by the fluids in a given volcano could be qualitatively evaluated in a ( 226Ra/230Th)-( 238U/230Th) diagram (Fig. 9) where there is a rough positive correlation defined by the presently available results. The importance of fluids increases with increasing ( 226Ra/230Th) and ( 238U/230Th) ratios from the still hypothetic (due to the lack of Ra

i-4 226Ra 230Th

i

Fig. 9. ( 226Ra/23” Th) versus ( 238U /230Th) diagram in historic volcanic rocks. The role of fluids increases as the ( 22hRa/23”Th) the and ( 23xU/230Th) ratios increase. Domain A represents possible field of MORB and OIB. References as in Fig. 6, and Mt. St. Helens, El Chichon, Arena1 [31], Merapi [24,38].

256 data) domain of MORB and OIB to the extreme domain of carbonatites. In a " d r y " environment the fractionation pattern would be Ra = Th > U, in a " w e t " one Ra >> U > Th and in intermediate situations Ra > Th >~ U. Although, as explained above, there are some arguments to think that these Ra (and U) enrichments can occur during partial melting of the mantle sources, there is no definite proof and the possibility cannot be excluded that the enrichments take place in crustal reservoirs. Many more data will be necessary before a clear answer can be given to this problem. Above all, specific and careful studies, on a given volcano, of the variations of (23°Th/232Th)o and (226Ra/23°Th)o ratios through the volcanic history should help solving these questions. It is the purpose of the following section to show the usefulness of radioactive disequilibria for the study of a single volcano, its magmatic evolution and m a g m a dynamics.

5. Transfer time and magma dynamics A classical application of the usual isotopic tracers (Sr, Nd, Pb, etc.) is the study of the evolution of terrestrial reservoirs (mantle, continental crust, seawater, etc.), by looking at the variations of the initial isotopic ratios in certain categories of rocks. The short-lived isotopes of the U and Th decay series in volcanic rocks can be used in the same way to follow the magmatic evolution and m a g m a dynamics in a single volcano [12,13]. Depending on the time constants of the phenomena we want to study, different isotopes with appropriate half-lives can be chosen in the decay series, covering a timescale from a few days to some 3 × 10 5 years. Some short-lived isotopes such as 222Rn ( T = 3.8 days), 21°Bi ( T = 5 days) and 21°Po ( T = 138 days) have been successfully used in studies of m a g m a outgassing and volatile emission (e.g. [49-51]), and they give fruitful information about these processes. We will only concentrate in this paper on isotopes of "long" and intermediate half-lives: 23°Th, 226Ra, 21°Pb in the 238U series, 2ZSRa and 228Th in the 232Th series. 5.1. (23°Th/23:Th)o ratio variations and magmatic evolution The possibility of a non-negligible time of transfer of the magmas from their formation by

partial melting to their eruption has been mentioned m a n y times in the preceding discussions. If the transfer time is in the range 1 0 4 - 1 0 6 years, this possibility can be tested by considering the evolution through time of the initial (23°Th/ 232Th)0 ratios in the volcanic rocks of a single volcano. These ratios are reported in a (23°Th/ 232Th)0vs. e - a t diagram, where t is the age of the volcanic rocks: this "isotopic evolution diagram" is similar to the 87Sr/86Sr vs. t diagram. The principle is simply based on the fact that in a closed system the (23°Th/232Th) ratio of the m a g m a will decrease or increase through time, according to the initial position of the m a g m a to the left or to the right of the equiline in the isochron diagram (Fig. 10). The radioactive decay equations show that the Th isotopic ratio variation is linear when plotted versus e x' [13]. If the time of transfer of the magmas is negligible compared to the 23°Th half-life, and if they are derived from a homogeneous mantle source with constant T h / U ratios, then the (23° Th/232 Th) 0 initial ratios should not vary with time during the history of the volcano (which means that the volcano has been fed by successive batches of m a g m a of similar Th isotopic composition). It is interesting to note that these two simple models (closed system evolution and negligible transfer time) correspond probably to different structures of the magmatic system under the volcano, mainly to the presence or absence of a deep permanent reservoir. Indeed if m a g m a stays at depth for 10 5 years or more, it implies a reservoir deep and large enough to prevent magma solidification by conductive cooling. Fig. 11 illustrates this concept in a very schematic way. In the case of closed system evolution, the residence time in the deep reservoir can be calculated, if we know the Th isotopic ratio of the m a g m a at the beginning of the evolution. This ratio can be estimated from the Sr-Th isotopic correlation, if the deviation from the mantle array is only due to the long time of transfer of the magma. Although this kind of study is very powerful, it has been applied to very few volcanoes until now, mainly because of the lack of precisely dated volcanic rocks covering the entire span of volcanic activity. The first studied case was the Irazu volcano (Costa Rica) [12]. Th data seemed to fit rather

251

Fig. 10. Variation of initial (z30Th/232Th)o ratios as a function of e A’ in a closed system evolution of a magma in its reservoir. Subscript M refers to the magma, MT to the initial value of the magma formed by partial melting. (a) (238U/230Th)M > 1; (b) (238U/230Th)M < 1.

well a linear variation in a (230Th/232Th)o vs. ehr diagram. More detailed was the study on Etna [13]. The variation of the initial Th isotopic ratios is more complex than in the previous example, and is interpreted as the result of closed system evolution of an alkaline magma in a deep reservoir for at least 150,000 years, with periodic injections into this reservoir of magmas of tholeiitic affinities and lower Th isotopic ratios. One of the main results of this work was to show that magmatic evolution occurs at two different levels corresponding to two different timescales: slow evolution in the deep reservoir for lo5 years or more, and more rapid evolution (< lo4 years) in the shallow reservoir(s), where the main differentiation by crystal fractionation takes place [13]. In contrast with those results, the data obtained by Newman et al. [21] do not show any significant

variation of the Th isotopic ratios in lavas from Mauna Kea (Hawaii) and Marion Island (Prince Edward hot spot) for the last 280,000 years. In these cases, a short transfer time of the magmas is indicated by the results. A rather similar situation is found for the Piton de la Fournaise volcano on Reunion island (Indian Ocean). This study is still in progress and the detailed results will be published elsewhere (Condomines et al., in preparation). But the first results indicate no significant trend in Th isotopic ratios variation through time for the last 170,000 years (Fig. 12). The small differences in ( 230Th/232Th)o ratios can be attributed to small source heterogeneities. This is confirmed by the slight difratios of the studied flows. ferences in a7Sr/%r Moreover, all the data from the Piton de la Fournaise are located inside the “mantle array” in the

258

1

,

e

b)

Fig. 11. Possible relationship between two simple models of Th initial isotopic ratio evolution and the structure of the deep “plumbing system” of the volcano. (a) Short transfer time of the magmas and successive reinjections = no permanent deep reservoir. (b) Closed system evolution with long transfer time = presence of a deep long-lived reservoir.

Sr-Th isotopic correlation diagram (Fig. 4) in agreement with other MORB and OIB data. Note that the Etna data plot below the “mantle array” as expected if the magma has stayed for a long time at depth (Fig. 6). It thus seems that the Piton de la Fournaise has no long-lived deep reservoir. The very constant and rather primitive chemical

t 0.6

0

I I / I 10 1

50

100 2

3

Fig. 12. Variation of initial ( 230Th/232Th)o tion of e” for the Piton de la Fournaise domines et al., in preparation).

typlcol error

bar

150

r10'y t

4

e Al

ratios as a funclava flows (Con-

composition of the transitional basalts from the Piton de la Fournaise is in agreement with this model. It is rather satisfying that the few detailed studies on hot-spot volcanoes in oceanic areas (Mauna Kea, Marion, Piton de la Fournaise) all indicate a short transfer time of the magmas towards the surface. This reinforces our earlier assumption that the MORB and OIB data defining the “mantle array” represent Th/U mantle values, the Th isotopic ratios being unaffected by a long transfer time of the magmas. It may be significant that only volcanoes located in continental (or intermediate: Costa Rica?) areas appear to have a long transfer time of their magmas and deep long-lived reservoirs. Perhaps the presence of a thick continental crust allows the formation of deep reservoirs at the crust/mantle boundary or in the lower crust. Testing this hypothesis will obviously have to await further detailed studies on individual volcanoes in different geodynamic environments. 5.2. Ra-Th disequilibria and magma dynamics The use of Ra-Th disequilibria to date magmatic processes was first proposed by Capaldi et al. [42] in a pioneering paper. The study, with of magma dynamics in active these isotopes, volcanoes is theoretically possible in a time span from a few years to lo4 years. The interest of the Ra-Th disequilibria is increased by the fact that we can use two isotopes: 226Ra and **’Ra (see Fig. 1) with respective halflives of 1600 years and 5.77 years. Thus the radioactive couples 230Th- 226Ra and 232Th- ***Ra will reach equilibrium at different rates (about 8000 and 30 years, respectively): it is then possible to constrain the time of the Ra-Th fractionation in the magma by measuring the disequilibria in both pairs. This can be illustrated in a very schematic way: if both ( 226Ra/230Th) and ( 22*Ra/232Th) ratios equal 1, we must conclude that either no Ra-Th fractionation has occurred in the magma or that this fractionation is older than about 8000 years. If only (226Ra/230Th) is different from 1, the fractionation took place between 30 years and 8000 years ago. And if both (226Ra/230Th) and (228Ra/232Th) differ from 1, it means that the Ra-Th fractionation is less than 30 years old. In that case, it is possible with some assumptions on

259

the initial state of the magma, to date precisely this fractionation (assuming an instantaneous process). This has been beautifully done by Williams et al. [20] for the Oldoinyo Lengai carbonatites. If the measurements are undertaken quickly after the eruption, the 228Ra-228Th disequilibria can give additional information as shown by Capaldi et al. [42] (228Th has a half-life of 1.9 years). Unfortunately the only volcano where undoubted 228Ra232Th disequilibria have been found is the Oldoinyo Lengai mentioned above. The early resuits by Capaldi et al. [42] indicated such disequilibria in lavas from Stromboli and Etna. But some other volcanic rocks from recent Stromboli eruptions do not show any 228Ra-232Th disequilibrium [43]. And our results on recent Etna lava flows demonstrate equilibrium between 232Th, 228Ra and 22STh [40]. It is fair to say that very few measurements are available on 232Th-228Ra disequilibria in present volcanic rocks. Apart from the above mentioned results, there are only four results on pyroclastic samples of the 1980 eruption of Mt. St. Helens [44], again showing 232Th-228Ra-228Th equilibrium. If this is confirmed in other volcanoes, where 226Ra-23°Th disequilibria exist, it would mean that the time interval between m a g m a formation and eruption is always greater than 30 years. Moreover, as 21°pb (T = 22 years) generally appears to be near equilibrium with 226Ra [31,40,44], this time interval is probably greater than 100 years. Our results on six recent lava flows (1950-1983) from Mr. Etna [40] also show that the Ra-Th fractionation must have taken place between 100 and 8000 years ago. Moreover, there are large variations in the (226Ra/23°Th) ratios between successive eruptions (from 1.73 to 2.56). These ratios are very well correlated with the Th content of the lavas, which is an index of differentiation. The highest ( 2 2 6 R a / 2 3 ° T h ) ratio is found in the 1974 basic magma, with the lowest Th content, and the lowest (226Ra/23°Th) in the more differentiated 1950 lava with a higher Th content. The very good linear correlation in a (226Ra/ 2 3 ° T h ) - 1 / ( 2 3 ° T h ) diagram (Fig. 13) is interpreted as the result of a mixing process between a deep basic magma, more Ra-enriched, and a shallow, differentiated, and less Ra-enriched m a g m a in the upper part of the " p l u m b i n g system" of the volcano [40].

9+/

3OThJ 2.5

1.5

~]]° ' ./.5

]50

.55

h 1/123OTh} .60

Fig. 13. Variation of (226Ra/23°Th) versus 1/(23°Th) in recent Etna lavas [40] and a schematic cross-section of the volcano (inset), parallel to its rift zone [58], showing the mixing between the two magmas (open and filled circles). The magma of the 1974 excentric eruption was tapped from a deep part of the magmatic system (see text for explanation).

We are now extending this study to the whole historic period of Etna's activity. This should help to understand the process of Ra enrichment and its kinetics (instantaneous or continuous process). It is worth noting that the most Ra-enriched lavas of the recent eruptions (1974, 1978, etc.) also exhibit significant excesses in K, Rb and Cs [28]. The latter have been attributed to selective contamination of the m a g m a by the wall-rocks of the sedimentary basement [28,52]. It is quite possible that the additional Ra excesses in these lavas from the last eruptions are due to this mechanism, although it should have occurred in the m a g m a more than 100 years ago and not between successive eruptions (228Ra-Z32Th and 21°pb-226Ra are in equilibrium). But Ra excesses are present in older lava flows which do not show any K, Rb or Cs anomaly, and thus the ultimate origin of the ubiquitous Ra excess in Etna cannot be accounted for only by the above process. Indeed our first results indicate that three lava flows from the 14th and 17th centuries have (226Ra/23°Th) ratios (calculated at the time of the eruption) similar to those found in the 1950 or 1964 laval flows. If the (226Ra/23°Th) has remained nearly constant for several centuries of Etna's activity, the m a g m a cannot have evolved in a closed system, with respect to Ra. Either successive reinjections of

260 magmas with constant (226Ra/23°Th) ratios occurred during this period; or the Ra enrichment is a continuous process affecting the m a g m a in its deep reservoir (for example, continuous migration of Ra-rich fluids from the wall-rocks), the (226Ra/ 23°Th) ratio reaching an equilibrium value higher than 1. The former hypothesis is in apparent contradiction with the conclusion derived from the study of (23°Th/232Th)o ratios, which favours the existence of a deep permanent reservoir for more than 10 5 years. If this reservoir evolves as a closed system with respect to Ra, 226Ra should be in radioactive equilibrium with 23°Th. This contradiction is only avoided if some Ra is introduced into the deep reservoir or at a higher level. It does not exclude the possibility that the primitive m a g m a was itself enriched in Ra during its formation by partial melting. Any further discussion would be premature but these results on Etna, as well as the spectacular data obtained by Williams et al. [20] on the Oldoinyo Lengai carbonatites, demonstrate the interest of Ra-Th disequilibria in the study of m a g m a dynamics in active volcanoes. 6. Summary and conclusions We would like in conclusion to underline the importance of the study of U - T h - R a radioactive disequilibria in magmatic processes and to summarize the main results obtained so far in this domain. - - The radioactive disequilibria allow us to study fractionation between the elements in a rather unique way, because the initial state of radioactive equilibrium can be assumed in some cases. They also can give time constraints on the problems of the genesis, evolution and transfer of magmas towards the surface. In the U and Th decay series, a number of nuclides can be used, with half-lifes from few days to more than 10 4 years. According to the timescale of the studied magmatic phenomena, one particular nuclide can thus be chosen in the series. - - The (23°Th/232Th) ratios can be used as isotopic tracers to infer the T h / U ratios of the mantle sources. The negative correlation between Th and Sr isotopic ratios in recent M O R B and OIB [15] is now substantiated by new results, defining a " m a n t l e array" similar to that observed in the Nd-Sr isotopic correlation diagram. The

Th-Sr correlation is a consequence of coherent T h / U and R b / S r fractionations due to partial melting events in the mantle sources. The time of transfer of the M O R B and OIB magmas from their sources to the surface is probably less than 10 4 years. - - The 238U-23°Th disequilibria are most probably created by partial melting of the mantle sources. M O R B and OIB display (238U/23°Th) ratios lower than 1, whereas volcanic rocks from subduction zones have (238U/23°yh) ratios both lower and higher than 1, and sometimes equal to 1. The few available data on 226Ra-23°Th disequilibria show that M O R B and OIB may be characterized by Ra-Th equilibrium or small Ra excess. In contrast m a n y volcanic rocks from subduction zone volcanoes show large Ra excesses. This fact, combined with the 238U enrichment over 23°Th in some of these volcanoes, favours a model in which fluids play an important role in the genesis of this volcanism, before partial melting (metasomatised mantle sources) a n d / o r during the melting event, or even later. - - The few results on U - T h - R a disequilibria in ultrapotassic and related volcanic rocks (carbonatites) confirm the role of the fluids in producing Ra and, to a lesser extent, U enrichments in the magmas. A rough correlation between (238U/ 23°Th) and (226Ra//23°Th) ratios in recent volcanics may indicate an increasing role of fluids in samples with high ratios. In a " d r y " environment, the enrichment pattern is R a - - T h > U, in a " w e t " one Ra > U > Th, and in intermediate situations Ra > T h - - U . The intervention of fluids can take place during partial melting or when the m a g m a is stored in a magmatic reservoir. --The variation of (23°Th/232Th)0 ratios through time in a single volcano permits study of the magmatic evolution and transfer time of the magmas which is related to the deep structure of the " p l u m b i n g system'. Volcanoes from oceanic islands (Fournaise, Marion, Manna Kea) appear to have short transfer times ( < 10 4 years) in agreement with their position inside the " m a n t l e array". Other volcanoes (Irazu, Etna) show long transfer times of their magmas ( >/10 5 years). This may be related to the presence of a continental crust. - - The Ra-Th disequilibria give additional information on m a g m a dynamics in active volcanoes

261

on a timescale from a few years to 10 4 years. From the very few available results, it appears that 228Ra is always in equilibrium with 232Th (and 21°Pb with 226Ra), except for the Oldoinyo Lengai carbonatites, which means that the transfer time of the magmas is longer than 100 years. The results on Etna demonstrate that the Ra enrichment is probably characteristic of the whole historic period and is a deep-seated phenomenon occurring in a deep reservoir a n d / o r in the zone of partial melting.

Acknowledgements We are grateful to J.D. MacDougall and F. Begemann for their careful reviews of the manuscript.

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