Iodine abundances in oceanic basalts: implications for Earth dynamics

Earth and Planetary Science Letters, 108 (1992) 217-227

217

Elsevier Science Publishers B.V., Amsterdam [UCl

Iodine abundances in oceanic basalts: implications for Earth dynamics B. D 6 r u e l l e a, G . D r e i b u s b a n d A . J a m b o n a Laboratoire de Magmatologie et Gdochimie lnorganique et Expdrimentale, Unit,ersitd Pierre et Marie Curie, 4 Place Jussieu, F-75252 Paris Cddex 05, France h Abteilung Kosmochemie, Max-Planck-lnstitut fiir Chemic, Postfach 3060, D-6500 Mainz, Germany

Received June 25, 1991; revision accepted November 13, 1991

ABSTRACT Iodine analyses by neutron activation have been performed on 32 oceanic basalts glasses, 1 phonolite and 3 subaerial arc basalts. The world-wide sample set encompasses all typical geodynamic settings [ridges (MORB), oceanic islands (O1B), arcs (IAB) and back-arc basins (BABB)] and the diversity of oceanic basalt types [depleted (N), intermediate (T) and enriched (P)]. Most basalts, including all N-types, all but one T-type, and some P-types, exhibit low iodine concentrations (2.5-13 ppb). Very high I concentrations (up to 363 ppb) in a small number of samples (all P-types) are interpreted to be the result of a recycled component which includes organic matter of sedimentary origin (sediment organic matter accounts for about 80% of total terrestrial iodine). Iodine appears to be the most incompatible element after the noble gases. Mass balance considerations permit the mantle iodine concentration and hence the bulk silicate abundance of the Earth to be constrained to 9-24 ppb, with a preferred value of 10 ppb. The low terrestrial iodine abundance, coupled with a chondrite-like chlorine/iodine ratio, strongly favours the late veneer model of Earth accretion. The scenario proposed to explain the terrestrial iodine distribution includes heterogeneous accretion, iodine extraction from the mantle simultaneous with (or even before) continent formation, and depleted mantle homogenization. As iodine is not recycled into the mantle by oceanic crust, heterogeneities in the mantle should result from organic sediment (C) recycling,of which iodine may be a good tracer.

1. Introduction A geochemical a p p r o a c h has b e e n used for decades in o r d e r to u n d e r s t a n d the evolution of the E a r t h (see n u m e r o u s references in [1]). O n e very clear result of the previous work is that the c o m b i n a t i o n of several tracers such as isotopes a n d trace e l e m e n t s p e r m i t s a b e t t e r d e l i n e a t i o n of constraints o n E a r t h geodynamics. U s i n g tracers that display distinct chemical properties is particularly efficient a n d allows new i n f e r e n c e s (e.g. the n o b l e gases have p e r m i t t e d a b e t t e r u n d e r s t a n d i n g of the early history of the E a r t h a n d its outgassing [2]). Similarly we have b e e n a t t e m p t i n g for several years [3,4] to c o m p l e m e n t data for volatiles with chemical species that have specific behaviour; halogens fall into this category of elements. This study was motivated by two major aspects. As iodine at the E a r t h surface is specificaIly c o n c e n t r a t e d in sediments, this eleElsevier Science Publishers B.V.

m e n t is a p o t e n t i a l tracer of s e d i m e n t recycling into the m a n t l e , like Li, B a n d ~°Be [5]. A very different aspect is 129I (with a half-life of 17 Ma), which is the p a r e n t of 129Xe, a n d as 129Xe a n o m a lies have b e e n discovered on E a r t h a n d in meteorites, the study of the I / X e system has yielded i n f o r m a t i o n o n the early stages of terrestrial formation. M o d e l l i n g of this p r o b l e m is, however, limited by i g n o r a n c e of terrestrial iodine a b u n d a n c e a n d its d i s t r i b u t i o n a m o n g the potential reservoirs. M o r e t h a n 90% of the surface iodine resides in s e d i m e n t s [6-8], a n d not in the c o n t i n e n t a l crust or oceanic crust. Limited data on iodine for volcanic rocks a n d m a n t l e - d e r i v e d xenoliths have previously b e e n f o u n d in the literature [9-11], but all estimates of the geochemical d i s t r i b u t i o n of iodine in igneous a n d s e d i m e n t a r y rocks [e.g. 12,13] refer to the same a n c i e n t m o n o g r a p h [14]. M o r e r e c e n t iodine m e a s u r e m e n t s on four

218

spinel-lherzolite nodules (from < 2 to 40 ppb) suggest a low mantle I content [15]. Using this very small number of ultramafic nodules to obtain the mean mantle concentration of halogens is questionable however, as iodine has probably been extracted from the mantle by liquids during partial melting and may not reflect the concentration of the mantle. Furthermore, several nodules have been shown to be affected by metasomatism, which could have strongly altered the concentration of trace elements. We believe that analyzing iodine in diverse types of oceanic basalt glasses is a more valid method for obtaining mantle iodine estimates because our knowledge of the relationship between basalts and the mantle is enhanced by detailed trace element and isotopic data [16],

2. Sample selection This work reports analyses of iodine (Table 1) in 32 glasses of oceanic basalts, 1 phonolite and 3 island-arc basalts (IAB). The samples were selected according to various criteria and the following characteristics: (1) The glasses are all fresh, as demonstrated by their unaltered aspect in thin sections. In cases

B. Df~RUELLE ET AL.

when olivine phenocrysts were present in the samples only the glassy matrix was analyzed. (2) They come from a world-wide selection of typical oceanic geodynamic settings (Fig. 1 ) - - t h e Mid-Atlantic Ridge, the East Pacific Rise, the Red Sea, the Lau back-arc basin, Fiji, the Marianas basin, Iceland and H a w a i i - - a n d represent basalts from mid-oceanic ridges (MORB), oceanic islands (OIB) and back-arc basins (BABB). Three subaerial samples from the Lesser Antilles representing an IAB were also analyzed. The number of samples from each location is by no means representative of their terrestrial occurrence, a point that will be returned to in the discussion. (3) Major and trace element concentrations are known in all the samples [4,17,20-24,27,29, 31,32]. Major volatile concentrations ( H 2 0 [4], CO 2 [33]) are known in all the oceanic samples. (4) They are typical basalts, as demonstrated by their overall composition, their Mg-number (0.46-0.70) and differentiation index [34] (17.233.8), with the Hawaiian samples being the most evolved. (5) On the basis of their K contents (Table 1), ( L a / T i ) N, ( L a / S m ) N or Z r / N b ratios [35], the oceanic samples can be conventionally divided into normal or depleted (N), plume or enriched

Fig. 1. Geographical distribution of oceanic basalts listed in Table 1. One phonolite (Society Islands) and three island-arc basalts (Lesser Antilles) are also indicated.

I O D I N E 1N O C E A N I C

BASALTS AND IMPLICATIONS

FOR EARTH

219

DYNAMICS

TABLE 1 Potassium and iodine data for the studied samples. P-, T- and N-types and IAB (see text). K data from [4] and measured by electron microprobe (CAMEBAX). Error (%): counting statistic + analytical error derived from standard deviation Sample

Type

ppm K

ppb I

I error (%)

Reference

P P P N N N N N

2530 5280 1825 805 540 480 700 790

6.5 363 49 12 11 6.3 8.4 5.1

17 5 5 10 10 16

[17] [17,18] [19] [17,18] [17,18] [17,18] [17,18] [17,18]

P

2655

7.3

13

[4]

P P P

7070 4015 5345

164 32 40

5 5 5

[20] [20] [20]

P P P P

5105 5790 5700 3595

61 27 20 8.1

5 5 10 15

[21,22] [21,22] [21,22] [21,22]

East Pacific Rise 981R26 CY82 18-01 CY82 31-02 CL D R 1 - 5 V CY82 27-01 CY82 09-03

N T N N T P

414 1485 620 605 1825 3570

2.5 8.1 13 6.6 4.4 9.4

22 11 10 15 18 15

[23] [24] [25] [26] [24] [24]

Society Islands Rocard P 4 - 3

(P)

47630

20

12

[27]

Marianas M A R A 04DR9 M A R A 07 10P

T T

3195 1775

26 6.6

9 12

[28] [28]

Lau Basin SO48-2 7GC SO48-2 42GC S O 4 8 - 2 18GA2

N N N

325 440 505

5.7 6.4 7.2

13 17 15

[29] [29] [29]

Fiji ST15-11

T

1305

5.7

16

[30]

Red Sea G M R KS04-A G M R DR07-B G M R KS12

T T T

2665 2930 1815

10 7.8 8.5

12 13 11

[31] [31] [31]

Lesser Antilles V 71008 VST79-70 VSTK 106

lAB lAB lAB

4815 4400 2655

11 13 27

15 15 5

[32] [32] [32]

Atlantic CH CH CH CH CH CH CH CH

Ocean 97DR02 97DR05 31DRI 1 98DRI0 98DR11 98DR12 98DR15 98DR17

Iceland IC ST1-89 Pacific Ocean

Loihi KK17-2 KK 20-14 KK 25-4 Kilauea ERZ ERZ ERZ ERZ

1712 1701 1699 1697

220 (P) and transitional or intermediate (T) types. The use of these criteria is consistent for all but one sample with a high potassium content ( M A R A 04DR9, K = 3195 ppm), and this sample is provisionally designated as a T-type. (6) All the M O R B , BABB and OIB samples (with the exception of the Icelandic sample) are of deep oceanic origin and were erupted at depths greater than 900 m, which prevents significant volatile loss effects [36]. The water contents of the samples leads us to suggest that they have retained their volatiles [4]. The Icelandic sample is from a subglacial pillow (the ice cap was probably more than 500 m thick at the time of eruption) and displays a normal water content, again suggesting retention of volatiles. In contrast, the IAB samples (Lesser Antilles) erupted under atmospheric pressure and may have lost some of their volatiles [37]. IAB glasses from deep oceanic origins ( > 1000 m) are not available at present. (Volatile loss at shallow depths has been proposed previously as a major factor controlling the volatile C1 and Br concentrations in oceanic basalts [36]. This possible loss effect is not documented in the case of iodine and will be addressed below. The four Kilauea samples are the only ones with some rough correlation between depth and iodine content, but the samples from nearby Loihi do not follow the trend.). (7) To obtain information on iodine concentration in the mantle rather than the behaviour of iodine during magmatic differentiation and partial melting, we analyzed mostly oceanic tholeiites, for which comparable degrees of partial melting are expected and fractional crystallization effects are a minor phenomenon.

B. D[~RUELLE ET AL.

tively extracts iodine and all other halogens which are in turn trapped in a sodium hydroxide solution. Wet chemical iodine separation from chlorine and bromine with a CC14-extraction of 12 is then performed for subsequent y-counting. This method is specifically designed to yield the abundance of all the halogens on one single sample, and especially iodine, at extremely low concentration levels. The iodine detection limit is 2.5 ppb. The precision varies from 20% through 10% to 5% for concentrations of 2.5 ppb through 10 ppb to 300 ppb respectively (Table 1). The isotope dilution mass spectrometry method [39] has been checked with the results obtained by R N A A [40,41]. Further details and results for the other halogens will be reported elsewere. 4. Results The analytical results are reported in Table 1. Iodine concentrations exhibit a very large range, from 3 to 363 ppb (Fig. 2). A range of more than two orders of magnitude is unusual for trace elements in oceanic tholeiites, and K concentrations in the same set of samples vary by a factor of 18. The great majority of samples exhibit low iodine concentrations: 24 of 35 samples contain less than 13 ppb iodine and only 2 samples have concentrations in excess of 70 ppb. In detail, it appears that all N-type basalts exhibit low iodine concentrations (mean = 7.7 ppb, o- = 3.2, 11 samples) and the T-type basalts are indistinguishable from N-types (mean = 7.4 ppb, o-= 1.8, 7 samples). This limited variability is comparable to that observed for moderately incompatible elements such as H R E E [42]. One characteristic of

3. Analytical method For each sample, we selected several unaltered, glassy chips. After crushing, coarse grains from these distinct chips were picked, reduced to powder and irradiated. The analytical technique [38] used in this study was developed at the Max-Planck-lnstitut fiir Chemie in Mainz and is based on neutron activation analysis (RNAA). Briefly, pyrohydrolysis of irradiated samples is combined with a wet chemical iodine separation. Irradiated (active) samples are admitted to a stream of water vapour at 1100°C which quantita-

N

i1'

,DDDD,~O,,,iI.,,O,,,I,,,iI~ ,

0

20

,I

40 4B 60 164 362 i o d i n e ppb

Fig. 2. Distribution of iodine concentrations among oceanic basalts of the N-, T-, and P-types (see text). One phonolite

(R) and three island-arc basalts (lAB) are also indicated. Most samples cluster around the low value of 8 ppb. Some P-type basalts display very high I contents. N = number of samples.

IODINE

IN

OCEANIC

BASALTS

AND

IMPLICATIONS

FOR

EARTH

incompatible elements is their significant enrichment in the continental crust and their depletion in the N-type basalts. It has also been shown that the variability of strongly incompatible elements becomes even more pronounced as their incompatibility increases. The iodine concentration variability among P-type basalts (mean = 63 ppb, ~r = 96, 13 samples), which have a very large range (7-363 ppb) completely overlaps the range for Nand T-types. This is similar to and even more notable than observations for the most incompatible elements such as Cs (the variability of which is about three times that of K in P-type basalts [42]). To summarize, iodine seems moderately incompatible when N-type basalts are considered; T-type basalts confirm this moderate behaviour of iodine, but P-type basalts point to an exceedingly large degree of incompatibility, in agreement with the strong crustal enrichment. 5. Interpretation and discussion Among the processes that can induce significant variability in trace element concentrations are those which affect the composition of the source material (e.g. recycling and metasomatism) [43], and those which affect the melt (degassing upon eruption or later, contamination, fractional crystallization and partial melting). This has been discussed in the case of halogens for basalts of the North Atlantic [44]. This aspect has also been discussed at length [4,42] for a set of samples including that studied here.

5.1. Degassing One problem of major concern is the possible partial degassing of iodine upon eruption. Iodine may be carried by major volatiles (CO 2 and H 2 0 ) when bubbles escape from the melt. It has been shown from the solution properties and natural abundances of these two species that CO 2 escapes first, whereas H 2 0 does not escape significantly as long as the pressure exceeds 50 bar [45]. CO 2 concentrations in glasses are highly variable because of different depths of eruption and possible significant supersaturation of the melt [46]. The three Loihi samples have experienced extensive CO 2 loss but iodine concentrations are high

221

DYNAMICS

depth (m) 5000

i

:.K

i

0 1K

4000 1-113

1K

30oo IL

2000

•

1000

[ ] N-type 0 T-type • P-type

• L

0 0

x

• 0

i i~

0

i~:~,

10

i

20

. ~

i

30

.

i

40

,

lAB

Phonolite i

50

,

i

60

,

70

I ppb Fig. 3. A plot of depth vs. iodine concentration for all basalt samples (and one phonolite) displays no correlation at all. This is a strong argument against a significant loss of iodine at shallow depth. K = Kilauea; L = Loihi.

and vary from 32 to 61 ppb in these samples. The N-type basalt samples, none of which is strongly degassed, as shown by their CO 2 and H 2 0 contents, all exhibit low iodine concentrations which cannot result from significant CO 2 degassing, and certainly not from H 2 0 degassing. Clearly, CO 2 escape does not control iodine concentration. Still, if iodine concentration in oceanic samples is not controlled by degassing, this process might have affected the subaerial samples. The Icelandic sample (ST1-89) is a glass from a subglacia[ pillow erupted under unknown pressure. Its high water content testifies that iodine was not affected by degassing of this sample. For the IAB samples, the measured iodine concentration may represent a residue from incomplete degassing, but we have no evidence of any iodine loss. As the concentrations (11-27 ppb) are not particularly low, we will not discard them until further evidence is provided. The effect of degassing with decreasing oceanic depth has been documented for CI and Br [36]. For the present set of samples, neither iodine (Fig. 3) nor the iodine/potassium ratio correlates with depth, which suggests that degassing of iodine did not occur. The presence of both C1 and I in lunar rocks [47,48] strongly suggests that these elements (unlike water) are not very volatile when incorporated in silicate melts. This might be the

222

B. DIZ'RUELLE ET AL.

2O7pb / 2O4pb 15.7 [] 0 • 15.6 •

15.5

istics that would entrap iodine. The high iodine content in sample C H 97DR05 (363 ppb) can only be explained by incorporation of organic sediments at some stage. Incorporation through assimilation at a late stage however seems very unlikely.

v

N-type [ T-type • P-type Phonolite

[]

," []

15.4

10

100

1000

I ppb Fig. 4. 2 ° 7 p b / 2 ° 4 p b vs. I for o c e a n i c samples. 2 ° 7 p b / 2 ° 4 p b is used as a t r a c e r of c o n t i n e n t a l crust m a t e r i a l (sediments). No c o r r e l a t i o n a p p e a r s , which suggests that iodine and lead carriers are not the same even t h o u g h s e d i m e n t s may be involved in both cases. L e a d data after [24,27,49-52].

result of low mobility of iodine, which limits its reaction kinetics.

5.2. Contamination One reason for choosing glasses is to avoid contamination of the samples. Only the IAB samples are not glasses (such samples were not available). Any late stage contamination would affect the samples independently of their chemical signature and not specifically P-type basalts; none of the N-type samples display iodine concentrations in excess of 13 ppb. Besides, high iodine contents in some P-type samples are not the result of contamination, as most of these samples do not display the signature that is normally expected when using the 2°7pb/2°4pb ratio as a tracer of continental crust material (Fig. 4). The most significant reservoir of iodine at the Earth surface is organic matter contained in sediments; sediments can thus carry large amounts of iodine (on the scale of parts per million or more). Incorporation of sediments in samples erupted at ridges is unlikely because of the general absence of sediments at the ridge crest itself. Incorporation of altered oceanic crust in the source is a more serious possibility, but altered crust is not expected to be a carrier of iodine as no alteration mineral displays the crystallochemical character-

5.3. Composition of the source Radiogenic isotope ratios in oceanic basalts have been interpreted as resultong from mixing of four end-members [1]: a depleted M O R B mantle component (DMM), a mantle with a high U / P b ratio possibly resulting from oceanic crust recycling (HIMU), and two enriched mantle components, a slightly modified bulk Earth component (EMI), and a component from pelagic (subducted) sediment recycling (EMII). If this is correct, it should be valid for trace elements as well. But concentration alone is not a very good signature of the source as it may be significantly affected by partial melting and fractional crystallization. Furthermore, D M M can be depleted to various degrees and isotope ratios are not affected by these processes. The same is true for the ratios of trace elements with the same incompatibility. For instance, the ( L a / S m ) N and K 2 0 / H 2 0 ratios of N-type basalts display limited variations, while the ratios of P-type basalts may be up to about five times higher [4]. As for radiogenic isotopes (and trace element ratios), it is more realistic to consider what possible mixture of the above components can explain the composition of the samples and to see whether the iodine variations can be explained within this framework. N-type basalts all display low iodine concentration and I / K ratios from 0.7 to 2.0 x 10 5 (Table 1). One cannot really consider the I / K ratio as constant as the K variations (a factor of 2.5, clearly outside the errors bars) are not well correlated with iodine variations which extend over a similar range (Fig. 5). When T-type samples are considered, the I / K ratio is extended to lower values as the iodine concentrations cover a similar range as for the N-types, despite higher K contents. It can be concluded that partial melting, crystal fractionation and degassing are not sufficiently efficient to produce the very strong variations in the iodine concentrations measured in

IODINE IN OCEANIC BASALTS AND IMPLICATIONS FOR EARTH DYNAMICS

I ppb

where C is iodine concentration, m the mass fraction of the reservoir, and BE, c, d, e and p refer to bulk Earth, crust, depleted, enriched, and primitive mantle respectively. As the concentration in all these reservoirs is not accurately known, an exact estimate is not possible. Three simple model cases can easily be explored. (1) The whole mantle is depleted. This corresponds to a lower limit of terrestrial content:

1000 [ ] N-type O T-type

•

P-type

x

lAB

100 •

• x 0

•

•

x x

O

008 i o •

0

2000

4000

CBE ~

6000

8000

K ppm Fig. 5. Plot of I vs. K for all basalt samples analyzed in this work. Note that all N-type basalts cluster at both low K and low I contents. The absence of a K - I correlation is thought to be due to distinct carriers for ! and K in the recycling process.

oceanic basalts. The heterogeneity of the source itself can only be explained by the incorporation of an exceedingly iodine-rich component. The only serious candidate is sediment rich in organic matter. 6. Bulk earth iodine estimate

Assuming that (i) the bulk mantle is depleted and (ii) iodine is significantly incompatible, the mean iodine abundance of 8 ppb in N-type basalts (Table 1) and a mean melting fraction of about 10% [4] suggests a lower limit of 0.8 ppb for the mantle. Higher I contents will result if a part of the mantle is still undepleted or enriched. As only a part of P-type basalts are I-enriched and the T-type basalts world-wide do not exhibit any I enrichment (Table 1), this suggests that I-enriched mantle is only a minor fraction of the mantle. For this reason, we believe that the value obtained from low iodine samples is close to the mean mantle value. We can calculate the bulk Earth iodine content as the sum of iodine in (1) the crust, and in (2) the depleted, (3) the enriched and (4) the primitive mantle: CBE =

m c C c + m d C d + m e C e + mpCp

223

mcC c

According to the values reported in Table 2, and with m c = 0.0059, one obtains a bulk Earth (silicate) concentration of 9 ppb, most of which is contributed by the crust. (2) N o primitive mantle exists. The fraction of enriched mantle is estimated from the bulk Earth K concentration of 270 ppm, the continental crust concentration of 9100 p p m [42], the depleted mantle concentration of 88 p p m [42] and an enriched mantle concentration of 350 ppm (the mean of P- and T-type basalts analyzed here). CBE =

m c C c + m d C d -}- m e C e

This yields: m d=

0.51

m e = 0.49 The K / I ratio in the enriched mantle varies between 1.5 x 10 4 and 4.6 x 105 ( K / I ratio for N-types is about 1.1 × 105). The lowest value yields an upper limit on the bulk Earth iodine concentration of 21 ppb. (3) N o enriched mantle exists: primitive mantle and depleted mantle only are considered. If the primitive mantle fraction is estimated to be a maximum of two thirds of the bulk mantle, we

TABLE 2 Estimation of iodine content in continental crust. Data for iodine in crustal components from [15] and unpublished data Component

Mass (1024 g)

Iodine (ppm)

Contrib. to Cont. Crust (ppm)

Seawater Mar. Sed. Non-mar. Sed. Igneous rocks

1.4 0.64 1.76 20.2

0.050 55 0.20 0.063

0.003 1.467 0.015 0.053

Cont. crust

24.0

1.55

224

B. DI~RUELI.E ET AL.

obtain from mass balance considerations a concentration of 24 ppb. These considerations give a bulk Earth iodine concentration of 9 < I ppb < 24. As any combination of the above models will yield a value within this range, no further model needs to be considered in order to bracket possible bulk Earth iodine estimates. Our preferred value can be obtained as follows: If we consider the mean concentration of P- and T-type samples to be reliable, then one obtains the following contributions: 2.1 ppb from the enriched mantle, 0.4 ppb from the depleted mantle and 7.8 ppb from the crust, which yields a bulk Earth concentration of 10.3 ppb. This result is quite close to the estimate of [15], simply because the continental crust contribution dominates the inventory. 7. Further c o n s i d e r a t i o n s

We now arrive at the peculiarity of iodine geochemistry. Iodine is exceedingly concentrated in the crust, and more specifically in the sedimentary layers rich in organic matter. The comparison with a typical incompatible element is very instructive: from crustal K (9100 ppm) [42] and I (1.55 ppm, Table 2) abundances, one obtains a crustal K / I ratio of 6000, while K-depleted mantle abundances suggest that the difference in I concentration between the surface reservoirs and the mantle is too large ( K / I ......tLe/K/I .... t = 1 8 ) to be explained by the conventional models of terrestrial differentiation. This point will be explored next. 7.1. Cosmochemical considerations We obtain in the above estimate a bulk Earth (BE) iodine concentration of ppb. Iodine is known to be a volatile element. Its degree of volatility V can be estimated after normalization to Si and C1 chondrites [15]: V(I) = ( I / S i ) u E / ( I / S i ) c ,

= 9+14 x 10 3

This can be compared with other volatile elements like K or CI: V(K) = 0.25 V(C1) = 0.017 This comparison exemplifies the strongly volatile character of iodine upon planetary accretion. A

very important point is the similarity (within error) of iodine and chlorine volatility, which means that the terrestrial I / C I ratio is similar to the chondritic ratio despite the strongly fractionated character of both C1 and I. One simple explanation is provided by the late v e n e e r / h e t e r o g e neous accretion model for the formation of the Earth [15]. In the first stage of this model, only refractory elements were quantitatively condensed and could be accreted. In the late stage, another component, rich in volatile elements, was added. In this type of model less than 1% of this component is required to provide all terrestrial volatiles with C1 chondrite proportions. 7.2. Geodynamical considerations The overall behaviour of iodine can be viewed in a continental mode[ of terrestrial geodynamics, involving three major processes: continental crust formation, oceanic crust formation, and recycling of crustal material into the mantle. Further complications are not necessary for the present discussion. Mantle depletions and enrichments are the result of these processes; N-type basalts are presently extracted from the depleted mantle which itself results from a previous stage of extraction corresponding to continental crust formation. The continental crust is quite stable and only a small fraction can be recycled into the mantle, most likely as sediments. Oceanic crust, with the possible addition of sediments, is almost completely recycled into the mantle to form the enriched mantle. From this iodine study, we can retain five major points that will constrain geodynamical models: (1) Iodine is a strongly incompatible element, as revealed by its exceedingly high concentration in the crust. The ratio of continental crust to bulk Earth iodine concentration exceeds 65, a very high value when compared to Th and Ba ratios which are of the order of 43 and 41 respectively [42]. The same point is illustrated when considering the ratio of present-day depleted mantle concentration ( A ) to the primitive concentration (A0). The ratio A / A o for iodine is 0.09-0.03, compared to 0.23 for Th and Ba. Only strongly volatile elements, like the noble gases, display lower A / A o ratios. (2) Iodine is not recycled efficiently with oceanic crust. Among all the studied P-type basalt sam-

I O D I N E IN O C E A N I C B A S A L T S A N D I M P L I C A T I O N S F O R E A R T H D Y N A M I C S

pies having the signature of crustal recycling, only a small number display iodine enrichment and most do not exhibit any enrichment compared to N-type basalts. No correlation appears with typical incompatible elements (REE, K, U, Th, Ba, etc.). In particular, the procedure of using the ratio of iodine to other incompatible elements to find its position in the incompatibility sequence [42] fails. This is not the result of analytical problems as the set of N-type basalts yields a variability in iodine concentration of about 40%, comparable to that in L R E E and smaller than that in Th and Ba (about 70%) [42]. The variable concentrations of Th and Ba in N-type basalts result from variable degrees of depletion. The lesser variability for L R E E (or iodine) is understood as resulting from a lower efficiency of extraction. The variability of Th and Ba concentrations in N-basalts is well explained by a model of extensive (and variable) degrees of extraction. The limited variability of iodine concentration in N-type basalts and its exceedingly high concentration in the crust cannot be reconciled by a simple model of crustal extraction from the mantle. (3) Iodine enrichment in basahs results from

sedimentary recycling (especially organic matter). This is observed only for some enriched (P-type) basalts, while some basalts with the signature of sedimentary contribution (high Pb isotope ratios, high Li concentration, etc.) do not exhibit any iodine enrichment. The strong enrichment in only some samples can only result from incorporation of a component even richer in iodine. The only potential component known so far is sedimentary organic matter. This suggests that recycling of sediments into the mantle may or may not induce iodine enrichment and organic C recycling. (4) The CI / I ratio is chondritic. From the bulk Earth chlorine estimate, we derive a C1/I ratio of about 9ano + 2131111 which is not distinctly different 00 ' from the C1 chondrite C I / I ratio, especially when its variability is considered. Both CI and I are strongly depleted on Earth (a factor of about 100 with respect to Si and C1 chondrite). One simple way of achieving these two features is by quantitative simultaneous addition of CI and I with a component carrying volatiles in C1 chondrite proportions, the heterogeneous veneer. (5) The specifity of iodine. Bulk Earth estimates for refractory elements (e.g. Ba, Th with A / A l l = ~

u

225

0.23) are known from abundances in chondrites. Mass balance considerations are in agreement with these estimates, with small differences resulting from both analytical uncertainties and Mighty different behaviour during recycling. The iodine depletion in the mantle ( A / A o ~ 0.9) must be explained by some peculiar behaviour. How is it possible that iodine alone is so strongly depleted compared to the most incompatible elements? This cannot result from different recycling of I compared to Ba and Th. Two possibilities may be suggested: (i) Incompatible elements (e.g. Ba and Th) need a continental crust to be retained in an external reservoir. During the first billion years of Earth genesis, such a reservoir would have been of negligible extent, so that the overall yield of extraction was nearly zero (instantaneous recycling). Iodine stabilization, in contrast, requires the occurrence of sediments rich in organic matter. It might be possible that the early crust provided these sediments at a very early stage. (ii) Iodine is a volatile element upon accretion. Because of late incorporation, iodine was never mixed with mantle material and its present-day abundance in sediments is not the result of extraction from the mantle. In this case, decoupling of iodine from Th, Ba, etc. is obvious. 8. Conclusion

In order to explain all these features of iodine geochemistry we propose the following scenario: (1) Iodine was incorporated into the Earth during the late stage of accretion like other volatile elements. The estimated turnover time of the Earth's mantle during accretion is less than 106 a [15], a short period relative to the accretion time of about 4.107 a [53]. Therefore, the late, volatile-rich accreted veneer material may have been efficiently mixed into a vigorously convecting mantle. (2) Iodine was extracted from the mantle during the process of continental crust formation and probably also during the early stages of the equivalent of present-day oceanic crust formation. Unlike other incompatible elements [43], but like rare gases, iodine was efficiently extracted because it is not recycled with the oceanic crust (or its former equivalent). More likely, iodine is

226

leached out of the crust during hydrothermal processes and concealed in organic molecules of surface sediments. (3) After the major stage of continental crust extraction (or, more specifically, iodine extraction) the residual mantle had been sufficiently stirred and h o m o g e n i z e d to explain why iodine concentrations in most oceanic basalts (either Nor P-types) vary in a restricted range. (4) After the continental formation rate drastically decreased, oceanic crust formation and recycling (with sediments) produced the observed major mantle heterogeneity. The high concentrations in some enriched basalts are understood to be the result of organic sediment incorporation in the source. Because of the high | abundance in this sediment, only a marginal amount is necessary, and the effect this has for all other trace elements will be negligible. This high iodine concentration of some of the enriched basalts is not the signature of an iodine-rich primitive source. If this were the case all P-type basalts would be iodine rich. The high iodine concentration of some enriched basalts may be indicative of organic matter recycling into the mantle. Acknowledgements

This work was supported by the Abteilung Kosmochemie of the MPI at Mainz and by URA 736, CNRS, Paris. Analytical work at Mainz was possible thanks to a grant from the MPG (BD) and the help of Direction des Relations Internationales of the UPMC. We gratefully acknowledge F. Albar~de, H. Bougault, D. Clague, A. Hofmann, T. Juteau, D.D. Karig, V. Lorenz, J.G. Moore, U. von Stackelberg and P. Watremez for providing rock samples. Reviews by Thomas Torgersen, Steve Galer and two anonymous reviewers were very helpful in improving the manuscript. References 1 S.R. Hart, Heterogeneous mantle domains: signatures, genesis and mixing chronologies, Earth Planet. Sci. Lett. 9(1, 273-296, 1988. 2 T. Staudacher and C.J. Allbgre, Terrestrial xenology, Earth Planet. Sci. Lea. 60, 389-406, 1982.

•. I),~RUELLE ET AL. 3 B. Marty and A. Jambon, C / 3 H e in volatile fluxes from the solid Earth: implications for carbon geodynamics, Earth Planet. Sci. Lea. 83, 16-26, 1987. 4 A. Jambon and J.-L. Z i m m e r m a n n , Water in oceanic basalts: evidence for dehydration of recycled crust, Earth Planet. Sci. Lea. 101,323-331, 1990. 5 J.D. Morris, W.P. Leeman and F. Tera, The subducted component in island arc lavas: constraints from Be isotopes and B-Be systematics, Nature 344, 31 36, 1990. 6 H.A. Kennedy and It. Elderfield, Iodine diagenesis in pelagic deep-sea sediments, Geochim. Cosmochim. Acta 51, 2489 2504, 1987. 7 H.A. Kennedy and H. Elderfield, Iodine diagenesis in non-pelagic deep-sea sediments, Geochim. Cosmochim. Acta 51, 2505-2514, 1987. 8 W.J. Ullman and R.C. Aller, The geochemistry of iodine in near-shore carbonate sediments, Geochim. C o s m o c h i m Acta 49, 967-978, 1985. 9 M. Yoshida, K. Takahashi, N. Yonehara, T. Ozawa and I. lwasaki, The fluorine, chlorine, bromine, and iodine contents of volcanic rocks in Japan, Bull. Chem. Soc. Jpn. 44, 1844-1850, 1971. 10 A. Brehler and R. Fuge, Iodine, in: Handbook of Geochemistry, K.H. Wedepohl, ed., Springer, Berlin, 1974. 11 G. Dreibus, E. Jagoutz and H. W~inke, Distribution of CI, Br, and iodine in mantle and crust, Terra Cognita 3, 122, 1983. 12 K.K. Turekian and K.H. Wedepohl, Distribution of the elements in some major units of the Earth's crust, Geol. Soc. Am. Bull. 72, 175 192, 1961. 13 K.H. Wedepohl, The contribution of chemical data to assumptions about the origin of magmas from the mantle, Fortschr. Mineral. 52, 141-172, 1975. 14 Chilean Iodine Educational Bureau, Geochemistry of Iodine, 150 pp., Shenval Press, London. 1956. 15 11. W~inke, G. Dreibus and E. Jagoutz, Mantle chemistry and accretion history of the Earth, in: Archean Geochemistry, A. Kr6ner et al., eds., pp. 1-23, Springer, Berlin, 1984. 16 C.J. Allbgre, Chemical geodynamics, Tectonophysics 81, 1(19 132, 1982. 17 H. Bougault, J.-L. Joron, M. Treuil and R. Maury, Local versus regional mantle heterogeneities: evidence from hygromagmaphile elements, Init. Rep. DSDP 82, 459-481, 1985. 18 H. Bougault and M. Treuil. Mid-Atlantic Ridge: zero-age geochemical variations between Azores and 22°N, Nature 286, 209-212, 1980. 19 R. Hekinian, M. Chaigneau and J.-L. Chemin~e, Popping rocks and lava tubes from the Mid-Atlantic Rift Valley at 36°N, Nature 245, 371-373, 1973. 20 J.G. Moore, D.A. Clague and W.R. Normark, Diverse basalt types from Loihi seamount, Hawaii, Geology 10, 88 92, 1982. 21 J.G. Moore, Petrology of deep-sea basalt near Hawaii, Am. J. Sci. 263, 40-52, 1965. 22 T.K. Kyser and J.R. O'Neil, Hydrogen isotope systematics of submarine basalts, Geochim. Cosmochim. Acta 48. 2123-2133, 1984. 23 R. Hekinian and D. Walker, Diversity and spatial zonation

IODINE

24

25

26

27

28 29

30

31

32

33 34

35

36

37

38

IN O C E A N I C

BASALTS AND IMPLICATIONS

FOR EARTH

of volcanic rocks from the East Pacific Rise near 21°N, Contrib. Mineral. Petrol. 96, 265-280, 1987. A. Prinzhofer, E. Lewin and C.J. All?~gre, Stochastic melting of the marble cake mantle: evidence from local study of the East Pacific Rise at 12°50'N, Earth Planet. Sci. Lett. 92, 189-206, 1989. R. Hekinian, M. Fevrier, F. Avedik, P. Cambon, J.-L. Charlou, H.D. Needham, J. Raillard, J. Boul~gue, L. Merlivat, A. Moinet, S. Manganini and J. Lange, East Pacific Rise near 13°N: geology of new hydrothermal fields, Science 219, 1321-1324, 1983. R. Hekinian, J. Francheteau, V. Renard, R.D. Ballard, P. Choukroune, J.L. Chemin~e, F. Albar~de, J.F. Minster, J.L. Charlou, J.C. Marry and J. Boul~gue, Intense hydrothermal activity at the axis of the East Pacific Rise near 13°N: submersible witnesses the growth of sulfide chimney, Mar. Geophys. Res. 6, 1-14, 1983. C.W. Devey, F. Albar~de, J.-L. Chemin6e, A. Michard, R. Miihe and P. Stoffers, Active submarine volcanism on the Society hotspot swell (West Pacific): a geochemical study, J. Geophys. Res. 95, 5049-5066, 1990. D.D. Karig, Pets. Commun., 1989. G. Loock, W.F. McDonough, S.L. Goldstein and A.W. Hofmann, Isotopic compositions of basaltic glasses from the Lau Basin, Mar. Min. 9, 235 245, 1990. J.-M. Auzende and 28 Co-authors, L'accr6tion r6cente dans le Bassin Nord-Fidjien: premiers r~sultats de la campagne franco-japonaise Kaiyo 87, C.R. Acad. Sci. Paris S~r.II 306, 971-978, 1988. D. Dupuis, P~trologie des basaltes de la fosse Mabahiss. Place de cette fosse dans le contexte de la Mer Rouge, M6m. DEA, Univ. Brest, 1986 (Unpubl.). W.M. White and J. Patchett, Hf-Nd-Sr isotopes and incompatible element abundances in island arcs: implications for magma origins and crust-mantle evolution, Earth Planet. Sci. Lett. 67, 167-185, 1984. J.-L. Zimmermann, Pers. Commun., 1990. C.P. Thornton and O.F. Tuttle, Chemistry of igneous rocks--I. Differentiation index, Am. J. Sci. 258, 664-684, 1960. A.P. Le Roex, Source regions of mid-ocean ridge basalts: evidence for enrichment processes, in: Mantle Metasomatism, M.A. Menzies and C.J. Hawkesworth, eds., pp. 389422, Academic Press, London, 1987. C.K. Unni and J.G. Schilling, CI and Br degassing by volcanism along the Reykjanes Ridge and Iceland, Nature 272, 19 23, 1978. N. Metrich, Chlorine and fluorine in tholeiitic and alkaline lavas of Etna (Sicily), J. Volcanol. Geothermal Res. 40, 133-148, 1990. G. Dreibus, B. Spettel and H. W~inke, Halogens in meteorites and their primordial abundances, in: Origin and Distribution of the Elements, L.H. Ahrens, ed., pp. 33-38, Pergamon, Oxford, 1979.

DYNAMICS

227

39 K.G. Heumann and H. Weiss, Mass spectrometry iodine trace determinations in geochemical standard reference materials and in meteorites, Fresenius Z. Anal. Chem. 323, 852 858, 1986. 40 G. Dreibus, H. W~inke and L. Schultz, Mysterious iodineoverabundance in Antarctic meteorites, in: Int. Workshop on Antarctic Meteorites, Houston Lunar Planet. Inst. Tech. Rep. 86-01, pp. 34-36, 1986. 41 K.G. Heumann, M. Gall and H. Weiss, Geochemical investigations to explain iodine-overabundances in Antarctic meteorites, Geochim. Cosmochim. Acta 51, 2541-2547, 1987. 42 A.W. Hofmann, Chemical differentiation of the Earth: the relationship between mantle, continental crust, and oceanic crust, Earth Planet. Sci. Lett. 90, 297 314, 1988. 43 E. Ito, D.M. Harris and A.T. Anderson, Jr., Alteration of oceanic crust and geologic cycling of chlorine and water, Geochim. Cosmochim. Acta, 47 1613-1624, 1983. 44 J.-G. Schilling, M.B. Bergeron and R. Evans, Halogens in the mantle beneath the North Atlantic, Philos. Trans. R. Soc. London A297, 147-178, 1980. 45 C.W. Burnham, Water and magmas: a mixing model, Geochim. Cosmochim. Acta 39, 1077-1084, 1975. 46 H. Sakai, D.J. Des Marais, A. Ueda and J.G. Moore, Concentrations and isotope ratios of carbon, nitrogen and sulfur in ocean-floor basalts, Geochim. Cosmochim. Acta 48, 2433 2441~ 1984. 47 G.W. Reed and S. Jovanovic, The halogens and other trace elements in Apollo 12 samples and the implications of halides, platinum metals, and mercury on surfaces, in: Lunar Science Conf. Proc., 2nd, pp. 126l, 1971. 48 G. Dreibus, B. Spettel and H. W~inke, Determination of lithium and halogens and the significance of lithium to the understanding of cosmochemical processes. J. Radioanal. Chem. 38, 391 403, 1977. 49 Ph. Vidal and N. Clauer, Pb and Sr isotopic systematics of some basalts and sulfides from the East Pacific Rise at 21°N (project RITA), Earth Planet. Sci. Lett. 55, 237-246, 1981. 50 B. Hamelin, B. Dupr6 and C.J. AIIbgre, Lead-strontium isotopic variations along the East Pacific Rise and the Mid-Atlantic Ridge: a comparative study, Earth Planet. Sci. Lett. 67, 340-350, 1984. 51 H. Staudigel, A. Zindler, S.R. Hart, T. Leslie, C.-Y. Chen and D. Clague, The isotope systematics of a juvenile intraplate volcano: Pb, Nd, and Sr isotope ratios of basalts from Loihi Seamount, Hawaii, Earth Planet. Sci. Lett. 69, 13-29, 1984. 52 W.M. White, A.W. Hofmann and H. Puchlet, Isotope geochemistry of Pacific mid-ocean ridge basalt, J. Geophys. Res. 92, 4881-4893, 1987. 53 G.W. Wetherill, Accumulation of the terrestrial planets, in: Protostars and planets, T. Gehrels, ed., University of Arizona Press, Tucson, 1978.