Noble gases in mafic phenocrysts and xenoliths from New Zealand

Geochimica et Cosmochimiea Acta, Vol. 58, No. 20, pp. 441 i-4427, 1994 Copyright Q 1994 Elsevier Science Ltd Printed in the USA. All kits rewved 0016-7037/94-$6.00 + .OO

Pergamon

0016-7037(94)00217-7

Noble gases in mafic phenocrysts and xenoliths from New Zealand D. B. PATTERSON,* M. HONDA, and 1. MCDOUGALL Sciences, The Australian National University, Canberra, ACT 0200, Australia

ResearchSchool of Earth

(Received July 1, 1993; accepted in revised form April 13, 1994)

Abstract-We have determined the elemental and isotopic compositions of noble gases in young subduction-related phenocrystic olivine and clinopyroxene samples from the Taupo Volcanic Zone, central North Island, New Zealand, and in behind-arc intraplate phenocrystic and xenolithic olivine samples from the Northland and Auckland Volcanic Provinces, northern North Island, New Zealand. Helium isotopic ratios range from MORB-like 3HefPHe values of about 11 X 10m6to lower values of about 6 X 10e6, consistent with previous measurements of helium isotopic ratios in subduction-retated samples. 40Ar/36Ar and 2’Ne/22Ne ratios range from atmosphere-like compositions to maximum values of about 700 and 0.033, respectively. In contrast, most of the *‘Ne/**Ne ratios are generally indistinguishable from atmospheric values. The variations in helium, neon, and argon isotopic ratios are interpreted as resulting from mixing of: (a) a primordial 3He-rich component derived from the upper mantle and characterised by MORB-like 3He/4He ratios of about 12 X 10e6; (b) a radiogenic 4He-rich component derived from crustal materials and character&d by a 3He/4He ratio of less than 2 X IO-?, and (c) a helium-poor atmosphere~e~ved com~nent which dominates the heavier noble gases. By combining the helium and argon isotopic results, it is possible to estimate the relative contribution of each of these three components to the total *Ar observed in each sample. Based on the present understanding of the origin and evolution of arc magmas, a simple qualitative model of the mechanisms which introduce noble gases to the parent magmas of the samples is developed. However, it remains uncertain whether the atmospheric and crustalderived noble gas components are introduced to the mantle source regions of the magmas by subduction, or alternatively, are introduced by interactions between the ascending parent magmas and the overlying crust. MARTY et al., 1989; NAGAO

ANALYSIS OF THE ELEMENTALand isotopic composition of the noble gases (He,

Ne, Ar, Kr, and Xe) trapped in mantlederived samples erupted at mid-oceanic ridges ( MORBs) and at intraplate hotspot plumes (e.g., Hawaii, Reunion, Samoa), has provided important constraints on the geochemical structure and evolution of the terrestrial mantle-crust-atmosphere system. This has led to the real&ion that the upper mantle, as sampled by N-type mid-oceanic ridge volcanism, is remarkably uniform in helium isotopes with a 3He/4He ratio of about 11.8 (+0.4) X 10m6(e.g., KURZ, 199 1). The isotopic compositions of Ne, Ar, and Xe in MORB samples are much more variable, and appear to require mixing of atmosphere-like and nonatmospheric heavier noble gas components (e.g., ALLEGRE et al., 1983, 1986; FARLEY and POREDA, 1993; HIYAGON et al., 1992; HONDA et al., 1993a; MARTY, 1989; SARDA et al., 1988; STAUDACHERet al., 1989). In contrast, the study of noble gases in su~u~ion-related volcanic arcs has been almost entirely restricted to the analysis of 3He/4He isotopic ratios of arc-related fluid samples such as water, steam, and gas collected from volcanic fumaroles, hot springs, and geothermal wells, with results from over four hundred and fifty samples reported in the literature (e.g., BASKOVet al., 1973; CRAIG et al., 1978; GIGGENBACH et al. 1993; HILTON and CRAIG, 1989; HILTON et al., 1993a;

* E+esenfaddress: Division of Geological and Planetary Sciences, California Institute of Technology, MSli’O-25,Pasadena, CA 9 1125, USA. 4411

et aI., 1979, 198 1; POLAKet al., 1982; POREDAand CRAIG, 1989; SANO and WAKITA, 1985, 1989; SANO et al., 1982, 1984, 1987, 1988; TEDESCOet al., 1990; TORCERSENet al., 1982). On the other hand, the helium isotopic composition of phenocrysts and xenoliths extracted from arc-related rocks has been reported for only about forty samples (HILTON and CRAIG, 1989; HILTON et al., 1992, 1993a; MARTY et al., 1989; POREDAand CRAIG, 1989; TO~TIKHIN et al., 1972, 1974). These studies have shown that the 3He/4He ratios of su~u~ion-~lat~ samples range from MORB-like values of about 12 X 10e6 down to values lower than the atmospheric value of 1.4 X 10m6.This variation has heen interpreted as the result of mixing of two dominant helium components: (a) a mantle-derived component rich in primordial 3He and characterised by a MORBlike 3He/4He ratio ofabout 12 X 10e6, and(b) a radiogenic 4He-rich component character&d by a crustal-like ‘He14He ratio of less than 2 X lo-’ (e.g., GIGGENBACHet al., 1993; HILTONet al., 1992,1993a; POREDAand CRAIG, 1989; SANO and WAKITA, 1985; SAKAMOTOet al., 1992; SANO et al., 1987; TORGERSENet al., 1982). Helium has been the focus of noble gas studies of arc-related samples in part because contamination by atmospheric He generally is insignificant, owing to very low abundance of He in the atmosphere. The hiavier noble gases (Ne, Ar, Kr, and Xe) in arc-related fluids appear to be almost entirely atmospheric in origin, dominating any magmatic heavier noble gases which may be pres ent (e.g., HULSTONet al., 1986; NAGAO et al., 1979, 1980a, 1981; STAUDACHERand ALL~GRE, 1988; TORGERSENet al., 1982).

4412

D. B. Patterson, M. Honda, and 1. McDougall

In an attempt to circumvent the problem of addition of atmosphere-derived heavier noble gases to fluid samples, and to extend the noble gas dataset for arc-related samples, we have measured the elemental and isotopic composition of all five noble gases trapped in olivine and clinopyroxene phenocrysts and xenoliths extracted from subduction-related and behind-arc intraplate basalts and basaltic andesites from New Zealand. It was hoped that such mafic phenocryst and xenolith phases might have trapped a component of ambient magmatic noble gases before their host magmas equilibrated with the atmosphere. This study confirms the presence of MORB-like He and identifies nonatmospheric components of Ar and Ne in arc magmas. The variations in the observed helium, neon, and argon isotopic ratios can be accounted for by mixing of three noble gas components: a mantle-derived component characterised by a MORB-like 3He/4He ratio, a crustal-derived component rich in radiogenic 4He characterised by low 3He/4He ratios, and a helium-poor atmospherederived component which dominates the heavier noble gases in the samples. On the basis of these observations and the present understanding of the origin and evolution of arcmagmas, we present a generalised qualitative model for the movement of noble gases through subduction systems.

Australian

Plate

BACKGROUND GEOLOGY AND SAMPLE SELECI-ION Samples were collected from two distinct geochemical and tectonic settings within North Island, New Zealand: the subduction-related Taupo Volcanic Zone of the central North Island, and the intraplate Northland and Auckland Volcanic Provinces of the northern North Island (Fig. I ). Details of sample number, location, and lithology of host and parent lavas are given in Appendix I. Subduction-related

Samples: Taupo Volcanic Zone

The Taupo Volcanic Zone (TVZ) extends some 300 km across the central North Island of New Zealand from Ohakune in the south to White Island in the north. The TVZ comprises an active volcanic arc and marginal basin which lies 200-270 km west of the structural trench (the Hikurangi Trough) and about 80 km above the Benioff Zone (COLE and LEWIS, 1981; COLE, 1986). The TVZ is a geologically young feature, with the onset of volcanism related to changes in the configuration of the boundary between the Indo-Australian and Pacific plates approximately 0.75 Ma ago (BALLANCE, 1976; COLE, 1986; WALLCOTT, 1984). Volcanic activity within the TVZ appears to have been focused within five geographically distinct volcanic centres: Tongariro, Taupo, Maroa, Rotorua, and Okataina. Over 95% of the andesitic lavas of the TVZ are found in the southernmost centre of Tongariro, whereas the four more northern centres are almost exclusively rhyolitic with small volumes of high alumina basalts. Numerous previous workers have documented the geology and geochemistry of the rocks of the TVZ (see for example: COLE, 1981; COLE et al., 1986; GAMBLE et al., 1990). Subduction-related basaltic to andesitic samples which contained sufficient fresh olivine and/or clinopyroxene for extraction were collected from five locations within the TVZ. Three samples yielded olivine, one sample yielded clinopy-

FIG. I. Location of young volcanic regions in the North Island, New Zealand. Subduction-related volcanics of the Taupo Volcanic Zone (TVZ), Alexandra Volcanic Centre (AVC), and Egmont Volcanic Centre (EVC) in stippled pattern. lntraplate basalt volcanics of the Northland Volcanic Province (NVP) and Auckland Volcanic Province ( AVP) in solid pattern. Labeled arrows show the motion of the Pacific Plate relative to the Australian Plate in mm a-’ (plate motions from WALLCOTT,1978). Also shown is the approximate location of the Hikurangi Trough where active subduction of the oceanic Pacific Plate beneath continental New Zealand occurs.

roxene, and one sample yielded both olivine and clinopyroxene. Two of the olivine samples were analysed in triplicate. Intraplate Samples: Northland and Auckland Volcanic Provinces The Northland-Auckland region comprises the northern 400 km of North Island, New Zealand. Within this region there are a large number of young (approximately 2 Ma to Recent) predominantly basaltic monogenetic volcanoes which do not appear to be associated with subduction-related processes. These young intraplate volcanoes are not considered to be the result of a mantle plume or hotspot because of the lack of any systematic age-location correlation (HEMING, 1980a). The intraplate volcanoes can be subdivided into the two geographically distinct volcanic provinces of Northland and Auckland (Fig. 1). Both provinces occur in tectonically quiet, gently extensional settings and are regarded as intraplate, well removed from the present-day convergent margin which lies between 350 and 700 km to the southeast. The geology and

Isotope ratios of noble gases in phenocrysts

geochemistry of these intmplate basalts have been discussed by various workers including ASHCRO~~ ( 1986), HEMING (1980b), and HEMINGand BARNET(1986). Some of the basalts of the Auckland Volcanic Province also contain mantle-derived olivine and clinopyroxene rich nodules. Individual nodules range from l-5 cm in size and are typically contained within aphyric undersaturated nepheline-no~ative alkali basalts. The nodules themselves are regatded as xenolithic fragments of upper mantle material and are not crystal cumulates, However, the question of whether they represent refractory residues of parental partial melts, or accidentalIy incorporated fragments of mantle material is less clear (see for example: RAFFERTY and HEMING, 1979; ROWERS et al., 1975). A totai of eight samples of young intrapiate basalts having sufficient divine for extraction were collected from the No~hIand-Auckland region. Five of these samples yielded phenocrystic olivine, and three samples yielded xenolithic olivine from dunitic mantle-derived nodules. Clinopyroxene was not extracted from any of the Northland-Auckland intraplate basal&. SAMPLE PREPARATION Samples were prepared by crushing l-2 kg ofwhole rock in a steel jaw crusher. The crushed rock was then placed in a sieve stack and separated into a range of grain sizes. The size fraction which yielded the maximum amount of clean phenocrysts was selected for further processing. Olivine and clinopyroxene concentrates were obtained by use of heavy liquid gravimetric separation with methylene iodide (CH&, s.g. 3.32 g cm-‘). Extreme care was taken to remove any residual methylene iodide by thorough ultrasonic washing in acetone and ethanol. A second stage magnetic separation was used to remove magnetic impurities and to partially separate olivine from clinopyroxene. Final mineral separation was performed by handpicking under a binocular microscope to remove remaining impurities and grains which showed signs of alteration. The final separates obtained comprised high purity olivine or clinopyroxene with a minimum grain size of 0.42 mm (Appendix 2). Under a binocular microscope the olivine separates comprised clean translucent grains with relatively rare fluid or melt inclusions. The clinopyroxene separates comprised clean grains which, along the thin semi-transparent edges of broken crystals also appeared to contain relatively few inclusions. A split of each mineral separate was pfaced in a polished epoxy resin mount and between 20 and 30 individual crystals were probed for chemical composition using energy dispersive analysis with a Cameca* electron microprobe. The forsterite content of olivine cores were typically between 80 and 90 mol’lbMg/Mg + Fe, although values ranged from as low as 74.6 to as high as 94.9 (Appendix 2). Normal zonation was common with a thin outer rim having forsterite contents between S-IO% lower than the cores. The clinopyroxene separates are augitic with Mg-Fe ratios of between 70.9 and 83.4 mot% Mg/

Mg+Fe (Appendix 2).

Details of the analytical techniques employed to measure noble gas compositions have been described previously in PATTERSON ( 1992) and HONDAet al. ( 1993a), and only a brief outline will be presented here. Between l-2 g of sample (Appendix 2) was split into four to six equal aliquots which were individually wrapped in tin foil balls ap proximately 4 mm in diameter. Up to three samples were then loaded into a sample holder attached to the top of the extraction furnace. To minim& atrn~p~e~c ~on~mination, the samples were then baked at about 1SOY, whereas the remaining ail metal extraction and purification system was baked at 250°C for at least 36 hours. After baking, the tantalum crucible and molybdenum liner of the

4413

double-vacuum resistively heated extraction furnace were outgassed at between 1700 and 1900°C for one to three hours. The samples were then sequentially dropped into the furnace and noble gases extracted by either total fusion (up to 1900°C for olivine and 1600°C for ciinopyroxene), or by dissolving the sample in a previously outgassed basaltic glass flux at 1500°C (Appendix 2). The glass fluxing technique was necessary because the high temperatures required to fuse divine ( 1900°C) resulted in the frequent mechanical failure of the resistively heated tantalum element. After extraction, the noble gases were purified by exposure io a series of three SAES@getters and a titanium filament flash getter. The gases were then collected on a cryogenic cold trap containing activated charcoal cooled to between 18 and 20 K by a commercially manufactured double-stage helium expansion refrigerator. At these tem~ratures (~20 K) all noble gases, including 3He, are physically adsorbed onto the surface of the activated charcoal within the trap, By means of a heater coil the temperature was incrementally increased and each of the five noble gases sequentially desorbed at successively higher temperature for analysis. Elemental and isotopic analysis was performed using an online VG.5400”noble gas mass spectrometer.

Full procedural hot blanks were measured before and after each sample. The blank amounts varied significantly over the period of this study: (in units ofcm3 STP) 4He from 6.4 X IO-” to 2.7 X 10m9, 20Ne from 4 2 X iO-12 to 8.7 X IO-“, @Ar from 5.7 X iOM9to 5.7 X lOas, 84K; from 2.1 X IO-” to 2.2 X 10-12, and ‘“Xe from 1.3 x 10mi4to 8.7 X IO-“. Gas amounts in the before-sample and aftersample blanks typically varied by between 10and 20% and sometimes by as much as 50%. The blank corrections were based on the average of the before-sample and after-sample blanks. and the un~e~inty was taken to cover the range in the before-sample and after-sample blank gas amounts. This variability in blank amount is the primary cause of the relatively large uncertainties in the isotopic compositions of some samples, particularly for samples with low gas abundances. The ‘He signal was below background for all the blank measurements, and the blank 3He/4He ratio was assumed to be atmospheric. The isotopic composition of the blanks for the four heavier noble gases (Ne, Ar, Kr, and Xe) were indistinguishable from atmospheric. Corrections for interference to %e from 4oAr’+ and H2180”, and to 2ZNe from CO;’ were determined by monito~ng the 4oAr+, H2160+, and CO; peaks. The size of the 4oAr++, H2180+, and CO;’ contributions to “%e and 22Ne were then calculated from the previously determined “Ar++/@Ar+, Hz’80+/H2’60+, and CO:‘/CO~ ratios in the mass spectrometer. The sensitivities and mass discrimination corrections for the noble gases were calibrated at regular intervals by the analysis of aliquots of known volume and isotopic composition delivered from a series of gas pipettes. The uncertainties arising from mass discrimination and neon interference corrections are generally insignificant in comparison to the uncertainties in the blank correction.

NOBLE GAS RESULTS The elemental abundances of all five noble gases, and the isotopic ratios of He, Ne, and Ar for the olivine and clinopyroxene samples from New Zealand are presented in Table 1. The isotopic ratios of Kr and Xe were indistinguishable from atmospheric within uncertainties(P~~~~~~~~, 1992). Listings of the measured krypton and xenon isotopic ratios are available from the authors. Elemental Abundances and Ratios The abundances of the noble gases varies between samples by more than an order of magnitude. The two clinop~xene samples are notable for their low He abundances, whereas two of the intraplate xenolithic olivine samples are notable for their high Kr and Xe abundances. Overall there is a weak positive correlation between the abundan~s of Ne and Kr with Ar, but there is no clear systematic correlation between the abundances of He or Xe with Ar.

[88207]

3He/4He

(10-6)

4We

(lo-*)

IO.?&)

[88244c]

9.02(24) 8.29(23) 6.07(18)

9.32(25)

1.34(6) s.oa(28t

1.43(131 8.93(67) l.%(6) 8.96127)

8.2%36) 6.65(29) 9.10(40)

13.1(6)

---1.4

0.42(31) 0.060(33) 7.0 (34)

22Ne

41.3(21)

16.2(7)

20.2(55) 4.02(63) 3.07(61)

3.28(32) 12.4(11) 7.95(75)

4X0(17)

24.4(41)

(10-12)

0.0284~47)

0.02992(73) 0.03 16(14) 0.0328(17)

0.0329(18) n.d. 0.0315(16)

3.1(25) 9.1(SO) 13.0(58)

1X3(63) n.d. 8.7(56)

2.t5(63)

2.907)

%

621

-1.1 (1.4) -2.1 (16)

2.5(21) 1.8141) 2.6(48)

8.1(X) n.d. 2.4(49)

lSl(45)

2.5(14)

%

620

9.80

0.0290

9.769(50) 0.02884(2.5) -0.35(51) -0.62Ca6) ---*

9.69 (13)

10.05(21) 9.98(40) 10.0X47)

10.&X50) n.d. 10.04(48)

9.951t44j 0.02964{18)

lU.O5(14) o.O29a6(49)

z%Je/2*Ne 2 t Ne(22Ne

348(Z)

302(Z)

JQrP6Ar

0.1833( 12)

0.1849123)

38ArP6Ar

352(53)

187(17j

1~wm 211y54) 128(34)

295.5

31l(3)

3w(3)

0.1880

0.1864(22)

0,1899(31)

380(61 0.1894(11) 526x83) 0.1912(41) 653(130) 0.188?(52)

145122) 588(59) 0.1908(47) 220(16) 490(16) 0.1880(40) 471(568) 1160(1320) 0.184(28)

449(65)

(lo-‘2)

%r

from Taupo Volcanic Zone, Central North Island, New Zealand: Noble gas data 84~

4.54(35)

3.7(11)

n.d. 4.54(27) 4.55(23)

3.92( 19) .5.76(31) 4.0X22)

8.68(75)

5.84(77)

(10-q

116(8)

@GO)

40.8(36) 59.9(49) 58.1(42)

78.1(60) 65.8(47) 55.5(37)

18.2135)

13*X, ( 10-14)

Gas abundances in cm%TP g-1. &XI(%)defiied as l~~~(~oNe~2~e) ~~~l~(zoNe~2Ne)~~ )- 11,and similariy for &I(%). Numbers in brackets give lcr uncertainty of last quoted digits and includes &heuncertainties in the cone&on factors for mass dictation and sensitivity determined by repeat analysis of the standard gases, as well as -ties in the blank correction and interference corrections in the case of Ne, Kr and Xe isotopic ratios are atmospheric within uncertainties. n.d. indicates not determined. Atmospheric ratios from OZIMAand P~DOSEK (1983). Suffixes a, b, and c after GNU number indicates repeat analyses, Extremely large uncertainties in argon results for QC547b are because of a large error in the blank correction.

Atmosphere

[88209]

P.H.47

WZ Pheno&~.~t~&~lin~p~r~xene Oh&me [88206] 0.50(20) 0.077(II) 6.5(24)

12.8(E) 12.2(6)

74.5(38) .5S.l(28j X5.2(29)

12216)

6.62(83) ~.aPa(84~ 7.38(6lj

[882&a] [88244b]

Waco

[9006%] [90067bI

[8820!3]

P.H.47

Mica

ongaroto 1882421

TVZ Phenortystic olivine

3He (10-14)

1a : Subduction-related phenocrystic olivine and c~~yroxene

Location [AN@]

Table

4He (108)

3He

(10-Y (10-6)

3He/4He

2.90)

37.8(19)

0.34(l)

3.66(16) 8.39(59)

10.33(29) 3.64(46)

3.40(37)

2ae (Iv’*)

10.1(5)

40.1(20)

0.91(4)

4.01(17) 11.15(33)

9.99(27)

1.4

5.3(9)

0.0326(4)

0.0290

10.33(9)

9.80

18.9(6)

1.8(11)

7.82(60)

0.0296(15)

9X8(13) 9.98(11)

3.8(32)

-0.6(10)

-048)

0.3(40)

36Ar

0.1902(18)

0.1880

614(4)

295.5

12.4(13)

0.1877(37)

571(6)

0.1859@7)

0.189504)

0.1860(35)

0.1855(28)

0.2013(97)

0.1998(91)

38Arp6Ar

1.9(51)

696G6)

403eO)

41Y3)

307(6)

330(3)

4OArp6Ar

5’6(6)

W8)

223(35)

119t5)

57(B)

58(3)

(10’2)

0.5(21)

-0.7(60)

7.3(15)

-1.7(18)

-5.8(49)

-2.9(57)

%

46

0.2(37)

621

620

0.8(13)

0.0288( 17)

0.03195(30)

0.02852(53)

O.a273( 14)

0.0282( 16)

*lNeme

o.cr291q62)

10.18(31)

9.75( 10)

6X(38)

13.3(12)

11.1(S)

9.76(S)

9.83(40)

9.82(36)

*wee

84Kr

45.3(19)

33.3( 14)

*.06(9)

3.21(13)

4.62(42)

3.60(15)

1.12(S)

1.22(5)

(IO-t*)

376(25)

597W)

37.9(39)

45.5(31)

55.9(77)

64.8(44)

16.8(10)

15.1(10)

lJ*Xe (10-14)

Gas abundances in cm3STP g-t. &0(9b) defined as lOO[((20Ne/r*Ne)~P&(~e~*Ne)~ )- 11. and similarly for $t(%). Numbers in brackets give 10 uncertainty of last quoted digits and includes the -ties in the correction factors for mass discrimination and sensitivity m by repeat analysis of the standard gases, as well as -ties in the blank correction and interference corrections iu the case of Ne. Kr and Xe isotopic ratios are atmospheric within uncertainties. n.d. indicates not determined. Atmospheric ratios from oZIh4A and PODCISEK (1983). Sufhxes ‘a’ and ‘b’ after ANU number indicate duplicate analyses.

Atmosphere

Auckland Volcanic Province, xenolithic olivine L. Pupuke [88276] 33.6(17) 3.42(E) 9.82(28) Ridge Rd. [88263] 14.1(S) 8.10(32) 1.74(S) Stone Rd. 32.2( 14) [88283] 294(ts) 9.14(24)

[88282]

Puketlllll

[88269]

ClatexHiu

Auckkmd Volcanic Province, phenocrystic olivine Wili [88265] 26.ql3) 2.81(12) 9.45(25) 1 lS(4)

[88238]

Pidilly

ww

MnMand Volcanic Province, phenocrystic olivine Rlketala

L4lcation MNUwl

Table lb: Intraplatephenocrystic and xenolithic olivine from Northland and Auckland Volcanic Provinces, northern North Island, New Zealan& Noble gas data

4416

D. B. Patterson,

M. Honda,

The relative elemental ratios 3He/36Ar, 22Ne/36Ar, “4Kr/ 36Ar, and ‘32Xe/36Ar, expressed as F-ratios relative to atmospheric compositions, are listed in Table 2 and are plotted in Fig. 2. The 3He/36Ar (F3) ratio varies by more than two orders of magnitude, from fifty-one to nearly thirty thousand times the atmospheric value. The F3 ratios of the two clinopyroxene samples are notably lower than the olivines, consistent with the low absolute abundance of He in these samples. In contrast, the 22Ne/36Ar ( F2*) and 84Kr/36Ar ( Fs4) ratios show a much smaller variation, and are mostly within a factor of three of the atmospheric composition. Importantly, the F2* and Fs4 ratios are inversely correlated, with low Fs4 ratios associated with high Fz2 ratios and vice versa. The ‘32Xe/3hAr (Fl32) ratios show greater variability, including F13? values of over 10. Because of the very low relative abundance of He in the atmosphere in comparison with the heavier noble gases (Ne, Ar, Kr, and Xe), addition of an atmospheric component would result in large variations in F3 ratios (owing to the variable addition of atmospheric 36Ar), but near-atmospheric F22, Fs4, and F132 ratios. As shown in Fig. 2, this is what is observed in the New Zealand samples. The variations in F22, Fs4, and F132values about the atmospheric composition could be accounted for by solubility controlled elemental fractionation of the noble gases during mass transfer of gases between basaltic melt and a separate vapour phase (HONDA et al., 1993a; PATTERSON et al., 1990, 199 1). This could occur in at least two ways: (a) during interaction of magma with a separate vapour phase carrying atmosphere-derived heavier noble gases, and (b) during subsequent exsolution and loss of dissolved volatiles from the melt during bubble formation. These elemental fractionation effects are shown in Fig. 2b and 2c. Within a factor of two, interaction between basaltic melt and a vapour phase carrying dissolved atmosphere-derived heavier noble gases, followed by bubble formation and

and I. McDougall

. 1(a)

100000

I

.

1

.

F : * 100

:

I

,o_.

3

1

. F22

3

2

3 0 Air

(b)

A

Seawater

n Magma 2

F22 1

0

0

2

1

3

F84

Table 2: Elemental abundance GI,IOS

Location

ANU#

Fi

Taupe Volcamc Zone. Phenocrystu oiiwne 631 (121) OtIp310 88242 8197 (533) Mahum 88207 21950(3470) PH.47 88209 10700 (950) 9006la 427 (59) Wmmanno 88244a 2590 (700) 88241b 3630 (980) 88244c

F22

FIG

hi4

(22) (7) (7) (14) (8) (11) (15)

0.63 ( 12) 0.65 (6) I.31 (21) I.?h(ll) n.d. I .09 (30) 1.71(46)

0.55 2.08 7.25 4.02 0.43 4.0 6.1

Taupe Volconrr Zone. Phenocrystrc clrwpyroxene 115 (47) 1.63 (17) OhZlklJlle 88206 51 (38) 2.20 (36) P.H.47 88209

0.94 (29) 0.62 (IO)

4.3 (22) 4.46 (74)

o/nine I.10 (II) 1.19 (60)

1.02 (6) 0.95 (47)

3.52 (29) 4.0 (20)

olivine I.81 (9) 0.94 (16) 1.84 (16)

1.46 (9) l.OO(lS) I.14 (8)

7.32 3.37 4.5

(58) (71) (4)

0.78 (5) 2.73(16) 2.28(13)

4.02 13.6 5.26

(451 (11) (41)

Provrncr Phenocrystic

Northland Volcam Puketona 88228 Picadilly 88238

27900( 1900) 2 I SO( 1070)

Auckland Volcanic Province. Phenocrystic 9550 (630) Wlli 88265 7700 (130) Crater Hill 88269 3180 (260) Puketutu 88282

1.02 1.26 0.42 1.06 0.30 0.37 0.45

Auckland Volcanic Provme’ Xenoiirhw olwine I 1330 (770) 0.96 Lake Pupuke 88276 1025 (75) 0.25 Ridge Road 88263 13080 (850) 0.37 Stone Road 88283

(7) (I) (2)

(13) (9) (12) (41) (5) (II) (17)

where IX 1s one of ZHe. 22Ne, “4Kr F ratios defined as F, = (1X~6Ar),,,gt~(‘~6Ar),,. or ‘QXe. Atmospheric ratios: iHe/’ Ar = 2.33 x IO- ‘, 22Ne/36Ar = 0.0531, R’+KrPOAr= 0.0208, 13ZXe/36Ar = 7.45 x lO-4 (alMA and PODOSEK, 1983). Numbers in brackets Its1 the lo uncertainties of the last quoted digits. “n.d.” indicates R”Kr amount was not determined. ANU 90067b is excluded awmg to large uncertainty in the lhAr amaunt.

0

15 5

F132

lo

FIG.2. Elemental ratios of the noble gases expressed as F-ratios relative to the atmospheric composition (Table 2). In Fig. 2b and 2c the large open square denotes the atmosphere composition (OZIMA and PODOSEK, 1983), whereas the large solid triangle labeled SW indicates the composition of seawater (Z2Ne/36Arsw = 1.44 x 10m2, 84Kr/~ArsW = 4.08 X lo-‘, “2Xe/36ArsW = 2.48 X IO-‘; these values are an average for Atlantic and Pacific deep ocean waters compiled from BIERI and KOIDE, 1972, BIERI et al., 1968, and MAZOR et al., 1964). The large solid black square labeled M in Fig. 2b shows the composition of a tholeiitic melt which has equilibrated noble gases with a seawater-derived vapour phase calculated using the Henry’s Constants of LUX ( 1987). In Fig. 2c, the point M is almost identical to the atmospheric composition and is not plotted for clarity. Tieline joining seawater SW to M denotes the fractionation associated with the solubility controlled exchange of noble gases during seawatermagma interaction. The curved lined extending from M (or Air in Fig. 2c) towards higher values of Fz2 illustrates the elemental fractionation of the noble gases remaining in the melt phase following bubble formation (see PATTERSON et al., 1990).

Isotope ratios of noble gases in phenoczysts loss of volatiies from the melt, can account for the observed variations in the relative elemental abundances of 22Ne, j6Ar, and 84Kr in the New Zealand samples. In contrast to the Fs4 and Fz2 data, the FgS2ratios are systematically two to ten times higher than that predicted on the basis of the simple vapour-magma interaction model (Fig. 2~). The origin of such high Xe/ Ar ratios has been discussed previously (PATTERSON et al., 1990, 1991), and seems likely to reflect an additional Xe-rich component in many of the samples. The source of this Xe-rich component is unclear, but possibly relates to a surface adsorption process which has enriched Xe with respect to the lighter noble gases. Thus, the elemental ratios of the noble gases in the samples are consistent with the presence of an atmosphere-derived component of heavier noble gases. The origin of this atmospherederived component will be discussed in a subsequent section. Helium Isotopic Ratios The helium isotopic measurements reported here show a range of 3Hef4He ratios from about I1 X 10e6 to 6 X 10v6 (Table 1). For the majority of samples the uncertainty in the He composition is less than lo%, and usuaily less than 5% at the la. level. The exceptions are the two clinopyroxene samples from TVZ which have unce~ainties of about 50% owing to their low He concentrations. Overall, the oiivine samples from the intraplate basalt provinces of Northland and Auckland have slightly higher 3He/4He ratios than the olivine samples from the subduction-related TVZ (the average 3He/4He ratios are 9.5 (f0.9) X 10m6and 8.3 (21.0) X lo+, respectively). It is important to note that atmospheric He is unlikely to be a significant com~nent in the samples. Based on an atmospheric 4He/2ZNe ratio of 3.12, and assuming all the 22Ne in the samples is atmosphere-derived, it is possible to estimate the amount of atmospheric 4He in the samples from 4He,i, = 22Ne,l X (4He/22Ne),i, (after CRAIG et al., 1978; SANO and WAKITA, 1989). Such a calculation shows that for the olivine samples atmospheric He is likely to account for significantly less than 1% of the total 4He. In contrast, owing to the observed low He abundances, atmospheric He could account for about 7 and 20% of the 4He in the two clinopyroxene samples ANU 88206 and ANU 88209, respectively. However, because of the large uncertainty in the measured 3He/4He ratios of the clinopyroxenes (Table 1 ), the 3He/4He ratios corrected for atmospheric 4He are indistinguishable from the measured 3He/4He ratios.

Because it is important to compare our new results with the large number of previously published 3He/4He isotopic data for subduction-related samples, a summary of the helium isotopic m~urements of phenocrysts and xenoliths from arc-related rocks reported in this and previous studies is presented in Fig. 3a. Figure 3b, 3c, and 3d present a compilation of 3He/4He ratios previously measured in arc-related fluid samples broken down according to tern~ratu~ (>2OO”C, <2OO*C, and temperature unspecified, respectively) and location (samples from New Zealand and elsewhere). Overall,

4417

‘* t(b) Fluids, T > 200 C 8 !Zl Other t n New Zealand

n = 35

6

4 6 5 3 L

2 25 (c) Fluids, T < 200 C 20 ’

n=241 x=30 .

15 .

3He/4He (1O-6) FIG. 3. Compilation of available 3He/4He ratios for arc-related samples, n = the number of samples included in each plot, x = the number of samples having 3He/4He < 0.5 X lop6 and excluded from plots. Approximate 3He/4He ratio in mid-oceanic ridge basalts (MORB) of 11.8 X IOa(Ku~ 1991) shown for comparison. Figure 3a: Phenocrystic and xenolithic olivine and clinopyroxene measured in this study (solid) and in previous studies of samptes from Andes, Guatemala, Aleutian Islands, Marianas, Indonesia, Japan, and Kamchatka (open and cross-hatched). Rata sources: HILTON and CRAIG ( 1989), HILTON et al. ( 1992,1993a), MARTY et al. ( 1989), POREDA and CRAIG ( 1989), and TOL~TIKHINet al. ( 1972, 1974). Figure 3b, 3c, and 3d: Arc-related fluid samples including gas, steam, and water from volcanic fumaroles, hot and cold springs, geothermal welts, earthquake prediction wells, petroleum and natural gas we& and soil gases from Italy, New Zealand, Japan, and other arcs from the circum-Pacific, broken down on the basis of temperature. Data sources: BASKOV et al. ( 1973), HILTON and CRAIG ( 1989), GIGGENBACH et al. ( 1993), HILTON et al. (1993a), HUL~TON et al. (1986), MAR-II’ and GIGGENBACH ( 1990). MARTY et al. ( 1989), NAGAO et al. (1979, 1980a, 1980b, 1981), POLAK et al. (1982), POREDA and CRAIG (1989), SAKAMOTO et al. (1992), SANO and WAKITA ( 1985, 1989), SANO et al. ( 1982, 1984, 1985, 1987, 1988, 1989), STAUDACHERand ALL%iRE ( 1988), TEoEsco et al. ( 1990), TORGERSENet al. ( 1982), WAKITA and SANOf 1983), and WILLIAMS etaL(1987).

the 3He/4He ratios for phenocrysts and xenoliths from New Zealand form a well defined group falling in the higher range of results for subduction-related phenocrysts and xenoliths

4418

D. B. Patterson. M. Honda, and 1. McDougall

(Fig. 3a). However, it is worth noting that the relatively low 3He/4He ratios reported by HILTON et al. (1992) are for samples from the east Sunda-Banda arcs of Indonesia, a zone of continent-arc collision. These low 3He/4He ratios are interpreted to reflect the underthrusting of radiogenic ‘He-rich continental material beneath the east Banda-Sunda arcs and. therefore, should not be considered repre~ntative of subduction-related samples in general (HILTON and CRAIG, 1989; HILTON et al., 1992; POREDA and CRAIG, 1989). Thus, our new results are in accord with previous measurements of arcrelated mineral samples. Similarly, Fig. 3b shows our results are consistent with helium isotopic ratios reported for high temperature fluids (>2OO’C), particularly for fluid samples from New Zealand which range in ‘He14He from about 7 x IO+’ to 10 X IO-’ (~IGGENBA~H et al., 1993; HULSTON et al., 1986; MAR-I-Y and GIGGENBACH, 1990; SANO et al.,

1987; TORGERSENet al., 1982) and support the general interpretation that He in high temperature fluids is essentially magmatic in origin. In contrast, fluids having lower or unspecified temperatures (Fig. 3c and 3d) show a much wider range in 3He/aHe ratios consistent with greater degrees of shallow or local addition of nonmagmatic radiogenie ‘He derived from local country rock (e.g., SANO and WAKITA, 1985; SAKAMOTO et al., 1992; SANO et al., 1987; TORGERSEN

10.6 10.4 10.2 10.0 9.8 9.6 0.026

0.028

0.030

0.032

0.034

0.036

21Ne/**Ne FIG. 4. Three isotope plot for neon in olivine and clinopyroxene samples from New Zealand. Atmosphere-large open square. Subduction-related phenocrystic samples from Taupo Volcanic Zone: olivine-solid circles, clinopyroxene-open circles. Intraplate samples from Northland and Auckland Volcanic Provinces: phenocrystic olivine-solid squares, xenolithic olivine-open triangles. The dashed line labeled ~?f/ shows the isotopic fractionation trend from the atmospheric neon composition, and the dashed line labekd MORB shows the MORB neon correlation line ( SARDAet al.. 1988). Error

bars show L(r uncertaintyas listedin Table I.

et al., 1982).

Neon Isotopic Ratios

ASSESSMENT OF POSSIBLE IN SITU NOBLE GAS COMPONENTS

Figure 4 presents the neon isotopic data for the New Zealand samples on a conventional neon three isotope plot. Overall the data show a cluster centred on the atmosphe~c neon composition with a small number of results scattering towards higher 2eNe/22Ne and 2’Ne/22Ne ratios of up to 10.60 and 0.0329, respectively. Because the CO:’ interference correction is typically of the order of 3% of the 22Ne peak, it is considered that neon isotopic results which are within 3% of atmospheric values can not be distinguished from atmospheric. In addition, to be regarded as distingu~shabIe from atmospheric, the measured neon isotopic ratios must not overlap with the atmospheric ratios at the la uncertainty level. On the basis of these criteria, seven olivine samples from four localities (P.H.47, Waimarino, Crater Hill, Stone Road) have “Ne/ 22Ne ratios which are clearly nonatmospheric. However, only one sample of xenolithic olivine from the Auckland Volcanic Province (Stone Road ANU 88283) has well constrained nonatmosphe~c 2eNe/22Ne and *‘Ne/“Ne ratios. Interestingly, the Ne results for this sample lie cIose to the MORB Ne correlation line as given by SARDA et al. ( 1988).

Argon isotopic Ratios As listed in Table 1, the 4”Ar/36Ar ratios of the samples range from atmosphere-like values (295.5) to as high as 696 4 26 (ANU 88282). One other sample (ANU 9~67b) has a higher 40Ar/36Ar ratio of 1160 but has an uncertainty of over 100% (& 1320) owing to an extremely large uncertainty in the Ar blank correction. The 38Ar/36Arratios of the samples are generally atmospheric within uncertainties.

Before attempting to interpret the isotopic data in terms of the noble gas composition of the parent or host magmas of the samples, it is necessary to consider the im~~ance of in situ noble gas components generated within the samples since eruption. Of interest are: (a) noble gases produced by radioactive decay, including radiogenic 4He and nucleogenic *‘Ne from the decay of U and Th, and radiogenic 40Ar from the decay of ‘OK,and (b) cosmogenic 3He and *‘Ne produced by the interaction of cosmic-rays with the samples ( KURZ, 1986a,b; LAL, 1987, 1988; LAL et al., 1989, MARTI and CRAIG 1987; PORCELLIet al., 1987; CRAIG and POREDA, 1986). In Situ Radiogenic ‘He and 40Ar and Nucleogenic *‘Ne In order to minim& the effects of in situ components, the samples were collected from youthful lavas. As summarised in Appendix 3, with the exception of one sample which is estimated to be 90 ka, all the TVZ samples are constrained to be younger than 20 ka. The intrapIate samples from Northland and Auckland are generally younger than 100 ka except for one of the Northland samples which is estimated to be 570 ka and two of the xenoliths which are from lavas about 1.3 Ma old. As a result of these young ages, combined with the low U and Th contents of the samples (less than 5 X 10e3 and 15 x 10M3ppm, respectively, as estimated by PATTERSON,1992), in situ radiogenic 4He is unlikely to account for more than 1 or 2% of the total ‘He observed in the samples. Similarly, low K contents (less than 60 ppm K, PATTERSON, 1992) mean that in situ radiogenic 40Ar is unlikely to account for more than approximately 0.75% of the nonatmospheric @Ama

Isotope ratios of noble gases in phenocrysts present in the samples (where nonatmospheric 40Ar,, is defined as equal to 36Arobserved [(40Ar/36Ar)0hNd - (40Ar/ 36Ar),i,]). The exceptions are the xenolithic olivine sample from Ridge Road (ANU 88263), in which radiogenic 4He could account for approximately 7.5% of the total 4He, and the two olivine samples from the Northland Volcanic Province (Puketona ANU 88228, Picadilly ANU 88238) where in situ radiogenic 40Ar could account for about 6.5 and 2.0% of the nonatmospheric 40Ar, respectively. However, these values are upper limits, and do not change the overall conclusion that in situ radiogenic 4He and @Ar are unlikely to have had an appreciable effect upon the observed 3He/4He or 40Ar/36Ar ratios. Similarly, in situ nucleogenic *‘Ne, produced via the reactions ‘*O(a,n)*‘Ne and 24Mg(n, Lu)*‘Ne ( WETHERILL, 1954), could generate no more than 0.1% of the *‘Ne observed in the samples (calculated assuming a production ratio of 10 X 106; KYSER ‘Heradiogenic/*‘Nenucleogenic and RISON, 1982). Thus, in situ nucleogenic *‘Ne can not have had a significant effect upon the observed *‘Ne/**Ne ratios. In Situ Cosmogenic ‘He and “Ne

Various estimates of the production rate of cosmogenic 3He in surface exposed olivine samples have been made (CERLING, 1990; KURZ, 1986b; KLJRZ~~al., 1990; LAL, 1991; POREDAand CERLING, 1992). For the following discussion we use a value of 107 3He atoms g-’ a-’ (corrected to sea level and latitudes 2 60”) as determined by CERLING( 1990) and POREDAand CERLING ( 1992) in well-dated lava flows from the western United States. Note that this value is intermediate to the larger range of values (82-190 atoms g-’ a-‘, corrected to sea level and latitudes > 60”; LAL, 1991) given by KURZ et al. ( 1990), but is higher than theoretical values (51-68 atoms g-’ a-‘) in endmember Fe2Si04 and Mg2Si04 calculated by LAL ( 199 1). Based on this value, and using the latitude correction of LAL ( 1991), we estimate the sea level production rate of cosmogenic 3He in the central and northern North Island of New Zealand (about 40’S) to be approximately 100 3He atoms g -’ a-’ or 5.5 X lo-‘* cm3STP g-’ a-‘. We have used the geographic latitude instead of the magnetic latitude because the wandering of the magnetic pole over periods greater than 15 ka should average the magnetic field intensity to that of the present-day latitude ( POREDA and CERLING, 1992). The production rate of cosmogenic ‘He in each sample, corrected for altitude ( LAL, 199 1) , and for depth-below-surface assuming the overburden to have an attenuation length of 160 g cm-* (LAL, 1991; YOKOYAMAet al., 1977), is listed in Table 3. In order to minimise the effects of cosmogenic 3He production, all the samples were collected from recently exposed outcrops such as roadcuts, quarry faces, and river banks having at least 2 m of dense rock overburden. By combining the estimated production rate with the maximum geological age (Appendix 2), an estimate of the maximum amount of cosmogenic 3He that could be produced in each of the samples can be made. As shown in Table 3, for the majority of samples cosmogenic production could generate at most 2 or 3% of the total 3He. The exceptions are the samples which may be

4419

Table 3: Estimate of maximum in siru cosmogenic ‘He.

Location

ANU#

Altitude

D.B.S

Prcducdon

(m)

(m)

Rate

Taupe Volcanic Zone: Phemmysric 88242 300 Mahuia 88207 840 P.H.47 88209 900

Wabnarino

Tat&mlolanic P.H.47

SQO

9CO67b 88244a 88244b 88244c

900 580 580 580

Zan;&Phenocryy’c 88209

’ HeJ'He, (%)

olivine

OngaIvto

9c067a

‘He, (IO-)

2.5 >2 3 3 3 >2 >2 >2

4.9 x lo-~ 2.1 x 10’9 2.9 x 10-2”

49 42 6

7.3* 0.35 0.08

2.9 x K-2”

6 6 34 34 34

0.10 0.10 2.6 2.7 3.2

2.9 x 1.7 x 1.7 x 1.7 x

10-m 10-19 10-19 I@‘9

clin;piyf;one

900

3

Norlhlmd Volcanic Province: Phenocrystic olivine Puketona 88228 200 >3 Picadilly 88238 200 >3 Auckland Volcanic Province: Phenocrystic olivine Wiri 88265 30 >2 Crater Hill 88269 50 1.5 Puketutu 88282 10 >2 Auckland Volcanic Province: Xenolithic Lake Pupuke 88276 5 RidgeRoad 88263 100 StoneRoad 88283 100

negligiblet 2.9 x IO-l9

6

14

1.6 x 1O-20 1.6 x 10-2”

91 9

2.3 3.0

1.0 x 10-19 2.9 x IO-l9

28 87

9.8 x lo-20

75

1.0 2.2 7.4”

olivine 1.5

2.7 x lo-l9

> 4 > 2

2.0 x 10-Z’ 1.1 x 10-19

270 24 1500

8.F 1.7 5.1

Gas abundances in cm’STP g-‘, production rates in cm’STP 8-l a-‘. ‘He. is the predicted abundance of cosmogenic ‘He in the samples calculated assuming maximum geological age as listed in Appendix 2. ‘Hq/‘He, is the calculated amwnt of cosmogenic‘He exprssed as a percentage of the total )Hq. D.B.S. denotes depth below surface. Altitude is the height above sea level as estimated from I : 50,CCJLl scale topographic maps. * indicates calculated amount of cosmogonic‘He is likely to be too high becauseof poorly constrained maximum geological age. # Production of cosmogenic ‘He in ANUl 88206 is negligible owing to large thickness of overburden (> 20m ).

significantly older (i.e., Ongaroto, Puketutu, Lake Pupuke, Stone Road), and one of the TVZ clinopyroxene samples which has very low total 3He (P.H.47). However, it is stressed that the calculated amount of cosmogenic 3He in each sample is a maximum value based on a maximum estimate of geological age and assuming there has been no removal of overburden during the history of each of the samples. Thus, production of cosmogenic 3He may have had a small effect upon the 3He/4He ratios of most of the samples, and is likely to have been less than 10% in the remainder. In addition to 3He, the interaction of cosmic-rays with the samples can also generate observable amounts of cosmogenic *‘Ne (GRAF et al., 1991; MARTI and CRAIG, 1987; NIEDERMANNet al., 1993; POREDA and CERLING, 1992; STAUDACHERand ALL~?GRE,1991). Taking the maximum cosmogenic 3He, as listed in Table 3, and assuming a cosmogenic 3He/2’Ne production ratio of 2.1 in olivine ( MARTI and CRAIG, 1987), it can be shown that cosmogenic *‘Ne could account for a maximum of 14% of the nonatmospheric 2’Ne,a observed in the samples (where *‘Ne, is defined as **Neobsclved [(2’Ne/22Ne)o~serv~ - ( 2’Ne/22Ne).i, I). Taking the more recent cosmogenic ‘He/*‘Ne production ratio of 2.44 estimated by POREDA and CERLINC (1992) results in slightly lower estimates of the maximum cosmogenic *‘Ne. It is, therefore, concluded that production of cosmogenic *‘Ne can not account for the observed nonatmospheric 2’Ne/22Ne ratios. Thus, the isotopic compositions of He, Ne, and Ar observed in the samples are not significantly influenced by in situ radiogenic/nucleogenic and cosmogenic components.

4420

D. B. Patterson, M. Honda, and I. McDougall ELEMENTAL FRACTIONATION: ‘He AND NONATMOSPHERfC 40Ar AND 2’Ne

7 (4

6

As noted earlier, the elementat F-ratios of the nonradiogenie noble gas isotopes (i.e., 3He, ‘*Ne, %Ar, %r, and ‘jzXe; Fig. 2) reflect the presence of dominant atmosphere-derived component of the heavier noble gases. Further, the variations in the F-ratios of the heavier noble gases are broadly consistent with solubility controlled elemental fractionation. Additional information regarding the elemental fractionation of the noble gases in the samples can be obtained by comparing the relative abundances of radiogenic 4He, and nonatmosphe~c ‘OArand 2’Ne ( referred to as ‘@Ar,,,and ’ INens, as previously defined ) observed in the samples with the likely ‘He/*Ar, ‘He/“Ne, and *‘Ne/@Ar nuclear production ratios in MORB-source upper mantle and crustal materials. Taking a typical MORB-source composition (K/U = 12,7OOandTh/U=2.5,J~~~~etal., 1983),anda’He/ *‘Ne production ratio of 10 X lo6 ( BALLENTZNEet al., 199 1; KYSER and RISON, 1982 ) , yields present-day instantaneous ‘He f *Ar and “Ne f *Ar nuciear pr~u~on ratios of 3.9 and 3.9 X lo-‘, respectively. For a typical crustal composition (K/U = 12,70OandTh/U = ~.~,TAYLOR~~~MCLENNAN, 1985), we obtain slightly higher present-day 4He/40Ar and *‘Ne/“Ar production ratios of 4.6 and 4.6 X 10W7,respectively. This value for the crustal ‘He/““Ar production ratio is consistent with the value of 4.08 +- 0.33 for the crustalderived component observed in geothermal fluids from Yel-

ANUX

4He/“oAr,,

4He/2’N%, (106)

Taupe Volcanic Gngsroto Mahuia P.H.47

Zone: ;8$;crystic

Waft

Ta~~~;dc

88209

Nort~c$wdc

Prf;l;i;c;:

Picadilly Volcanic

Ahland Voicmic Lake Pupuke L%%

:oz

Estimated production MORB-source CNStal

Province: 88276

(IO-‘)

6.4 ad. 4.6 0.79 1.31 1.14

(311 130) (66) (22) (24)

3.0 (16) n.d. § 1.7 (14) ;::

I:!;

clinopyroxene 0.85 (49) 0.11 (6)

Phenoc~stic olivine 6

I:;

Phenoey;~;

88282 88269

~~~~’

2; (2)

8.13 (1) 0.30(13) 0.29(15)

88238 Pm;v6;

2’Ne,&“‘Ar,,

(8)

;: 1:s

Zone: ;8teevstic

P.H.47

A~~~

olivine 2.9

88207 88209 90067a 90067b 88244a 88244b 88244~

olive 0.17 1.7

(4) (2)

1.23 (IO)

13.7 (36)

Xenolithic &vine 1.21 (7)

88263 88283 ratios 4.0 Oa Present-day 4.0 Ga

Present-day

4.73 (26)

::; 1.9 4.6

10 10 IO t0

12

3

4

5

6

7

8

91011

411e/21Ne,, (106) FIG. 5. 4He/40Ar,. and “He/“Ne,, ratios of phenocrystic and xenoiithic olivine and clinopyroxene from New Zeatand. Symbols as for Fig. 4. Error bars show I (I uncertainty as tisted in Table 4. Figure

Table 4: Ratios of total 4He, and non-atmospheric 4oAr,,, and *‘Nt+ Location

0

2.2

(2)

1.6 3.9 1.9 4.6

Numbers in brackets give fa uncertainty in last quoted digits. 2iNe,, and *AI., are non-atomspheric *tNe awl 4o.k as d&c& in text. n.d. indicates neon isotopic ratio was not detemdoed. 0 iodicates that @AT,, was not calculated owing to an uncertainty in %rPAr of greater than 100 percent. Estimated production ratios calculated for an accumulation time of 4.0 Oa and for prcwuda imaotaocoos production in the upper maotk MORB-source (K/U = 12,700 and Thh = 2.5, JOCHUM et al 1983) and in typical crustal material (K/U = 12,700 and TbRl = 3.8, TAYLOR ad MCLENNAN, 1985). assuming a radiogenic/nucleogenic 4He121Ne production ratio of 10 x 106 (BAIxzmm~ et al., 1991: KYSER and RtSON. 1982) and using the decay comaots for 231Tb, 23%J, U’IJ, and % and isotopic data for U and K of STEIL~ERand JAmR (1977).

4a: Comparison of ‘He/*Ar”,, ratios with the estimated ‘HejQAr nuclear production rates in typical crustat and MOREsource mantle materials for the present-day and for an accumulation time of 4.0 Ga (Shown as horizontal lines; see Table 4 and text for discussion of these production ratios). Figure 5b: Plot of 4He/40Ar., vs. %e/ “Nena for samples having nonatmospheric “Ne/“Ne ratios.

lowstone (KENNEDY et al., t 985). Alternatively, taking an integrated accumulation time of 4.0 Ga (a maximum upper limit) yields lower 4He/WAr and ZiNe/40Ar production ratios; 1.6and I .6 X lo-‘, respectively, in the upper mantle MORBsource, and 1.9and 1.9 X 10e7, respectively, in typical crustal materials. These limits for the nuclear production ratios of ‘He/‘*Ar, 4He/2’Ne, and 2’Ne/40Ar are listed in Table 4, along with a compilation of 4He/40Ar,,a, 4He/2’Nena, and Z’N%,/40Ar,, ratios observed in the New ZeaIand samples. As shown in Table 4 and plotted in Fig. Sa, the 4He/40Ar,,, ratios of the samples are highly variable. The minimum ‘He / 40Ar,, ratio of 0.11 is an order of magnitude lower than the likely limits of the 4He/40Ar nuclear production ratio, whereas the maximum ‘He/‘*Araa ratio of 18 is at least four times higher. Note, however, that the majority of samples have ‘He/“Ar,, ratios that are low relative to the estimated production ratio. The 4He/2’Ne,, ratios of about 0.8 X IO6 to 6.4 X 106 in the samples having clearly nonatmospheric *‘Ne/ 22Ne ratios are systematically lower than the estimated ‘We/ “Ne nuclear production ratioofabout IO X 106. Importantly, these low ‘He/*‘NG, ratios are correlated with low 4He/40Ar,,a ratios (Fig. Sb). In contrast, the observed Z’Ne,/40Ar, ratios of I .7 X lop7 to 3.0 X 10T7 (excluding ANU 88269 which

Isotope ratios of noble gases in phenocrysts has a relatively high 2’Ne,,,/40Ar,, ratio of 13.7 X 10s7; Table 4) are similar to the estimated nuclear 2’Ne/40Ar production ratio of between 1.6 X 10 -’ and 4.6 X IO-‘. Taken together, these results are indicative of a significant and systematic elemental fractionation of 4He from nonatmospheric 21Ne, and *Ar,=, but relatively little elemental fractionation of ‘IN%, from @Ar, in the New Zealand samples. Note that i similar systematic loss of 4He from nonatmospheric 40Ar has been previously observed in ultramafic phenocrystic and xenolithic samples from a variety of locations (FISHER, 1983; KANEOKA and TAKAOKA, 1978,198O; KANEOKAetal., 1983, 1986; KYSER and RISON, 1982; POREDAand FARLEY, 1992; STAUDACHERet al., 1986, 1990). It is interesting to briefly consider the cause ofthis variation. At least three po~ibili~es can be suggest& (a) the samples equilibrated noble gases with magmas that had low 4He/WAr,, ratios, (b) the low 4He/40Ar,,, ratios reflect preferential trapping of Ar with respect to He in the crystals, and (c) there has been preferential loss (by diffusion?) of He with respect to the heavier noble gases from the crystal. However, although it is possible to suggest these possibilities, it is unclear which, if any, is the cause of the observed low 4He f #Ar,,, (and 4He/ “Ne,,,) ratios. For the moment we simply note that there appears to be a systematic loss of He with respect to Ne and Ar in the samples. HELIUM RESULTS: MIXING OF MORBLIKE CRUSTAL

AND

COMPONENTS

As noted previously, the ‘Hef4He ratios of arc-related fluids range from MORB-like values of about 12 X 10V6down to values that are lower #an the atmospheric ratio of 1.4 X 10-6 (Fig. 3). Thus, the range in 3He/4He ratios of about 11 X 10m6to 6 X 10m6for mafic phenocrysts and xenoliths from New Zealand is consistent with the published He results for arc-related fluid samples and can be interpreted in the same fashion. Specifically, mixing of a mantle-derived component rich in primordial ‘He and character&d by a MORB-like ‘He14He ratio of about 12 X 10e6 with a radiogenic 4Herich component derived from crustal materials and characterised by a low 3Hef4He ratio (<2 X lo-‘) can account for the observed ratios. The suggestion that the upper mantle sources of the samples includes a MORB-like component is extremely important and is based primarily on the observations that (a) the upper mantle, as sampled by N-type mid-ocean ridge volcanism, is believed to be globally uniform with respect to helium isotopic com~sitions with 3Hef4He ratios of about 12 X 10m6(e.g., KURZ, 1991), and (b) none of the fluid, phenocryst, or xenolith samples reported in this or previous studies (Fig. 3) have 3He/4He ratios higher than the MORB value. Further, theoretical modelling of the flow of mantle material above a subducted slab suggests that viscous drag along the upper surface of the downgoing slab leads to an induced comerflow regime (RIBE, 1989; SPIEGELMANand MCKENZIE, 1987). Becausethis induced flow is likely to introduce fresh upper mantle material into the mantle wedge, it is argued that noble gases observed in basaltic glasses from the backarc environment are representative of the mantle wedge material before it became metasomatised by fluids released from

4421

the downgoing slab. Thus, direct evidence for MORB-like noble gases in the upper mantle near subduction zones is found in the observation of MORB-like helium isotopic compositions in basaltic glasses dredged from the Mariana Trough ( POREDA, 1985; SANO et al., 1986), and of MORBlike He, Ne, and Ar components in basalt glasses from the Lau Backarc Basin (HILTON et al., 1993b; HONDA et al., 1993b; POREDA, 1985). NEON AND ARGON RESULTS: MIXING OF THREE COMPONENTS

As discussed previously, the ““Ar/36Ar and 2’Ne/22Ne ratios of the samples range from atmosphere-like to hiier-thanatmosphe~c values. Because in situ generated 40Ar and “Ne is unlikely to be significant, this variability must reflect mixing of atmospheric and nonatmospheric noble gas components. However, given that the helium in the samples is a mixture of two nonatmospheric components derived from mantle and crustal sources, it is reasonable to argue that the nonatmospheric Ne and Ar in the samples is also a mixture of mantle and crustal-derived components enriched in radiogenic 40Ar and nucleogenic “Ne generated within the earth over geological time. Thus, the Ne and Ar in the samples is interpreted to be a mixture of three components derived from the MORB source upper mantle, crustal materials, and the atmosphere. As noted in Table 4, the 2’Ne/40Ar nuclear production ratio is likely to be similar in the MORB-source upper mantle and crustal materials. Therefore, although we can identify the presence of nonatmospheric Ar and Ne components in the samples, it is not possible to distinguish the relative importance of the MOR~iike and crustal sources on the basis of the observed nonatmospheric 40Ar/36Ar and 2’Ne/22Ne ratios. However, by using helium isotopic signatures, which can be attributed to a mixture of MORB-like mantle and crustal components, it is possible to quantitatively deconvolve Ar in the samples into three distinct components: atmospherelike, MORB-like, and crustal-derived. ~ontibution of Each Com~nent to Total *oAr Because the 40Ar/36Ar ratio in MORBs is commonly very high (up to 28,000, STAUDACHERet al., 1989) compared with the atmospheric value of 295.5, and crustal-derived Ar is predominantly radiogenic in origin, it is reasonable to assume that virtually all the 36Ar in the samples is atmospherederived. Also, because the presence of significant in situ radiogenic argon in the samples can be exduded, the relative cont~bution of each com~nent to the total *Ar ( 4aArt) can be expressed as follows: 40Ar,/40Ar, = A(36Ar/40Ar),, 40Ar,/40Ar, = [l - A(36Ar/40Ar)0](40Arm/40Ar,.),

(1) (2)

and @ArJ40Art = [l - A(~Ar/~Ar)~][l

- (40Ar,/40Ar~)],

(3)

where &Ar., 40Ar,,,, and 4oArCare the contributions derived from the atmosphere-like, MORB-like and crustal-derived components, respectively, 4oAr, = WAr,, + “OAr,,, and 4oAr,,

D. 8. Patterson. M. Honda, and 1. McDougall

4422

= 40Ar,,,+ @Ar,, A is the atmospheric 40Ar/36Ar ratio (295.5 ), and ( 36Ar/40Ar)o is the observed 36Ar/ 40Ar ratio (after PATTERSON, 1992). Because atmospheric He is not an important component in the samples and helium isotopic ratios in the MORB-like and cm&-derived components are significantly different, it is possible to deconvofve the MORB-like and crustal-derived components of He. Assuming all 3He in the samples is primordial and derived from an upper mantle source having a MORB-like 3He/4He ratio of 12 X 10m6, and the crustalderived component is pure 4He (i.e., 3He/4He,,,,,,l < 0.1 x 10e6; MAMYRIN and TOLSTIKHIN, 1982), the relative contribution of the MORB-like component can be expressed as

0 A

a!

4He,f4He, = M(3He/4He),,

(41

where 4He,,,/4He, is the relative contribution of the MORBlike component to the total observed 4He, A4 is the 4He/3He ratio of MORB (83,330), and (3He/4He), is the observed 3He/4He ratio. Given the similar K/U ratios of upper mantle and crustal materials, we can make the reasonable assumption that the 4He/40Ar ratios of the MORB-like and crustal-derived components are likely to be similar, (Table 4). thus 40Ar,/ 40Ar,, = 4He,/4He, = M( 3He/4He), and Eqns. 2 and 3 can be rewritten as 4oAr,/40Art = [ 1 - A(36Ar/4”Ar),]M(3He/4He),,,

(5)

and

= [I - A(36Ar/4~Ar)*][l - ~(3He/4He)*].

(6)

Clearly this calculation is dependent upon the assumption that the 4He/40Ar ratios of the MORB-like and crustal components in the sample are similar. Elemental fractionation of He from Ar will have little effect provided that the noble gases derived from both MORB-like and crustal sources are fractionated in the same event. If the systematic loss of He with respect to Ar (and Ne) observed in the samples occurs at a late stage and shallow levels (e.g., by diffusive loss of He from the crystals), then this assumption is likely to be valid. Further, note that this calculation only applies to the 40Ar in the samples, and not to the other noble gas isotopes in general (e.g., the calculation assumes all the 36Ar is atmosphere-derived and all the 3He is mantle-derived). Figure 6 presents a ternary plot showing the relative contributions of each of the three components to the total 40Ar in the samples calculated on the basis of Eqns. 1, 5: and 6. The dominance of the atmospheric com~nent is clearly apparent, accounting for between 50 and 100% of the total @Ar. Of the nonatmospheric components, 40Ar derived from the assumed MORB-like mantle source generally dominates, accounting for a maximum of 50% of the total 40Ar. The question of the mechanisms by which these noble gas components are introduced to the samples is discussed in the following sections. THE NOBLE GAS CYCLE IN SIJBDUCIION

Atmospheric

SYSTEMS

The presently favoured model for the generation of arc magmas involves a dominant mantle wedge component as

MORB-iike

0

f

Crustal

FIG. 6. Ternary plot showing the relative percentages of atmospheric, MQRB-like, and crustal-derived ““Ar in the samples as calculated from Eqns. I, 5, and 6. Symbols as for Fig. 4. Plot excludes the two TV2 clinopyroxene samples which plot almost at the atmospheric apex owing to their atmosphere-like *Ar/36Ar ratios, and the TV2 olivine sample ( ANU# 90067b) which has an uncertainty of greater than 100 percent in its measured 40Ar/sAr ratio.

well as input from subducted crust, sediment, and seawater (e.g., WOODHEAD,1988; TATSUMI, 1989). In addition, subsequent processes such as crystal fractionation and settling, assimilation of crustal material, and magma mixing, can significantly alter the composition of arc magmas and obscure the initial primary composition. It is, therefore, necessary to consider how these processes may influence the noble gas composition of the parent magmas of the New Zealand sampies, and more specifically, the mechanisms by which each of the three noble gas components identified above may be incorporated in the samples. Based on the present understanding of the origin and generation of arc magmas, a qualitative schematic box model of how noble gases may be transported through subduction systems and introduced to the parent magmas of the samples is presented in Fig. 7. The three identified noble gas ~rn~nen~ (MORB-like, atmosphere, and crustal) are represented by the boxes labelled in the large upper case letters. The box labelled SAMPLE represents the noble gases trapped in the samples as measured in the laboratory. The arrows trace the possible routes by which noble gases may enter the samples. There is clearly no difficulty in incorporating MORB-like noble gases in the samples if the parent and host magmas are derived from partial melting of an upper mantle source which includes a MORB-like noble gas component. In contrast, the mechanisms by which atmosphere-like and crustal-derived noble gases may be incorporated in the magmas are more complex, and include addition to the source regions of the magmas through subduction processes and shallow interaction between ascending magmas and crustal materials. These possibilities are described in greater detail below. Atmospheric and crustal-derived noble gases may be introduced to the mantle source regions of the magmas through the process of subduction. Seawater contains relatively iarge concentrations of dissolved atmospheric heavier noble gases (Ne: 1.8 X lo-‘; Ar: 3.5 X 10e4; Kr: 8.4 X lo-*; Xe: 1.1

Isotope ratios of noble gases in phenocrysts

4423

Arc-related magma in crustal magma

Partial me/thg and ascenf

FIG. 7. Schematic model of the pathways by which atmosphere, crustal, and MORB-like noble gases may enter the parent magmas of the samples. The three main components are shown in large upper case letters. The box labelled SAMPLE refers to the noble gases trapped in the samples as measured in the laboratory. See text for discussion of this model.

X 10-s cm3 STP per gram of seawater; listed concentrations are an average of Atlantic and Pacific deep ocean waters compiled from BIERI and KOIDE, 1972, BIERI et al., 1968, and MAZORet al., 1964). Seawater is likely to be incorporated into the downgoing slab in a number of ways including adsorption onto the surface of sediments, direct incorporation in pores and fractures in sediments and crust, and chemical incorporation in hydrous authigenic and alteration mineral phases in sediments and hydrated oceanic crust (Fig. 7). Crustal-derived radiogenic 4He and &Ar and nucleogenic “Ne are likely to be present in the old subducted oceanic crust as well as being incorporated in the downgoing slab by the subduction of terrigenous sediments derived from continental crust. Subsequent generation of hydrous and/or siliceous mobile fluids from the dehydration and/or partial melting of the crust and sediments of the downgoing slab will transport subducted atmospheric and crustal-derived noble gases into the overlying mantle wedge where they will become incorporated in the mantle-source of the parent and host magmas of the samples. However, not all noble gases which are initially subducted may become available for incorporation in the mantle source of the samples. For example, on the basis of theoretical and experimental studies, TATSUMI (1989) suggested that the majority of the dehydration of the downgoing slab occurs at relatively shallow depths below the forearc region, probably at depths of less than 50 km. It is, therefore, possible that there would be significant release of noble gases from the subducted slab at shallow depths, with recycling to the atmosphere through the forearc. In addition, it is possible that some subducted noble gases may be retained within the downgoing slab and recycled into the deeper mantle. Both of these po~ibilities would reduce the amount available for inclusion in the SOUIXX region of arc magmas, and are included in Fig. 7. An alternative way of introducing atmospheric and crustal noble gases is by interaction between crustal materials and

the magmas during their ascent through the crust. This might occur by direct assimilation or stoping of crustal material into the ascending magma, or by diffusion of crustal noble gases into the magma from the wall rocks of magma conduits and chambers. Interaction between the magma and downward percolating crustal fluids that carry dissolved atmospheric noble gases as well as radiogenic and nucleogenic component derived from the surrounding crustal country rocks (PATTERSONet al., 1990, 199 1) can produce similar effects. Thus, it is argued that there are several ways by which atmospheric and crustal-derived noble gases may be incorporated in the parent magmas of the samples. In the following section we discuss whether or not it is possible to distinguish which of these mechanisms is likely to be the most important. DISCUSSION

As previously shown, the isotopic and elemental ratios of the noble gases enable us to identify the presence of crustal and atmosphere-derived noble gas components in the samples and to conclude that these components have undergone some degree of elemental fractionation. However, the noble gas data alone do not provide much information as to how these components are introduced. In this section we further discpossible sources for the crustal and atmosphere-derived noble gas components observed in the samples. As previously noted, dehydration reactions at shallow depth (~50 km) may extract noble gases from a downgoing slab and recycle them to the atmosphere through the forearc, thus preventing subduction processes from introducing these noble gas components into the mantle source regions of arc magmas. Direct evidence for shallow dehydration is provided by the discovery of seamounts consisting of highly serpentinised ultramafic material in the forearc region of the Mariana Trench. FRYER et al. ( 1985) suggested that the major source of water involved in this metamorphism is from dehydration

4424

D. B. Patterson, M. Honda, and I. McDougall

of the downgoing slab, and they argued that this dehydration

should be complete before the slab descends below depths of about 30 km. More recently, in a study of samples from the Valu Fa Ridge in the Lau Backarc Basin, HILTON et al. ( 1993b) noted that although the lead and strontium isotopic data clearly indicate addition of a slab-derived (sedimentary) component to the samples, the helium isotopic systematics are in no way coupled to the Pb and Sr results. As a result, HILTONet al. ( 1993b) argued that the crustal and atmospheric noble gas components observed in these samples can not be derived from the subducted slab. Because the Valu Fa Ridge lies at about the same distance (approximately 150 km) above the slab and the region of melt or fluid generation as most other subduction arcs, HILTON et al. ( 1993b) further suggested that the downgoing slab must have been stripped of noble gases at shallower sub-forearc depths. If this interpretation can be applied to subduction systems in general, it implies that the atmospheric and crustal-detived components in the New Zealand samples are more likely to be derived from interaction between the ascending magmas and surrounding crustal materials. Although there is no direct evidence suggesting stripping of the subducted noble gases from the downgoing slab at sub-forearc depths in the case of New Zealand, a number of arguments can be used to support the suggestion that the atmosphere and crustal-derived components in the New Zealand samples are introduced to the samples by interaction with crustal materials at relatively shallow depths. First, petrological modelling of the major and trace element compositions and 87Sr/86Srisotopic ratios of lavas and crustal xenoliths indicate the addition of a significant crustal component to many of the subduction-related magmas of the Taupo Volcanic Zone (GRAHAM and HACKETT, 1987; PATTERSONand GRAHAM, 1988). In the case of the intraplate samples, MCDOUGALL et al. (1969) found that the K/Ar ages of the basalts of the Auckland Volcanic Field were significantly higher than associated 14Cages and incompatible with the youthfulness of the lavas inferred on geomorphological grounds. MCDOUGALL et al. (1969) attributed this to the presence of excess radiogenic 40Ar, derived from incorporation of crustal materials in the basaltic magmas. These two observations are at least consistent with shallow addition of a crustal-derived noble gas component to the samples. Second, one of the primary reasons that this study included a suite of intraplate samples from the behind-arc Northland and Auckland Volcanic Provinces was to test whether there was a systematic change in the noble gas composition of the samples with increasing distance from the subduction system. However, there is no systematic difference in the noble gas results from the subduction-related and intraplate samples; both sets of results are dominated by the atmosphere-like heavier noble gas component. This suggests that the processes which introduce atmospheric and crustalderived noble gases to the samples are independent of distance from the subduction zone, and may be related instead to the ascent of the magmas through the overlying crust. It is important to also consider an alternative situation where the downgoing slab is not stripped of noble gases at subforearc depths, but rather, releases subducted noble gases into the source region of the samples. If this were the case,

then the similarity of the heavier noble gas results from the intraplate and subduction-related samples may be simply because the upper mantle source for both the subduction-related and intraplate samples is dominated by a subducted atmospherederived heavier noble gas component, and would not require the shallow addition of atmospheric and crustal-derived noble gases to the samples or their precursors. Such an explanation would be consistent with the observation that the intraplate samples were collected from a region where subduction is believed to have occurred over an extended period between about 15 and 6 Ma ago (BALLANCE,1976; BALLANCEet al., 1982; COLEand LEWIS, 198 1). This would also more easily explain the atmosphere-like heavier noble gases observed in the intraplate mantle-derived ultramafic xenoliths. In conclusion, we argue that it is not possible to confidently distinguish between shallow interaction or subduction as the most likely origin for the crustal and atmosphere-derived noble gas components in the New Zealand samples. SUMMARY This study has confirmed the presence of MORB-like helium in arc magmas, and has shown that the heavier noble gases in subduction-related and intraplate samples from New Zealand can be accounted for by mixing of a dominant atmosphere-like component with MORB-like and crustal-derived components. However, the evidence regarding the mechanisms by which the atmospheric and crustalderived noble gas components are introduced to the samples remains equivocal. Clearly, further studies of phenocrystic and xenolithic samples from other subduction-related lavas are required to help resolve these important questions. Acknowledgments-We

would like to acknowledge the assistance of Drs. J. A. Gamble, 1. E. M. Smith, and R. M. Briggsin identifying and collecting suitable samples in New Zealand. We greatfully acknowledge the reviews of the manuscript by Drs. B. Marty, D. R. Hilton, and R. J. Poreda; their comments have helped us to considerably improve the paper. D. B. Patterson also wishes to acknowledge the support ofthe Japan Society for the Promotion of Scienceduring the preparation of this manuscript.

Editorial handling: D. E.

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4426

D. B. Patterson, M. Honda, and 1. McDougall

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Appendix I: Sample location and rock type Sample

ANU#

4427

WALLCOTTR. I. ( 1984) Reconstructions ofthe New Zealand region for the Neogene. Paleogeogr... Puleoclimatol., Paleoecol. 46,2 17231.

WETHERILLG. W. ( 1954) Variations in the isotopic abundances of neon and argon extracted from radioactive minerals. Phys. Rev. 96,679-683.

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Appendix 2: Sample masses. extraction temperatures, grain sizes. and Mg/Mg+Fe

Ma@! Grid Ref.

Rock Type*

Location

ANU#

Mass fg)

Temperature (‘C)

Grain Size5 (mm)

ratios

Mg/Mg+Fe+ (mole percent)

Taupe Volcanic Zone: Subduction-related samples OlElkUlle Mahuia P.H.47 Ongtinlt0 Wuimatino

88206

s20

173979

88207 882c9.90067$ 88242 88244

S19 S19 T17 T19

268257 296300 668075 643387

Taupe Volcanic Zone: Phenocrystic olivine ongatoto 88242 I.511 1875 Mehuia 88207 1.547 1900 P.H.47 88209 1.424 1875 90067a 1.445 1450’ 90%7b 0.973 1450* Waimarino 88244a I.571 1875 88244b 1.975 1500* 8824‘k I.555 14M)*

cpx endesite cpx-pl undesite cpx-pl en&site 01 basalt ol-cpx basalt

Northland and Auckland Volcanic Provinces’ Intraplafe samples Puketone Picadilly Wiri Crater Hill Puketutu Lake Pupuke Ridge Road Stone Road

88228 88238 88265 88269 88282 88276 88263 88283

PO5 Nl5 RI1 RI1 RI I RI1 R12 RI2

998559 315258t 759651 732667 652692 674894 868365 872646

al-pl basalt PI-01 basalt al-cpx basalt ol-cpx basalt 01 basalt xenolith bearing lapilli tuffg xenolithic aphyric basult xenolith bearing 01 basalt

0.42

-

85.5 - 89.5 86.3 89.2 91.2 94.9

- 0.70

89.9

Toupo Volcanic Zone: Phenocry&clinop~;ar&ne Ohektlne 88206 P.H.47 88209 1.596 1600

0.70 I.2 0.70 - I.7

Northland Volcanic Province: Phenocrvsdc o/iv&! Puketona 88228 1.525 1500’ Picadilly 88238 1.598 1500’

1.0 I.0

Auckland Volcanic Province: Phenmysric &vine Wiri 88265 1.488 1500’ Crater Hill 88269 1.754 1500: Puketutu 88282 I.613 1500*

0.42 0.42 0.40

Auckland Volcanic Province: Xenolirhic olivine I.710 Lake Pupuke 88276 1500’ Ridge Road 88263 I.210 1500’ 1.546 Stone Road 88283 1500’

1.0 - 3.0 0.42 I .O 0.70 I .7

’

* Rock type descriptions ate based on the dominant pknccrysts present: 01 - olivine. cpx clinopyroxene, pl - plagioclese. $ P.H.47 wes sampled twice because the initial sample (ANUX 88209) did not yield sufficient olivine to allow multiple analysis. + All grid references are for New Zealand Map Series 260 (1:50000). except ANU# 88238 which is for New Zealand Map Series I (1:63360). k The Luke Pupuke xenolithic sample consists of lava encrusted dunitic nodules and fragments handpicked from crystal lithic lapilli tuff beds surrounding the Lake Pupuke explosion crater. The Ridge Road end Stone Road xenolithic samples were extracted from dense basaltic host laws by ctushing.

0.42 - 0.70 0.42 - I.0 0.70 I .7

- 1.7

-

- 91.2

81.5 83.4 70.9 - 72.9

1.7

81.0 87.7 80.5 - 86.9

- I .7 - I .O - 0.70

83.3 83.0 83.2

-

- 86.4 85.3 87.0

90.5 91.3 74.6 - 92.8 90.8 93.5

g Range in grain sizes us defined by standard sieve mesh. t Range in MgiMg+Fe ratio! expressed in mole percent, determined for between 20 and 30 individual grains usmg energy dispersive electton microprobe. * Indicates gas was extracted from sample by dissolving the olivine in approximately log of previously outgassed basaltic glass flux.

Appendix 3: Maximum geological ages

Sample

ANU#

Age (ka)*

Dating Technique

source**

Taupe Volcanic Zone: Subducrion-relared samples Oh&Ule Mahuia P.H.47 0tYgtU0t0 Waimurino

88206 88207 88209.9co67 88242 88244

20 19.8 19.8 91 (30) 19.8

Tephra stratigraphy : K-At Tephta stratigraphy

Northland and Auckland Volcanic Provinces: lnlraplare samples Puketona Picadilly Wiri Crater Hill Puketutu Lake Pupuke Ridge Road Stone Road

* ** $

t

g

88228 88238 88265 88269 88282 88276 a8263 88283

K-Ar K-At c-14 c-14 K-Art @ewhnlw5 K-At

b b

Numbers in brackets listed quoted uncetluinties in last digits. ~-TOPPING (1973) as listed by GRAHAM and HACKRT (1987). b SDPP (l968), c PA-SON (1992). d - MCDOUGALL et al. (1969), P SEARLE (1961). Direct age construints on these samples are nut available. Based on similarity of genchemical composition (PATERSON, 1992). these samples we assumed to have a similar age to Pukeonake cinder cone which is consttuined by tephrachronology to be less then 19.8 ka (TOPPING (1973) us listed by GRAHAM and HACKETT (1987)). Age is likely to be overestimated as K-Ar dating of the Auckland Province bus&s gives ages which ere systematically too high: possibly because of the incorporation of crustal derived rediogenic argon (MCDOUGALL et al., l%9). No direct age constmhus available. This is a maximum geological age based on lack of erosional features (SEARLE, I %l).