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Hobbs Coast Cenozoic volcanism: Implications for the West Antarctic rift system
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CHEMICAL GEOLOGY fMC~DING
ELSEVIER
ISOTOPE GEOSCIENCE
Chemical Geology 139 (1997) 223-248
Hobbs Coast Cenozoic volcanism: Implications for the West Antarctic rift system Stanley R. Hart a,*, Jerzy Blusztajn a, Wesley E. LeMasurier b, David C. Rex c a Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA b Department of Geology, University of Colorado, Denver, CO 80217, USA c Department of Earth Sciences, University of Leeds, Leeds, LS2 9JT, UK
Received 31 December 1996; accepted 8 January 1997
Abstract
Basaltic lavas were erupted from a 40-km-long lineament near the Hobbs Coast of Made Byrd Land, Antarctica, over the period from 11.7 m.y. to 2.3 m.y. ago. The lavas from the southernmost locality, Coleman Nunatak, are virtually constant in major, trace dement and isotopic composition over this entire age span. Their high FeO-low A120 3 character indicates melting of garnet peridotite at about 140 km depth. There is no evidence for the involvement of ancient continental lithosphere or MORB asthenosphere in the magmatism. Isotopically, the lavas show the highest 2°sPb/2°4pb ratios (up to 20.7) of any of the Cenozoic volcanism associated with the West Antarctic rift system (WARS). This HIMU isotopic signature is also clear in the trace element patterns, which closely mimic end-member HIMU basalts from the oceanic islands of Tubuai and Mangaia. From the other localities along the Hobbs Lineament, the earliest volcanism, which is coeval with that at Coleman Nunatak, is of shallower derivation ( ~ 110 km), and isotopically like the oceanic FOZO end-member (Z°tPb/2°4pb ~ 19.5). The trace-dement patterns are similar to those at Coleman, but less enriched in the most incompatible elements by a factor of two. Modeling of the trace element data is consistent with a uniform mantle source composition, depleted in major dements, but hydrous and mildly enriched in the incompatible and LREE. Inversion for the bulk distribution coefficients of the source mantle reveals a spidergram with a marked negative Ti anomaly and marked positive anomalies for K, Sr, Zr and Hf. From this modeling, the extent of melting at Coleman is inferred to be ~ 1.6%, as compared to ~ 3.2% during the earliest volcanism elsewhere on the lineament. With time, the volcanism from these other localities progresses to greater depth, becomes more HIMU in character, and lower in extent of melting (i.e., approaches the character of basalts from the Coleman locality). The FOZO component is prevalent as a mixing end-member in WARS volcanism from numerous other Made Byrd Land (MBL) and Northern Victoria Land (NVL) localities. It is also the main constituent of the three nearby oceanic plumes (Balleny, Scott, Peter I islands). The HIMU component is at best a minor constituent of these oceanic plumes, but is present at several other MBL and NVL localities, as well as in pieces of Zealandia which were adjacent to this coast of Antarctica prior to fragmentation of Gondwana. We propose that this HIMU mantle source was emplaced under Gondwana lithosphere prior to breakup, as a large weak plume head, with little or no accompanying volcanism. This 'fossil-plume' proto-lithosphere is now being sampled during WARS extension. Likely mechanisms for the volcanism relate either to small-scale
* Corresponding author. FAX: 508-457-2175; E-mail: [email protected]. 0009-2541/97/$17.00 ~9 1997 Elsevier Science B.V. All fights reserved. PH S0009-2541 (97)00037-5
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
convection associated with strong basal topography of the lithosphere (such as that recorded by the Hobbs Lineament volcanism), or to emplacement of a new plume, which may in part be driving the extension. © 1997 Elsevier Science B.V. Keywords: volcanism;rifting; plumes; isotopes;trace elements
1. Introduction Cenozoic volcanism in Antarctica is largely confined to west Antarctica and ranges in age from ~ 30 m.y. to recent (Armstrong, 1978; LeMasurier and Rex, 1983; Smellie et al., 1988; LeMasurier and Thomson, 1990), though the preponderance of volcanism is less than 6 - 8 m.y. old. Aside from the subduction-related volcanic rocks of the northern Antarctic Peninsula and Scotia Arc, the remainder are alkaline basalts and related differentiated rocks; tholeiitic-series rocks are not common. While the tectonic setting of the subduction volcanic rocks is clear enough, the alkalic volcanics occur both as isolated 'hotspots' (Balleny Island, Scott Island, Peter I Island) and as an alignment of volcanic provinces along the edges of the West Antarctic rift system. As defined by LeMasurier (1990a), this rift system extends from the Ross Sea to the Bellingshausen Sea; the volcanic provinces along the Transantarctic Mountain front in Victoria Land lie on the south side of the rift, and the volcanic provinces in Marie Byrd Land (MBL), and extending to Jones Mountains, lie on the northern side of the rift (Fig. 1). While the bulk of the mainland volcanism is presumed to be related to rifting and extension (LeMasurier, 1978; Cooper and Davey, 1985; Behrendt and Cooper, 1991), the underlying causes are not well understood. Some researchers subscribe to variants of conventional active or passive rift magmatism models (Fitzgerald et al., 1986; Tessensohn and Wtmer, 1991). Others appeal to lithospheric effects associated with the continental plate overriding a mantle 'hot line' associated with earlier continental break-up (Smith and Livermore, 1991). Then there are a diversity of models involving some connection to mantle plume activity (LeMasurier and Rex, 1989; Behrendt et al., 1991, 1992; Kyle et al., 1992; Hole and LeMasurier, 1994). On a grander scale, this volcanism forms part of a major arcuate belt of alkaline volcanism that extends from southern South America (Stem et al., 1990) via the Antarctic Peninsula, Marie Byrd Land and Victo-
ria Land to New Zealand (Gonzalez-Ferran, 1972; Cooper et al., 1982). We believe that all of this volcanism shows a common geochemical and isotopic 'fingerprint'. While a full understanding of the WARS volcanism cannot be achieved without viewing it in this larger context, regional studies are important for elucidating the temporal and spatial evolution of the volcanism. We report here on detailed chemical studies of one key lineament in Marie Byrd Land, where the HIMU isotopic signature which characterizes this whole Cenozoic volcanic belt is most prominently displayed.
2. Regional setting The basaltic rocks of the Hobbs Coast region are part of a large alkaline basalt-trachyte volcanic province that lies within the West Antarctic rift system (Fig. 1), about 2000 km from the nearest active plate boundary. The province has been volcanically and tectonically active from about 25-30 m.y. to the present. It is centered over a large Neogene tectono-magmatic dome whose growth and concurrent volcanism are believed to be related to the inception of mantle plume activity around 30 m.y. (LeMasurier and Landis, 1996). Extension and block faulting, related to this doming, have produced basin and range topography in the Hobbs Coast area and throughout Marie Byrd Land. Basaltic rocks (basalt, basanite, tephrite, hawaiite) are the most widespread and voluminous of the volcanic rocks, and have been erupted almost continuously throughout Neogene time. They are found as basal sections up to several thousand meters thick, underlying large felsic shield volcanoes, and as isolated flows, hyaloclastite sheets, cinder cones and tuff cones that rest on basement rock (satellite centers), or on the flanks of the shield volcanoes (LeMasufier, 1990b). The basement rock in this region is part of a large Devonian to Cretaceous magmatic arc terrain that
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Fig. 1. Index maps. (At Map of Antarctica showing the position of map B (square), the position of the Marie Byrd Land dome (dotted line), and the southern boundary of the West Antarctic rift system (dashed line). Note that, for maps of this scale, the Greenwich meridian is oriented up by internatiolml convention, but that for maps (B) and (C), north is up. Abbreviations are: AP, Antarctic Peninsula; B, Balleny Islands; P, Peter I Island; S, Scott Island. (B) Made Byrd Land volcanic province, showing major central volcanoes and their inferred potential for future activity (circles). 'Satellite centers' are relatively small volcanic centers that consist of cinder cones, tuff cones, flows and hyaloelasfite sequences of predominantly basaltic compositions. (C) Reconnaissance geologic map of the area described in this study, modified from LeMasurier (1990c).
extends along the coast of West Antarctica from the Antarctic Peninsula to the Ross Sea, representing a prolonged interval of subduction prior to the 85 m.y. breakup of this sector of Gondwanaland. In the Hobbs Coast and its environs the basement is dominated by a 9 5 - 1 0 2 m.y. suite of A-type granitoids (Weaver et al., 1994), and less common talc-alkaline granitoids with U - P b ages of 110-322 m.y. (Mukasa,
1995). These rocks intrude early Paleozoic metaclastic rocks. Granitoids from this region have Nd model ages of 900-1200 m.y., suggesting the presence of a Proterozoic basement that is not exposed anywhere in Made Byrd Land (Pankhurst et al., 1995). Thus, the work done so far suggests a relatively juvenile lithosphere for this region. Previous trace element and Sr and Nd isotopic studies of the Cenozoic
Coleman Nunatak Coleman Nunatak Coleman Nunatak Coleman Nunatak
Coleman Nunatak
Coleman Nunatak
Coleman Nunatak
Cousins Rock
Patton Bluff Shibuya Peak
Kouperov Peak
Holmes Bluff
Holmes Bluff
35K 35E 35M 35J
46B
46,I
46D
36B
37A 38A
47A/B
48B
50A
1.33
1.644
1.468
1.487
1.482
1.458
%K
2130 2311
3038 3040 3020 3022 3461 3480 3021 3.23 3366 3376
Ar Nr.
0.04262 0.04219
0.01593 0.01758 0.01714 0.01915 0.0654 0.0700 0.04550 0.04482 0.0342 0.0336
Vol. 4°Ar rad.
39.2 28.2
15.2 8.0 11.3 12.8 6.9 7.0 11.0 11.3 41.6 37.6
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rad.
6.27 + 0.25
8.17 + 0.33
8.18 4- 0.33
9.97 + 0.40 4.66 ± 0.50
5.3 ± 0.2
7.90 ± 0.36
11.7 ± 1.0
3.174-0.13
2.63 + 0.11 3.19 ± 0.33 2.34 5: 0.11 2.95 ± 0.12
Age (M a)
Basalt flow on WAES (LeMasurier and Rex, 1983). Basalt flow in valley cut into WAES (LeMasurier and Rex, 1983).
Basalt flow on granite. (LeMasurier and Rex, 1983). Basalt clast in hyaloclastite, middle of exposed section (LeMasurier and Rex, 1983). Basalt flow on granite.
Dike in eroded cinder cone.
South cliff, basalt clast from base of hyaloclastite section.
Flow rock, base of south cliff.
~ Same horizon as 35E and 35M, different locality.
Top of section, tuff cone (LeMasurier and Rex, 1983). Top hyaloclastite underneath tuff cone (LeMasurier and Rex, 1983). Top hyaloclastite underneath tuff cone (LeMasurier and Rex, 1983). ~ Same horizon as 35E and 35M, different basalt clast.
Comments
Constants: ~8 = 4.962 × 10- I 0 y r - t ; E = 0.581 × 10- l o y r - t ; 40K = 0.01167 atom%; 4oAr/36At = 288.5. AH dated materials are whole rock. Analyses were done at the University of Leeds, following procedures described by LeMasurier and Rex (1989).
Locality
Sample Nr.
Table 1 K - A t data for Hobbs Coast Nunataks
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
basalts have found no evidence for contamination by this lithosphere (Futa and LeMasurier, 1983; Hole and LeMasurier, 1994).
3. Field and age relationships The samples described were collected from a north-south line of nunataks, extending from Coleman Nunatak, in the south, to Holmes Bluff, that project above the continental ice sheet just inland of the Hobbs Coast (Fig. 1C; Table 1). We will call this the Hobbs Lineament. Coleman Nunatak is flattopped, about 1.3 krn long, and composed of subhorizontaUy bedded basaltic hyaloclastite breccias with a total exposed thic~mess of ~ 150 m. This section is, in turn, overlain by a low tuff cone at the north end. Glacial striations on the highest points indicate that the whole nunarak has been overridden by ice. Locality 35 is from the vicinity of the tuff cone, and represents the highest stratigraphic levels at Coleman Nunatak. Locality 46 is at the south end of the Nunatak, at the top of a ~ 100 m cliff with limited access to the base. The cross-section exposed here contains multiple stratigraphic discordances, interbedded flow rock, and foreset-bedded hyaloclastite with redeposited tuff clasts. There is, therefore, some uncertainty about the stratigraphic relationships of the samples because of the possibility of redeposition of clasts. However, the ages are generally concordant with stratigraphic position; 46D and 46J are from the bottom of the section, 35K from the top, and the remainder in between (Table 1). All samples are from lava clasts or entire flows, interbedded in the hyaloclastite. The K - A r ages were obtained from glass--free samples. The ages suggest that this section formed by multiple, small-scale eruptions over a pe~riod of ~ 5 - 9 m.y., in marked contrast to the relatively large magma production rate at a shield volcano like Mount Murphy (Fig. 1B), where it appears that ~ 500 km 3 of basalt was erupted in < 1 m.y. (LeMasurier, 1990c). Field relationships at the remaining localities are described in Table 1 and Fig. 1C. At Patton Bluff and Holmes Bluff, basalt flows rest on granite that has been truncated by the very low relief West Antarctic erosion surface (WAES). Displacement of the WAES to different levels in adjacent nunataks
227
provides evidence for the effects of Neogene block faulting (LeMasurier and Landis, 1996). At Kouperov Peak and Holmes Bluff (sample 50A) basalt flows lie on the granite floor of valleys cut into the WAES. Samples 48B and 50A therefore satisfy a stratigraphic test of their ages, in that flow 50A, with an age of 6.27 m.y., fills a valley cut through an older flow (48B, 8.17 m.y.). To summarize, volcanism along the Hobbs Lineament started 11.7 m.y. ago at Coleman Nunatak; volcanism continued at Coleman until 2.34 m.y. ago. Coleman volcanism was joined at 9.97 m.y. at nearby Patton Bluff, then spread to the north and central parts of the lineament at 8.17 m.y. (Holmes Bluff), and 8.21 m.y. (Kouperov Peak).
4. Sample and analytical details The number associated with each sample is a locality number (see Table 1 and Fig. 1); the letters refer to different hand-specimens from a given locality. Megascopically, most of the samples are finegrained to aphanitic; the locality 46 samples are somewhat coarser-grained. Most are dense to only slightly vesicular; 35J and 35K are somewhat more vesicular. Phenocrysts of olivine and clinopyroxene are both usually present in small amounts (0-5% each); 35E and 35K are slightly richer in clinopyroxene (5-10%); plagioclase is present as a trace phenocryst in 37A, and 38A, and at about 5% in 47A. Eight of the samples are classified as basanites, using the LeBas et al. (1986) nomenclature (see Table 2). There are two hawaiites (but close to the basanite boundary), two alkali basalts and only one differentiated rock, a mugearite. All but one of the samples are nepheline-normative (48B has just a trace of nepheline; 37A shows 5% hypersthene in the norm). Except for the mugearite, the Mg numbers range from 49 to 61. All of the samples appear fresh and unweathered, though several contained thin oxidative finds which were removed before powdering. Though we did not do volatile analyses, various trace elements which are typically sensitive to weathering and alteration show no hint of such effects. For example, only one sample (38A) shows an anomalous R b / C s ratio (41), whereas the other twelve show values which cluster very tightly around a value of 112 ( + 8 % ,
0.702843 0.512873 20.637 15.715 39.732 43.55 3.95 15.00 12.85 0.221 7.02 10.21 4.55 1.74 0.90 6.29 0.520 32%
531 37.6 0.36 1045 294 98.4 36.5 17 264 52 98 20 24 94 64.3 123.4
87Sr/S6Sr 143Nd/144Nd 2°6pb/2°4Pb 2°7pb//204pb 2°spb/2°4pb SiO 2 TiO 2 Al203 FeO* MnO MgO CaO Na20 K20 P205 Na20+K20 Mg number Olivine, Mg number 73
Ba (ppm) Rb Cs Sr Zr Nb Y Sc V Ni Cr Ga Cn Zn La Ce
Rock type: Age(m.y.):
46D Coleman Nunatak Basanite 7.9(0.36)
Sample Nr.: Location:
533 36.7 0.35 1047 297 97.8 35.7 27 255 51 95 21 24 95 64.8 124.1
0.702809 0.512919 20.582 15.688 39.641 43.59 3.85 15.08 13.02 0.222 6.97 10.08 4.52 1.77 0.89 6.29 0.515 33%
46I Coleman Nunatak Basanite 11.7(1.0)
309 64 103 19 44 83 64.5 122.6
533 38.2 0.34 1076 301 97.3 36.4
0.702815 0.512883 20.619 15.727 39.733 44.32 3.92 14.95 13.06 0.12 7.15 10.08 3.84 1.75 0.81 5.59 0.520 33%
4613 Coleman Nunatak Basanite 3.17(0.13)
516 36.7 0.33 1028 293 95.0 35.8 29 262 53 95 20 29 99 64.4 122.6
0.702858 0.512908 20.652 15.746 39.778 43.37 4.00 14.91 13.12 0.218 7.23 10.20 4.35 1.71 0.89 6.06 0.522 33%
35E Coleman Nunatak Basanite 3.19(0.33)
508 36.5 0.36 1046 293 96.6 37.0 28 260 55 96 16 25 97 63.4 120.7
0.702830 0.512894 20.629 15.708 39.692 43.42 4.00 15.07 12.92 0.222 7.25 10.24 4.26 1.72 0.90 5.98 0.526 32%
35I Coleman Nunatak Basanite 2.95(0.12)
Table 2 Chemical data for volcanic rocks from the Hobbs lineament, Antarctica
313 67 102 18 43 91 65.4 124.2
529 39.1 0.37 1071 296 103.0 37.3
0.702807 0.512908 20.689 15.769 39.904 43.52 3.79 15.34 13.25 0.27 6.89 9.86 4.47 1.73 0.88 6.20 0.507 35%
35M Coleman Nunatak Basanite 2.34(0.11)
517 37.1 0.36 1040 295 96.5 36.9 19 256 51 94 15 22 97 64.1 123.3
0.702833 0.512900 20.641 15.711 39.729 43.77 3.91 15.05 12.92 0.219 6.96 10.09 4.41 1.77 0.89 6.18 0.516 33%
35K Coleman Nunatak Basanite 2.63(0.11)
117 12 12 18 29 100 68.2 128.3
489 35.7 0.30 953 344 94.8 35.2
0.702883 0.512892 20.450 15.689 39.614 50.17 2.31 16.62 12.50 0.21 3.30 6.44 5.30 2.01 1.13 7.31 0.343 50%
3613 Cousins Rock Mugearite 5.3(0.2)
237 192 329 19 63 83 35.5 68.7
291 22.5 0.19 652 222 52.2 29.6
0.702758 0.512961 19.502 15.643 38.882 47.01 2.69 15.20 11.73 0.08 9:42 9.43 2.58 1.28 0.58 3.86 0.614 19%
37A Patton Bluff Basalt 9.97(0.40)
235 99 211 18 66 98 46~9 90.4
375 24.8 0.61 829 234 70.4 29.7
0.702831 0.512911 20.195 15.639 39.242 46.52 2.84 15.14 12.28 0.18 7.25 9.63 3.92 1.44 0.79 5.36 0.539 29%
38A Shibuya Peak Hawalite 4.66(0.50)
232 51 82 20 47 90 43.9 86.7
368 23.7 0.21 762 252 62.6 33.4
0.702700 0.512952 19.863 15.627 39.004 46.74 2.78 17.17 11.56 0.22 5.64 9.26 4.27 1.61 0.77 5.88 0.491 32%
47A Kouperov Peak Basanite 8.21(0.33)
224 154 285 20 63 82 31.1 60.2
257 14.7 0.11 576 203 46.1 27.6
0.702535 0.513008 19.566 15.631 38.857 48.85 2.20 15.74 10.89 0.10 7.84 9.58 3.43 0.93 0.44 4.36 0.588 21%
48B Holmes Bluff Basalt 8.17(0.33)
227 113 220 17 57 84 48.2 87.5
463 27.6 0.24 900 215 70.0 28.6
0.702716 0.512945 20.032 15.688 39.339 46.12 2.45 15.83 12.18 0.21 7.65 9.41 4.17 1.14 0.85 5.31 0.554 27%
50A Holmes Bluff Hawalite 6.27(0.25)
::t
tx~ t,o I t~ 4~ oo
ga
5.40 88.2 14.13 8.25 384 27.2 104.4 2.99 16.34 1.68 44.9 3.81 5.49 0.575 62.3
Ba/Nb Ba/Th Ba/Rb Ba/La K/Rb K/Ba Rb/Cs Zr/Nb Nb/Th Sm/Hf Ce/Pb Th/U La/Sm U/Pb Nb/U
5.45 89.9 14.54 8.23 401 27.6 104.7 3.04 16.49 1.70 44.9 3.68 5.48 0.583 60.7
15.18 58.3 11.83 3.54 9.24 1.35 6.87 1.25 3.14 0.44 2.62 0.40 5.76 5.93 6.94 1.61 2.76 5.48 88.0 13.97 8.26 381 27.3 112.2 3.09 16.06 1.65 50.0 3.76 5.56 0.657 60.5
14.99 57.3 11.60 3.51 9.13 1.34 6.78 1.24 3.17 0.44 2.59 0.39 5.89 6.06 7.04 1.61 2.45 5.43 90.1 14.08 8.01 387 27.5 111.1 3.08 16.58 1.71 47.7 3.60 5.41 0.619 59.8
15.17 57.7 11.90 3.57 9.50 1.36 7.08 1.30 3.24 0.44 2.63 0.41 5.59 5.57 6.97 1.59 2.57 5.36 87.6 13.97 8.02 391 28.1 101.5 3.03 16.65 1.70 44.0 3.74 5.39 0.566 62.3
14.86 58.1 11.76 3.57 9.62 1.34 7.10 1.27 3.16 0.44 2.69 0.42 5.65 5.80 6.91 1.55 2.74 5.14 87.7 13.53 8.09 368 27.2 105.7 2.87 17.08 1.75 54.9 3.61 5.54 0.739 61.7
14.48 57.9 11.80 3.76 9.65 1.38 7.40 1.41 3.50 0.44 2.66 0.40 5.62 6.03 6.76 1.67 2.26 5.36 86.5 13.92 8.07 396 28.4 103.2 3.06 16.13 1.68 59.6 3.71 5.65 0.778 59.9
15.26 58.3 11.34 3.49 9.19 1.34 6.78 1.23 3.15 0.45 2.64 0.39 5.66 5.98 6.77 1.61 2.07 5.16 73.5 13.71 7.17 468 61.2 118.9 3.63 14.26 1.38 51.1 3.91 6.24 0.677 55.8
15.52 57.7 10.92 3.27 8.50 1.23 6.55 1.19 3.10 0.45 2.75 0.43 5.86 6.65 7.90 1.70 2.51
Following the age, number in parentheses is the uncertainty, in m.y. 87Sr/86Sr and 143Nd/I'~Nd normalized to 0.71024 (SRM987) and 0.511847 (La Jolla). Majors recalculated to 100% volatile-free, all Fe as FeO, Mg number calculated with FeO = 0.9 FeO *. Olivine, Mg number 73 is the percent of equilibrium olivine added incrementally to bring Mg number to 73. Zr, So, V, Ni, Cr, Ga, Cu, Zn by XRF; all other traces by I C P / M S .
15.09 59.0 11.71 3.63 9.24 1.33 7.06 1.27 3.17 0.44 2.65 0.41 5.47 6.02 6.99 1.58 2.75
Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Ta Th Hf U Pb 5.57 86.4 12.95 8.21 473 36.5 118.3 4.25 15.49 1.42 26.4 3.79 4.80 0.342 58.7
8.52 33.8 7.38 2.26 6.45 0.96 5.38 1.02 2.63 0.38 2.31 0.37 3.13 3.37 5.21 0.89 2.60 5.33 88.2 15.12 8.00 481 31.8 40.7 3.32 16.56 1.66 30.0 3.54 5.29 0.399 58.7
10.61 42.8 8.86 2.88 7.74 1.09 5.94 1.11 2.91 0.36 2.12 . 0.30 3.87 4.25 5.34 1.20 3.01
10.09 42.1 8.90 2.96 7.37 1.04 5.93 1.10 2.70 0.35 Z07 0.31 3.41 3.96 4.69 0.99 2.46 6.61 116.9 16.78 9.60 342 20.4 115.0 3.07 17.68 1.90 35.6 4.00 5.42 0.402 70.7
7.15 28.5 6.49 2.15 6.02 0.90 5.2i 1.02 2.74 0.36 2.17 0.32 2.52 2.80 4.57 0.78 2.15 5.57 91.8 17.48 8.27 523 29.9 133.6 4.40 16.46 1.42 28.0 3.59 4.79 0.363 59.1
10.56 41.8 8.91 2.99 7.81 1.13 6.50 1.22 3.18 0.42 2.50 0.37 3.57 4.09 5.69 1.13 3.58 5.88 90.0 15.55 8.38 566 36.4 112.7 4.03 15.30 1.57 24.2 3.62 4.93 0.316 55.4
;zt
~D
oo
I
t,~ t~
',O
g~
230
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
ltr). This compares well with the canonical value of 85-95 found for fresh ocean island basalts (Hofmann and White, 1983; Morris and Hart, 1983). The T h / U ratio of all the samples is constant at 3.72 ( + 4%, l tr); Th is quite robust to alteration, whereas U can be significantly 'mobile'. Sr, Nd and Pb isotopic analyses were carried out with conventional techniques (Hauri and Hart, 1993), using sample powders which had been leached in warm 6 N HC1 for 1 h. Sr and Nd data carry precisions of _ 0.003%, and are reported relative to 0.71024 (NBS 987) and 0.511847 (La Jolla). Pb is corrected ratio by ratio for fractionation, using NBS 981 and the values given by Todt et al. (1996). Major and trace elements were determined on unleached powders prepared in an agate shatterbox. Major elements and some trace elements (Ba, Sr, Cr, Ni, Nb, Ga, Y, Rb, Zr, Cu, Zn, V) were determined by XRF; the other trace elements (plus Ba, Sr, Rb, Nb, Y) were determined by ICP/MS at Washington State University. Much of the XRF data has been previously published (LeMasurier, 1990c). The chemical and isotopic data are given in Table 2, with the samples arranged in order of locality, from south to north along the Hobbs Lineament.
50 4948" 47" © 4645" ~2
t
@A
43 £ 42 • " " ' I . . . . 5 10
[]
ib I ....
15
20
11
L) O&
m l[] ! ........
5
5
lO
5
10
lJ5
20
15
20
5. C o n d i t i o n s o f m e l t i n g a n d d i f f e r e n t i a t i o n
The major element data are shown in Figs. 2 and 3 in selected variation diagrams, both as raw data and as olivine-corrected data. Noteworthy in the raw data is the very tight clustering for samples from the two localities on Coleman Nunatak (46, 35; squares), despite their considerable range in age from 11.7 m.y. to 2.3 m.y. Typically, this Coleman group will lie outside (SiO 2) or at the end (Na20) of the data fields for the other localities. The other localities show more dispersed data arrays, with one extreme (low Na20) being the hypersthene-normative basalt from Patton Bluff (37A). The raw data for the mugearite (36B) is not shown, as it lies well away from the basalt data and would expand the plots excessively (it has only 3.3% MgO, and 50.2% SiO 2, for example). Comparison of the raw basalt data is hindered by the significant range in Mg number (49 to 61). For this reason, we have attempted a first-order fractiona-
2
MgO Fig. 2. Correlation of SiO 2, CaO and N a 2 0 with MgO, Hobbs Lineament volcanics. The raw data (Table 2) is enclosed in fields;
the other symbolsare for data which has been correctedto Mg 73 by incremental addition of equilibrium olivine (see text). The curves are the near-solidus garnet lherzolite melt compositions from the Herzberg and Zhang (1996) algorithms, labeled with pressure in GPa. Herzberg and Zhang (1996) do not give an algorithm for SiOz, but state that it is fairly constant at 45-47% over the pressure range 2.5-10 GPa. The same symbols will be maintained in all figures. Squares are Coleman series 35 and 46 samples, undifferentiated; X -- 38A; Y = 47A; dots = 50A and 48B; triangle = 37A. The mugearite is not plotted.
tion correction, on the assumption that only olivine has been a fractionating phase. Clearly this is a simplification, as clinopyroxene is commonly present as a phenocryst phase, along with olivine. We modeled possible Cpx fractionation effects by using V -
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
18 16-. ~
I18" 16"
12-
~..,~12-~ 5
10-
10"
86
8" ....
5
231
I' 10
•
6 15
MgO
20
.... 10
I .... I .... I .... 11 12 13 14 FeO
Fig. 3. Correlationof AI:~O3, MgO and FeO, HobbsLineamentvolcanics.The 'as is' data (Table 2) is enclosed in fields; the other symbols are for data which has been corrected to Mg 73 by incremental addition of equilibrium olivine (see text). The curves are the near-solidus garnet lherzolite melt compositionsfrom the Herzbergand Zhang (1996) algorithms, labeled with pressure in Gpa. Squares are Coleman series 35 and 46 samples, undifferentiated; X = 38A; Y= 47A; dots = 50A and 48B; triangle= 37A. The mugeariteis not plotted. C a O / N a 2 0 relationships. C a O / N a 2 0 is virtually independent of pressure of melting (Herzberg and Zhang, 1996), and of olivine fractionation, but is very sensitive to Cpx fractionation. V, with a C p x / m e l t partition coefficient of ~ 4 (Hart and Dunn, 1993; Jenner et al., 1993), was modeled with fractional crystallization. The data do not lie along a Cpx fractionation tJ~jectory, and there is at most a 3% Cpx fractionation effect between the average Coleman samples and the other basalts from the Hobbs Lineament. While we could in principal normalize the data for Cpx control using V contents, the effects are small enough to overlook. Normalization for olivine control was achieved by incrementally adding olivine back to each composition, while maintaiLning a constant F e / M g g o = 0.30, until the Mg number was 73. This is equivalent to producing melt,; in equilibrium with a mantle olivine of Fo90 composition. This is a reasonable mantle composition; other choices could be made, but the key point is to normalize all the samples to the same Mg number. The amount of olivine required to reach Mg 73 ranged from 32 to 35% for the Coleman locality, and from 19% (37A) to 32% (47A) for the other localities (the nominal value for the mugearite was 50%, though this data is not shown in Figs. 2 aud 3). Both the olivine correction and the Cpx fractionation issue were evaluated further by carrying out liquid-line-of-descent calculations using the MELTS program (Ghiorso and Sack, 1995), starting with a 4 GPa melt from Herzberg and
Zhang (1996), cooled with variable amounts of water and K 2 0 at pressures in the 0.1-0.5 GPa range. These calculations confirm the simple approach we have taken above. In most cases, the olivine correction decreases the scatter in the data, typically generating fight curvilinear arrays or compact clusters. As we will argue later, the relative homogeneity of isotopic and trace element ratios in all of these samples argues strongly for a homogeneous mantle source composition. The trends in the olivine-corrected data are thus unlikely to be related to differences in source composition. Based on the high abundances of incompatible trace elements, the extent of melting must also be very small for these basalts, so the compositional trends seen in Figs. 2 and 3 are not related to wide ranges in the degree of melting. It is now well accepted for MORB melting that a principal control on elements such as FeO and SiO 2 is pressure of melting (e.g., Kinzler and Grove, 1992; Hirose and Kushiro, 1993). The trends shown in Figs. 2 and 3 are consistent with such a pressure-dependent melting model, though the actual pressures are likely to be much higher than for the MORB case; this is based on the need for a large residual garnet component during melting, to be discussed in Section 7. Recently, Herzberg and Zhang (1996) have reported melting experiments on a peridotite with near-primitive composition over a large range of pressures in the garnet stability field. Using their algorithms for near-solidus melt compositions, we
232
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
have plotted on Figs. 2 and 3 the compositional trajectory for these melts, in the 2-5 GPa pressure range. While our data arrays are displaced from the Herzberg and Zhang curves, the general parallelism is quite stalking--see especially the AI203-MgOFeO data in Fig. 3. In part, the displacement of the arrays may be due to the fact that their source is a fertile peridotite, whereas we show below that the Hobbs Lineament mantle source is most probably fairly depleted in elements such as CaO and AI203. We also show that the Hobbs mantle source probably contains hydrous phases, in contrast to the anhydrous experiments of Herzberg and Zhang. Considering this source difference and the difficult nature of the experiments, we feel that the general correspondence with our data is adequate, and supports a model where the principal control on our compositions is pressure (i.e., depth of melting). For the Coleman locality, the pressure estimates derived from the Herzberg-Zhang curves are ~ 4.7 GPa (FeO), ~ 4.1 GPa (MgO), > 3.6 GPa (A1203), < 6.1 GPa (CaO), and ~ 4.1 GPa (Na20). For the lowest FeO sample from the other localities (48B), the comparative pressures would be 3.6, 3.3, 3.0, 5.5 and 3.5 GPa; the pressure range exhibited by our samples is thus about 1 GPa (based on the MgO and FeO parameters). Taken at face value, these considerations suggest that the Coleman locality samples represent melt-equilibration at about 140 km depth, compared to a depth of about 110 km for sample 48B, Holmes Bluff. It might also be noted that, whereas the Coleman samples appear to have maintained their 140 km melting depth over a significant time span, melting at Holmes Bluff appears to migrate deeper with time, from ~ 100 km at 8.2 m.y. to ~ 115 km at 6.3 m.y. While this may be stretching the use of the H - Z algorithms in absolute values, we feel that the sense of the trends is certainly correct (higher FeO is higher pressure, etc.), and that the differences in indicated pressure are robust estimates (comparable to a 30 km depth difference between Coleman and Holmes). For ease of discussion we will use the 110 km and 140 km depth estimates, though we recognize that the absolute values must be tentative due to the different source compositions in their experiment and in the Hobbs case. We will make use of these melting depths later, in talking about models for WARS rifting.
Finally, to evaluate the conditions under which these magmas fractionated in the crust, we carded out MELTS modeling (Ghiorso and Sack, 1995) on samples from several of the localities, cooling them at various pressures and H 2 0 contents to look for the conditions of multiple-saturation (olivine + cpx, or in the case of sample 47A, ol + cpx + plagioclase). In several cases (37A, 50A), the melts had to be virtually dry to achieve multiple saturation at crustal pressures. For dry melts, multiple saturation was achieved at 0.41 GPa (35M), 0.7 GPa (37A), 0.28 GPa (47A), 0.37 GPa (48B) and 0.71 GPa (50A). These all represent mid-to-deep crustal levels; apparently little or no differentiation occurred in shallow magma chambers, perhaps suggesting the lack of any magma chamber for the small eruptions which make up the Hobbs Lineament.
6. Isotopic signatures One of the unusual aspects of the heavy isotope data, shown in Fig. 4, is the large range in 2°6pb//2°4pb isotopic ratios (19.50-20.69) coupled with a very restricted range in 87Sr//86Sr (0.702540.70288) and 143Nd/144Nd (0.51287-0.51301). Much of the data fall in areas of isotopic space not previously populated with data; at first glance the isotope arrays appear to be aligned between HIMU and MORB mantles. The HIMU mantle component (or at least a mantle component with fairly radiogenic Pb) is quite ubiquitous in Antarctic Cenozoic volcanism, first identified in the Balleny Plume (Hart, 1988; Lanyon et al., 1993), and also noted in Marie Byrd Land (Futa and LeMasurier, 1983), Northern Victoria Land (Hart and Kyle, 1994; Rocholl et al., 1995) and Peter I Island (Hart et al., 1995). The HIMU signature is found as well in Gondwanan pieces which were once in proximity to the Pacific coast of Antarctica (e.g., Tasmania, Ewart et al., 1988; Lanyon et al., 1993; New Zealand, Barreiro and Cooper, 1987; S.R. Hart et al., unpubl, data). To date, the 2°6pb/2°4pb value of 20.69 reported here for Coleman Nunatak is the highest reported for any of these 'HIMU' localities (and was one of the motivations for this detailed study of the Coleman area). A possible HIMU-MORB mixing explanation for
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
40.5
233
St.HELI~IA
IA
j;
40.0 39.5 R
39.0
E
U
N
I
O
N
~
~
oO
38.5 38.0 ~
37.5
P/A RIDGE
37A 48B
47A 50A 38A
36B .
18.0
18.5
19,0
19.5
20.0
.
.
20.5
.7045
.
21.0
IB
.7040,
JONES~rms~
%
PAGALU~ t ~ PRINCEEDWARD
\ ~\
~
CANARIES CAMEROONS
o0 .7030
.7025.
~
~
~C~SXON BALLENY/$COTT
.7020
.... 18.0
P/ARIDGE , .... , .... 18.5 19.0
37A 48B 47A 50A 38A 36B , .... , .... , .... 19.5 20.0 2,0.5
2,1.0
206/204 Pb Fig. 4. 20spb/204pb_ s7Sr/86 S r - 206Pb/204pb isotope correlation plots for Hobbs Lineament volcanics, in comparison with other Antarctic localities (Balleny/Scott, Peter I, and Jones Mountains), and a nearby spreading ridge ( P / A = Pacific Antarctic). Selected oceanic basalt data are also shown for comparison (both Atlantic and Pacific N-MOP, B, and OIB). Data sources are as given in Halliday et al. (1992) and Hart et al. (1992, 1995). Squares are Coleman series 35 and 46 samples, undifferentiated; X = 38A; Y = 47A; dots = 50A and 48B; triangle = 37A; Z = 36B mugearite; sample numbers appear at bottom of each panel, directly underneath symbol.
the Hobbs array can be tested in three-isotope space, using the mantle tetrahedron plot of Hart et al. (1992); see Fig. 5A. Here again the Hobbs array projects neatly to MORB, and in particular to the MORB from the nearest segment of the Pacific/Antarctic Ridge (the projection is excellent from other viewing angles as well). This viewpoint also shows that the ,'u'ray does not exactly back-project to HIMU, but to a location slightly up the HIMU-EM 1 edge.
When compared with the data arrays for Scott/Balleny, Peter I, Jones Mountains, and Mt. Early/Sheridan Bluff, all show a tendency to converge in the south-central part of the tetrahedron, i.e., the other arrays are clearly not mixing toward MORB. This is the same characteristic shown for the ocean island basalt (OIB) data set as a whole (Hart et al., 1992), with the convergence zone labeled FOZO (Focus Zone). To examine this phenomenon more closely, we show in panel B of Fig. 5 an enlargement
234
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
Nd•6/4 ~ ~ / f r EM1
Early/Sherid~
P/A Ridge Hobbs
DMM I
HUDSON
HIMU
SIPLE
~ID
of the Hobbs array portion of the tetrahedron, with the addition of some isotopic data from other regions of Marie Byrd Land, for cases where we have more than one sample from a given locality (see Table 3 for the isotope data on these samples). There is a remarkable fan-shaped nature to these arrays, with a convergence or focus just at the left end of the Hobbs array. This suggests that the Hobbs array is not mixing to MORB, but to a more central component common to many other MBL arrays (and very close to the FOZO component of Hart et al., 1992; Hauri et al., 1994a). We note that there is no evidence from nearby ridges or ocean crust for this component, only MORB with normal 2°6pb/2°4pb ratios ( < 19.0) have been reported (Ferguson and Klein, 1993; Pyle et al., 1995). This FOZO component will figure prominently in our WARS model, to be discussed below.
FOSDICK
7. Trace element patterns
Fig. 5. Three-dimensional S7Sr/S6Sr, 143Nd/144Nd and 2°6pb/2°4pb isotope correlation plots. The end-member mantle compositions of Zindler and Hart (1986) define a tetrahedron (Hart et al., 1992), and the data are projected onto the DMMEM1-HIMU base plane of this tetrahedron. (A) View of whole mantle tetrahedron, showing the Hobbs Lineament data (Table 2) compared to arrays from Jones Mountains and Peter I Island (Hart et al., 1995), Scott/Balleny islands (Hart, 1988), Mount Early/Sheridan Bluff (S.R. Hart et al., unpubl, data) and MORB from the Pacific-Antarctic Ridge (P/A, Ferguson and Klein, 1993). This MORB locality is only 650 km northeast of Scott Island. Projections of the three orthogonal isotope axes are shown for reference; end-member compositions are as in Hart et al. (1992), except FOZO has been moved to slightly lower 143Nd/144Nd. (B) Enlargement of the south-central region of the tetrahedron comparing the Hobbs array with several other localities from Made Byrd Land (Table 3). The arrays converge on the FOZO region of Hart et al. (1992); all except Jones Mountains lie very close to the DMM-EM1-HIMU base plane (the Jones Mountains trend, panel A, intersects the base plane at FOZO).
The trace element data are shown in Fig. 6 as primitive-mantle-normalized spidergrams. Because all of the Coleman 35 and 46 locality samples are so similar, we show the average of all seven; for compadson, we show the lowest spidergram for the Hobbs group, 48B, Holmes Bluff. All the other samples lie intermediate to these two spidergrams. All show the typical negative K and Pb anomalies of OIB; these anomalies, coupled with the marked depletions in Cs and Rb, argue against any significant role for crustal contamination of the lavas (as of course do the very low 87Sr/86Sr isotope ratios). Hole and LeMasurier (1994) also provided compelling arguments against significant crustal or mantie lithospheric components in alkaline volcanics from Made Byrd Land and the Antarctic Peninsula. The principal difference between the bounding spidergrams in panel A of Fig. 6 is the more pronounced Z r - H f dip in the Coleman average, and the larger Pb anomaly for Coleman, suggestive of a 'buffering' of Pb in the Hobbs volcanies by a Pbcompatible residual phase. The overall nature of the spidergrams is consistent with melting to variable degrees of an approximately homogeneous source. As noted before, the Sr and Nd isotopes support this notion, in the sense that only small R b / S r or S m / N d variations are needed to
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
235
Table 3 Isotopic data for selected Marie Byrd Land localities Sample Nr.
Locality
87Sr/86 Sr
E-27-34c 66-D-91 66-D-30 W83-5C W83-1 62 85-32B 24-1-A 9-1-C 20-2B
Scott Island Fosdick Mountains Fosdick Mountains Mt, Siple Mt, Siple Mt. Mmphy Mt. Mmphy Hudson Mountains Hudson Mountains Hudson Mountains
143Nd / 144Nd
0.512989 0.512859 0.512948 0.512847 0.512792 0.512893 0.512839 0.512980 0.512859 0.512928
0.702849 0.703338 0.703106 0.703487 0.703723 0.702917 0.703165 0.702904 0.703366 0.703026
1000
~B, HOLMES•
,Iktlll
]
CsRbBa Th UNbTa K l.a Ce PbNd Sr Zr HfSmEuTi GdTbDyHo y Er Yb Lu
1000
'B'
[I
1 CqRb Ba Th U NbTa
il '1:
La Ce Pb Nd Sr Zr FLfSmEu Ti GdTbDyHo Y Er Yb Lu
ELEMENT
Fig. 6. Trace element spidergrams for Hobbs Lineament volcanic rocks. (A) The average of seven Coleman Nunatak samples is compared to the most depleted sample, 48B, Holmes Bluff. (B) The average of the Coleman Nunatak samples compared to the oceanic HIMU end-mem~r. The latter is based on seven basalts from Mangaia and Tubuaii (S.R. Hart and E.H. Hauri, unpubl. data). The Coleman average is as in panel A, but adjusted uniformly down by 21%, which is the amount of olivine which would need to be added to bring the Mg number of Coleman (51.8) up to that observed in the HIMU suite (67.5). Primitive upper mantle (PUM) normalizing factors are from McDonough and Sun (1995); for the refractory elements, these are ~ 10% higher than the values of Hart and Zindler (1986).
2°6pb / 204pb
2°7pb / 204pb
19.638 19.054 19.224 19.669 19.753 19.764 20.010 19.406 19.138 19.333
15.601 15.580 15.607 15.665 15.725 15.667 15.691 15.623 15.603 15.692
2°8pb / 204pb
:
39.197 38.489 38.595 39.251 39.494 39.195 39.463 38.840 38.669 38.984
generate the small differences observed in 87Sr/86Sr and raNd/144Nd. Similarly, while the observed Pb isotope range is large (e.g., 19.5-20.7), the time-integrated differences in U / P b required to produce this in a few billion years is only 25%. We may test the concept of a single mantle source, as well as derive information about the source composition and bulk partition coefficients, using the modeling techniques for partial melting pioneered by Treuil and Joron (1975) and Minster and All~gre (1978). Numerous refinements and successful applications of these techniques have been reported subsequently (Albar~de, 1983; Hofmann and Feigenson, 1983; Albar~de and Tamagnan, 1988; Maalee, 1994; Zou and Zindler, 1996). We will adopt the formuladon given by Sims and DePaolo (1997). For modal batch partial melting: C
1
Co = F + o ( 1 - F )
(1)
where C is the concentration of a trace element in a melt fraction F; CO is the initial source concentration, D is the bulk partition coefficient. When two trace elements A and B are plotted against each other as reciprocals ( x = I / A , y = I / B ) , the resulting correlation is linear, with a slope and intercept defined as: slope
CoA(1 --DB) CoB(1 _DA)
( D s --DA) intercept --- Con ( 1 - DA)
(2)
(3)
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
236
For non-modal batch melting, the array is still linear, but the slope and intercept contain an additional parameter P, which is the bulk partition coefficient of the phases in the proportion in which they are melting:
co (1 slope = Con(1 _ eg)
intercept =
(4)
DB(1 --PA) --Oh(1 - - e a ) CoB(1 - PA)
mildly incompatible element (Gd). For La and Nb, since D << 1, the slope, from Eq. (2) gives the source ratio of N b / L a (here = 1.383); the intercept from Eq. (3) is proportional to the difference in partition coefficient between Nb and La. Because the intercept is positive in this case, DLa > DNb (though the difference is quite small). For Th and Gd, again DTh << 1 but DGa may not be small, so the slope from Eq. (2) gives:
(5)
Co
0.2966 = ~oGd(1 -- DGd) For fractional melting, the equations are not linear, but may nevertheless by treated as linear in limited situations, as we discuss below. Two examples of this approach are illustrated, for the Hobbs data, in Fig. 7; panel A shows two fairly incompatible elements (La and Nb) and panel B shows one highly incompatible element (Th) and one 0.040 : 0.035 ".: 0.030 0.025 ~ 0.020' 0.0150.010' 0"005' ~.,.00181 :~ .00088 0.000
,*,,l.,,i
i,,.,i,.,,
~" ii i f , l , ,
.0(30 .005 .010 .015 . 0 2 0 .025
1/Nb 0.20-' 0.15'
.2966* .0264
0.10' 0.05'
•
and the intercept, from Eq. (3), gives: DGd
0.0756 = - -
Cod
Thus fixing one parameter, say Co~d, allows one to calculate DGd and C~. And from a Th-Nb plot, the additional source concentration of Nb could be determined (and from this, La from Fig. 7A). In principle, for modal melting, by fixing only a single source element concentration, complete spidergrams for both the source concentrations and bulk partition coefficients can be derived, subject only to uncertainties relating to the actual scatter of data in plots such as Fig. 7. From Eqs. (2) and (3), it is clear that the right-hand end of the spidergram is well constrained for both Co and D; the highly incompatible elements are well constrained for CO, but substantial error magnification can occur in the D's, due to the (D B - Da) term. For non-modal melting, some ambiguity arises for the most compatible elements, but by judicious choice of elements the rest of the spidergram is still well constrained. For example, for a highly incompatible element A (where DA and PA << 1), relative to a mildly incompatible element B (D B < 1), Eqs. (4) and (5) become:
•
Co (1 - e . ) 0.00
....
I,,,, I,,,, I,,'" l, ,, ,I, , , , | , , , , I , , , ' m '~'~ ~
.00 .05 .10 .15 .20 .25 .30 .35 .40 .45 1/Th
Fig. 7. Reciprocal L a - N b and G d - T h plots, with York (1966) two-error regression lines plotted through the Hobbs Lineament data. Input uncertainties given as + 5% for each element, with concentrations from Table 2 corrected for olivine fractionation (see text). Output errors for slopes and intercepts are absolute values, quoted at + 2 t r , and are independent of the input errors, to first order (Brooks et al., 1972). The mugearite 36B is not plotted.
slope
CoB
DB intercept = C°a While a global inversion of all the data could be performed, as done by Minster and All~gre (1978), it is difficult to properly weight the 'geologic scatter' on the input errors; the output errors can be signifi-
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
cantly impacted by this aspect (All~gre et al., 1983a,b). Instead, we have iteratively worked our way through the whole data set with the two-error regression treatment of York (1966), which generates error estimates for :slope and intercept which are dominated by the actual data scatter and are relatively insensitive to the input weighting factors (Brooks et al., 1972). This approach capitalizes on the best-constrained data arrays. The results of this process are listed in Table 4 and shown in Fig. 8,~ for both the source concentration (panel A, top, normalized to primitive upper mantle: PUM) and the bulk partition coefficients (panel B, bottom), tks mentioned above, all of the data floats on the choice of one C O value; due to the relative constancy of the heavy REE in peridotltes, we selected Dy (the heavier REEs in our data set, and Yb and Lu in particular, are somewhat decoupied from more incompatible elements and form poor data arrays in plots such as Fig. 7). First, the
237
C s R b k Th U NbTa K La CeNd $r Zr HfSmEuTi GdTbDyHo Y Er Yb Lu 1
Table 4 Source concentrationsand partitioncoefficients Element
C O(ppm)
Co, P U M
20" ( % )
Cs Rb Ba Th U Nb Ta K La Ce
0.00314 0.486 7.08 0.0768 0.0208 1.199 0.0616 258.1 0.864 1.703
0.150 0.gl0 1.(173 0.966 1.023 1.822 1.666 1.075 1.334 1.017
6.9 4.1 3.1 0.4 0.7 1.1 0.8 4.2 1.0 0.5
Pb
.
Nd
0.842
Sr Zr Hf Sm Eu Ti Gd Tb Dy Ho Y Er Yb Lu
.
19.12 8.94 0.1757 0.2163 0.0822 384.9 0.2544 0.0403 0.337 0.0756 1.642 0.255 0.1355 0.01582
.
.
DO
-0.00359 0.00133 0.00167 0.00126 0.00129 0.00032 -0.00143 0.0165 0.00183 0.00245
20" ( % )
26.1 52.9 40.6 6.8 7.4 81.0 8.6 10.1 15.8 6.0
.
0.673
-
0.00332
-
0.960 0.851 0.621 0.533 0.534 0.319 0.,168 0.407 0.5 0.:507 0.1382 0.582 0.'.307 0.234
2.5
0.00816 0.0250 0.0183 0.00879 0.01471 0.00702 0.0206 0.0244 0.0486 0.0644 0.0446 0.0813 0.0526 0.0366
8.2
2.2 2.7 2.3 4.6 2.1 5.8 -
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C.sRbBa Th U NbTa K La C.eNdSr Zr HfSm EuTi C-dTbDyHo YEr Yb Lu ELEMENT F i g . 8. T r a c e e l e m e n t ( p a n e l A ) a n d b u l k p a r t i t i o n c o e f f i c i e n t (panel B) spidergrams for the Hobbs Lineament mantle source
(inversionbased on Dyo = 0.500 PUM; data is in Table 4). PUM normalizingfactors are from McDonoughand Sun (1995). Cs and Ta points are not plotted in panel B because the D values ended up being negative (see Table 4). No line is drawn connectingEr and Yb because of the poor correlationsbetween Yb and Lu and the rest of the spidergram.Dashed lines have been drawn between the various REE to highlight the anomalies of intervening elements. Error bars are not shown but are very small for the elements in panel A; some of the elementsin panel B have sizable uncertainties(Rb, Ba, Nb).
enriched nature of the Hobbs source mantle spidergram (Fig. 8A) is a robust feature independent of the choice of any C o (for example, L a / S m is virtually constant for Dy input choices ranging from 0.5 × PUM to 2 × PUM). Second, as has been known since the classic work of Frey and Green (1974), peridotites with enriched LREE patterns are typically depleted in heavy REE, and also less fertile (in CaO, A I 2 0 3, etc.) than PUM. Utilizing the peridotite
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database of McDonough (1990) (195 spinel and 57 garnet peridotite xenoliths), for samples with L a / S m > 2.5 × PUM ( L a / S m of the source in Fig. 8A is 2.5), only a very few samples (5) have heavy REE > 1 × PUM, and the average is about 0.5 × PUM. For the 68 massif peridotites in the McDonough data base, none have L a / S m > 2 × PUM; the average heavy REE content is ,-, 0.62 × PUM. We believe that our choice of Dy = 0.5 × PUM for Fig. 8 is unlikely to be uncertain by more than ___50%. This uncertainty will translate almost directly into a + 50% variation in the source spidergram and the bulk partition coefficient spidergram. The various anomalies on the spidergrams will not be materially affected. Most of the features of the Fig. 8 spidergrams appear reasonable. First, the source pattern is consistent with a mantle which was once depleted (positive slope on left end), as suggested by the Sr and Nd isotopes, but has subsequently been enriched, to account for the high concentrations in the Ba to La portion of the spidergrams of the erupted basalts (Fig. 6). The bulk partition coefficient spidergram is very steep, as would be consistent with melting in the garnet facies (depths of melting inferred earlier to be ~ 110-140 km). There is a marked negative Ti anomaly, and marked positive Z r - H f anomalies in the partition coefficient spidergram, again, typical signatures of garnet melting (Hauri et al., 1994b; Salters and Longhi, 1997). The absolute values of the heavy REE partition coefficients are quite low, as would be expected for a depleted (harzburgitic) source, with low modal garnet and Cpx abundances. Using the garnet/melt Er partition coefficients of Salters and Longhi (1997) or Hauri et al. (1994b), which have values of 3.5-4.0, no more than ~ 2% garnet can be present during melting. The negative source anomalies at Ti and Y are compensated by negative partition coefficient anomalies, resulting in almost anomaly-free Ti and Y in the melts (see Fig. 6). The positive Zr-Hf anomalies in the source are over-compensated by the large positive partition coefficient anomalies for Zr and Hf, leading to mild negative anomalies in the erupted basalts. The negative K anomaly in the source is exacerbated by a large positive partition coefficient anomaly for K, leading to melts with marked negative K anomalies. A positive source anomaly for Sr is largely offset by
a positive partition coefficient anomaly, leading to melts with only a minor positive Sr anomaly. There are no Eu anomalies in either the source or bulk partition coefficients, and there are no Eu anomalies in the melts (consistent with high-pressure melting). There is a slight bowl-shape to the REE source spidergram (ignoring Lu and Yb) as is frequently found in xenolith suites where an enrichment event is superimposed on a depleted source. The REE slope of the partition coefficient spidergram is not as steep as would be expected for pure garnet control, so it is likely that some Cpx is present as well. Cpx cannot be the dominant phase melting, however, as it carries a negative Zr partitioning anomaly (Hart and Dunn, 1993), which will tend to cancel the positive Zr anomaly in garnet (and in Fig. 8B). The positive K anomaly is indicative of minor amounts of phlogopite or amphibole; the lack of enhanced Rb and Ba partition coefficients argues for amphibole as opposed to phlogopite. It would require 1-2% amphibole to generate the K anomaly in Fig. 8B, for the range of amphibole/melt partition coefficients cited by Green (1994). The origin of the mild positive Sr anomaly in Fig. 8B is less clear, though it may also be attributable to amphibole (Green, 1994). The negative Y anomaly is unexplained, as neither garnet, Cpx nor amphibole appear to fractionate Y from the heavy REE (Hart and Dunn, 1993; Green, 1994; Salters and Longhi, 1997). We ran a series of numerical tests with an aggregated fractional melting model, using the source composition and bulk partition coefficients of Table 4, to see how closely the recovered slopes and intercepts agreed with those derived by assuming a simple batch modal melting model. The data are no longer linear on 1 / A - 1 / B plots, but for our particular data set the degree of curvature is very small and would not be noticeable given the scatter of data on plots such as Fig. 7. A linear regression (of the non-linear data) produced intercept values which were generally very close (within 10%) to those calculated using Eq. (3). The recovered slopes, however, typically underestimated slopes calculated from Eq. (2) by 5-40%. Thus, the bulk partition coefficient spidergram (Fig. 8B) will be virtually unaffected, were the actual melting process for the Hobbs Lineament to be fractional as opposed to batch melting. The general appearance of the source mantle
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
spidergram (Fig. 8A) will also be little changed, though it will be rotated clockwise (about the fixed Dy point), with the highly incompatible elements becoming more enriched by up to 40%. Given the data in Fig. 8, it is straightforward to calculate F, the melt fraction, for any of the Hobbs Lineament lavas. For the average Coleman Nunatak lavas, after olivine-correction, F--1.6%; for the lower bounding spidergram, sample 48B (Fig. 6), F = 3.2%. These melt fractions are directly dependent on our estimate for DY0; for example, if Dy 0 were as low as 0.25 × PUM, then the melt fractions would become 0.8% and 1.6%, respectively. In any event, the two bounding spidergrams in Fig. 6 can be derived from a homogeneous source with a melt fraction for 48B that is twice that for Coleman. The HIMU 'end' of the Hobbs array, represented by the average Coleman basalt in Fig. 6A, may be compared with the spidergram for 'classic' HIMU as defined for OIB. This is done in Fig. 6B, where the HIMU pattern is that of seven basalts from the islands of Mangaia and Tubuaii; the Coleman data have been adjusted down by 21%, to account for the difference in observed Mg number (Mg 51.8 for Coleman, Mg 67.5 for the HIMU average). The patterns are amazingly similar, and more or less parallel throughout; both have the negative K and Pb anomalies typical of OIBs in general, both have mild Zr-Hf anomalies arid again the Pb anomaly looks 'pinned' as it was i~LFig. 6A. The only significant difference seems to be in the Ba to Nb region, where the HIMU pattern is relatively smooth in contrast to the Coleman pattern which has a clear depletion of Th and U relative to Ba and Nb. This leads to a high B a / T h ratio in Coleman (av. 88.3) relative to HIMU (52.4). The exact significance of this difference is unclear. At face value, Ba and Th partitioning data for garnet and cpx suggests a much lower K o for B a / T h in cpx than in garnet (Hart and Dunn, 1993; Hauri et al., 1994b), so a smaller residual garnet component during Coleman melting would tend to lead to higher B a / T h ratios (note that Beattie, 1993 does have garnet B a / T h < Cpx Ba/Th). However, the Coleman and HIMU spidergrams are very nearly parallel throughout the whole REE range, suggesting that the modal gamet/cpx ratio during melting was very similar for both HI/vlU and Coleman. It is possible
239
that one can appeal to minor phases to fine-tune these elements; for example, Class and Goldstein (1997) have proposed that phlogopite and amphibole from mantle lithosphere may modulate trace element patterns in some OIB. We have suggested above that amphibole is a likely residual phase during Hobbs Lineament melting. Ba is compatible in phlogopite; residual phlogopite in the HIMU source could lead to the low B a / T h in HIMU relative to Coleman. However, K and Rb are also compatible in phlogopite, and it is not clear the B a / T h can be fractionated without upsetting Ba/Rb, etc.
8. Elemental and isotopic correlations First, as implicit in the previous section, many of the incompatible trace elements in the Hobbs suite show very good correlations with each other. These trace elements are found also to be well correlated with major elements such as FeO, particularly after the olivine normalization to Mg 73. There are also clear correlations between major and trace elements and isotopes (e.g., F e o - l g a N d / 144Nd, 2°apb/2°4pb-Nb, etc.). A few of these various correlations are shown in Fig. 9. Panels A and E show that trace element concentrations (e.g., Nb and Ce), which are argued above to be controlled by degree of melting, are correlated with major elements such as FeO and MgO. Since these major elements were argued in Section 5 to be pressure (depth) controlled, this leads to a correlation between extent of melting and depth, with the lowest extents of melting at the greatest depth (Coleman group, F ~ 1.6%, Z ~ 140 km). Panels C, D and F show that the St, Nd and Pb isotopes are also related to major elements, and therefore depth. In this case, the HIMU (Coleman) end of the isotope array (see Fig. 5) is correlated with the greatest depth of melting, and the FOZO (Holmes) end of the isotope array is from the shallower depths of melting. It might be noted that, of the other MBL localities plotted on Fig. 5B which show trends toward FOZO, the Balleny/Scott, Murphy and Siple arrays also show decreasing FeO (shallower depths) in the FOZO direction; Hudson and Fosdick show little or no change in FeO along the arrays.
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
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Finally, Fig. 9B demonstrates one of the isotopetrace element correlations which is implicit from the above correlations: the HIMU (Coleman) component represents the highest trace element concentrations, i.e., the lowest degrees of melting. This lower degree of melting at Coleman leads to slightly elevated R b / S r and N d / S m ratios (16% and 11%, respectively), which then produce Sr and Nd pseudo-iso-
chron correlations (87Sr/86Sr versus Rb/Sr, etc.) for the Hobbs Lineament volcanics (since Coleman has higher 87Sr/a6Sr and lower 143Nd/144Nd than the rest of the Hobbs localities). There are good pseudo-isochrons for the U / P b and T h / P b systems also, though in this case normal melting processes cannot produce the large range observed for U / P b and T h / P b (both a factor of 2.3). If Pb is strongly
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
buffered as a compatible element by some trace phase (sulfide?), this large range in parent/daughter ratio could be produced from a homogeneous source. Alternatively, and we feel, more likely, the Pb concentration in the HIMU and FOZO sources is different by a factor of ~Iwo to begin with (and is the fundamental reason for the large Pb isotope differences between the HIMU and FOZO sources). During the latest melting event, Pb has behaved in a normal magmatic way (i.e., like Ce), and the original factor of two deficit in Pb content of the HIMU source has been compensated for by the lower degree of melting. The result is comparable Pb contents in all the Hobbs volcanics. In this case, the U / P b and T h / P b pseudo-isochrons are in fact real mantleisochrons, and their approximate age of 230 m.y. may relate to the time of Pb depletion in the Coleman (HIMU) end-member. Thus the HIMU signature is not related to a U + Th enrichment process as frequently asserted, but rather to a Pb depletion process. Peucker-Ehrenbrink et al. (1994) have argued that a significant fraction of the Pb in the ocean crust is transferred into metalliferous sediments during ridge crest hydrothermal processes. If HIMU mantle is recycled ocean crust, as originally postulated by Hofmann and White (1982), then the Pb depletion we hypothesize at Coleman may relate to such a ridge crest process. While speculative, the pseudo-isochron age of 230 m.y. would point to fairly recent crust recycling, then, and would seem to obviate a deep-mantle transport cycle for this particular HIMU source. In addition to these numerous major-trace-isotope correlations, there are also important correlations with age of eruption of the volcanics; some of these are shown in Fig. 10. In all cases, since the basalts from Coleman show virtually no chemical change over their 9 m.y. eruption span, horizontal arrays are formed for the Coleman suite. In contrast, the rest of the Hobbs group show decreases in A1203 (and increases in FeO, MgO), increases in Nb (and other LIL-elements), and increases in 2°6pb/:°4Pb with younging of eruption age. In other words, Coleman erupts deep, low-degree HIMU-type melts through time, whereas volcanism elsewhere on the lineament starts shallow, with higher-degree FOZO-type melts. Through time, the melting becomes deeper and smaller in degree, thus approaching the Coleman
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suite in all aspects. To summarize these relationships in terms of our estimates for depth and extent of melting: (1) Coleman group: 1.6% melting at 140 km depth; HIMU in character; constant from 11.7 m.y. to 2.3 m.y.
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
(2) Cousins to Holmes group: starts at 10 m.y. with 3.2% melting at 110 km depth; FOZO in character; through time the melting gets deeper and the HIMU component increases, F decreases.
9. Implications for the WARS One of the central issues regarding the West Antarctic rift system is the role, if any, played by mantle plumes. The geochemical case for plumes rests on the chemical similarity between WARS volcanism and that of oceanic islands, where plumes are the reigning paradigm. This similarity has been argued in numerous papers (Kyle et al., 1992; Lanyon et al., 1993; Hole and LeMasurier, 1994; Hart et al., 1995), and the essence can be seen in the trace element patterns (e.g., Fig. 6) and isotopic signatures (e.g., Fig. 4), especially for the HIMU end of WARS volcanism. There are several garden-variety oceanic plumes in the WARS vicinity also, with appropriate chemical signatures to serve as analogues for some of the WARS volcanism. The Balleny Islands (see location B, Fig. 1A) have been in most plume catalogues for years (Morgan, 1972; Vogt, 1981), and, as shown in Figs. 4 and 5, the Balleny volcanics have isotopic characteristics appropriate for the 'FOZO'end of the Hobbs Lineament volcanism, as well as other volcanism in Marie Byrd Land (see Fig. 5) and Northern Victoria Land (see fig. 1 in Hart and Kyle, 1994). Scott Island (location S, Fig. 1A) has never been formally designated a plume, but its intraoceanic setting, young volcanism (WiSrner and Orsi, 1992), gravity signature (Marks and McAdoo, 1992), and striking isotopic similarity to the Balleny plume strongly implicate a plume attribution. And finally, Hart et al. (1995) argued for a plume origin for the young volcanism on Peter I Island (see location P, Fig. 1A). Note, however, that while these 'plumes' provide a FOZO-type mantle component, they do not strictly provide evidence for the HIMU component as seen at Coleman Nunatak (see Fig. 5). Aside from the geochemical case, a variety of strong arguments have been made for a large plume underlying Made Byrd Land. These include the doming and associated patterns o f volcanism (LeMasurier and Rex, 1989; Kyle et al., 1992; LeMasurier and Landis, 1996), the association with numerous geo-
physical manifestations of rifting (Behrendt et al., 1992), and the well-defined relationship between domal uplift of MBL and the volcanic history over the past 30 m.y. (LeMasurier and Landis, 1996). All of this seems quite persuasive as long as a single plume source is not argued for all of the Cenozoic volcanism in the WARS. As shown by Hart and Kyle (1994) and Rocholl et al. (1995), the volcanism in Northern Victoria Land exhibits a significant HIMU component, and to include this as part of the same plume as MBL requires a plume more than 3000 km in dimension (see fig. 2 in Behrendt et al., 1992). More importantly, as noted by Weaver et al. (1994) and by Blusztajn and Hart (1995), Mesozoic and Cenozoic HIMU-type volcanism is also present in Tasmania, New Zealand and the Campbell Plateau. These localities clearly cannot be embraced by a single plume presently centered under MBL. Short of postulating separate HIMU-type plumes for each area, we note that these areas were all contiguous prior to the breakup of Gondwana (see reconstruction in Lawver et al., 1992). As one resolution to an awkward multi-plume scenario, we suggest that a single large HIMU plume head arrived prior to breakup and underplated the lithosphere throughout this broad area. It is tempting to associate breakup with the arrival of this plume head, but as noted by Behrendt et al. (1992) and LeMasurier and Landis (1996), there is no large-scale volcanism or uplift related to the breakup event at about 85 m.y. The major volcanic event to which a plume head could be related is the Ferrar Group Large Igneous Province (Brooks and Hart, 1978), which, at ~ 185 m.y. (Encarnaci6n et al., 1996), significantly preceded NZ-MBL breakup (Hergt et al., 1991). Alternatively, because the area to be underplated by a single plume head is still very large ( > 5000 km), one could postulate a plume head which had thermally entrained such large fractions of mantle on the ascent that its thermal anomaly (and buoyancy flux) was highly dispersed, and in fact insufficient to cause plume melting, especially if its rise was terminated at depth by a thick Gondwanan continental lithosphere. Further flattening and radial expansion would provide a young proto-lithosphere under a vast region which could, after a suitable 'incubation' period (Kent et al., 1992), provide a fossil-HIMU mantle
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
source during subsequent rifting events (Halliday et al., 1990; Stein and Hofmann, 1992). This underplating/rifting scenario, origirtally proposed by Brooks et al. (1976) as a general phenomenon, can also be accompanied by metasomatism and enrichment of overlying mantle. (We concluded earlier that the Hobbs mantle source is both enriched and hydrous.) Thus the original plume need not have been HIMU in character but could have grown in the HIMU character subsequent to underplating, as postulated for the Cameroons by Halliday et al. (1990) and for the Antarctic Peninsula alkalic basalts by Hole et al. (1993). By postulating that the HIMU mantle component was emplaced prior to breakup, we can focus back on the Cenozoic volcanism in the MBL region. While the high 2°6pb//2°4pb ratios make the HIMU component particularly conspicuous, the FOZO component is probably more ubiquitous throughout the WARS, and potentially more important in understanding the evolution of WARS volcanism. Part of this is a nomenclature problem; in the past, volcanics with 2°6pb//2°4pb ratios of 19.5 were identified with a HIMU component. Based on the evidence of Fig. 5, we propose that only 2°6pb//2°4pb ratios greater than 20.5 be assigned HIMU 'status' (the oceanic HIMU end-member has a 2°6pb/204pb of 21.8; Hauri and Hart, 1993). The great bulk of Cenozoic volcanism in the west Antarctic region in fact has 2°rpb//2°4pb ratios in the 19.3-19.8 range, and this is more appropriately identified with the FOZO end-member. This is the dominant component in the three 'real' plumes in the area (Scott, Balleny and Peter I), and is therefore central to the issue of whether the WARS volcanism is plume-related, as argued by Kyle et al. (1992), Behrendt et al. (1992), Hole and LeMasurier, 1994, and LeMasurier and Landis (1996). Based on the isotopic arrays in Fig. 5, it would appear that much of the MBL volcanism contains a FOZO component, and therefore is consistent with some 'plume-connection'. For help in understanding this connection, we return to the evidence derived from the Hobbs Lineament volcanism: melting of the FOZO mantle is shallower than that of the HIMU mantle (110 km versus 140 km), represents a larger degree of melting (3.2% versus 1.6%), and migrates to the deeper
243
HIMU regions during the period from 10 m.y. to 5 m.y. Note, however, that these two mantle sources are erupting simultaneously during the 10-12 m.y. period (Fig. 10), from almost contiguous locations on the Hobbs Lineament (Coleman Nunatak and Patton Bluff are only l0 km apart). FOZO eruptions then commence at the far northern end of the lineament at 8.2 m.y. (Holmes Bluff, 48A). If a single large plume underlies all of MBL, how do we explain coeval volcanism at contiguous localities from different mantle sources, of very disparate depths? One solution would be to assert a heterogeneous two-component plume, with different solidus temperatures for the two components. For example, veined-mantle models (Wood, 1979; Zindler et al., 1984; All~gre et al., 1984; Hirschmann and Stolper, 1996) propose that pyroxenite or eclogite layers in a lherzolite might begin melting 100°-150°C lower in temperature than the lherzolite host; this would translate to a 25-40 km depth difference (for a difference in solidus slope and adiabat slope of 4°/km). If the pyroxenite layers are HIMU in character, and the lherzolite is FOZO, this would plausibly explain the deeper HIMU melts at Coleman relative to the shallower FOZO melts at Patton Bluff. With time, the melting gets deeper and the FOZO component becomes less dominant all along the lineament, perhaps because the upper surface of the plume becomes cooled upon impingement with the lithosphere. In this scenario, however, part of the logic unravels, as the depth differences we calculated were based on major element constraints from lherzolite melting experiments (Herzberg and Zhang, 1996). Similar melting experiments on marie (i.e., eclogitic) compositions (Yasuda and Fujii, 1994) show little or no dependence of melt composition on pressure, in the range 3-14 GPa. Also, the Yasuda and Fujii melt compositions are most similar to the FOZO compositions from Patton Bluff (37A) and Holmes Bluff (48B), suggesting that the melts of HIMU affinity (Coleman) derive from lherzolite, not pyroxenite. This argument is not firm, however, because highpressure melt compositions in mafic systems are quite dependent on the composition of the starting material. Yasuda and Fujii (1994) used an N-MORB starting composition, which at pressure is a quartzeclogite; a lower-silica starting material might conceivably lead to melt compositions more like the
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Coleman basalts. Also, in this model, it becomes difficult to explain our apparent success in Section 7 in modeling the full trace element data set in terms of a single uniform source. Surely the trace element partitioning from lherzolite and pyroxenite sources will look radically different! In any event, this heterogeneous plume scenario resurrects the implausibility problem of having HIMU-like volcanism occurring in many areas now well removed from one another (Tasmania-MBL). Alternatively, in a plume scenario, one could resort to major 'topography' at the base of the lithosphere (Saunders et al., 1992), to fit the constraints posed by the Hobbs Lineament data. If WARS rifting removed a significant thickness of a fossil HIMU lithosphere (30 km layer over a lateral distance of 10 km), then a FOZO plume could rise to ~ 110 km and begin melting, whereas the deeper adjacent lithosphere would melt to a lower degree during minor adiabatic uplifts associated with the rifting. Once on the solidus, the small extents of melting proposed here (1.6-3.2%) can be produced by only 3 to 5 km of adiabatic decompression (approximately 0.6% melt per km of uplift, using the preferred melting parameters of Hart, 1993). Finally, we explore a scenario in which no currently active plumes are needed. This accepts the relative depths of the FOZO and HIMU mantles (110 km versus 140 km) and locates both as part of a layered fossil plume lithosphere, a HIMU layer underlying a FOZO layer. This layering becomes a natural consequence of a plume-head model, if we adopt the Hart et al. (1992) explanation for the FOZO component. They showed that many oceanic plumes form elongated arrays in 3-D isotopic space, which converged on a region named FOZO (Focus Zone). FOZO was interpreted as lower mantle which was thermally and viscously entrained (coupled) to a plume arising from the core-mantle boundary layer. The dynamics of this entrainment were treated for steady-state plume tails by Hauri et al. (1994a); the case for plume heads may be similar (Griffiths and Campbell, 1990), in that the outermost sheath of the plume will be the entrained material. In the Hart et al. (1992) model, this entrained material is identified with FOZO. Therefore, a plume head comprised initially of HIMU mantle will arrive at the base of the lithosphere accompanied by a sheath of FOZO
attached to it; on underplafing and spreading, this proto-lithosphere will be a FOZO layer underlain by a HIMU layer. We do not envision that any melting is necessary during this underplafing event. If melting did occur, it would be dominated by the HIMU plume component, as Farnetani and Richards (1995) have demonstrated that plume head melts will be dominated by the original plume material, not the entrained material. This model is then beset by the same melting problem as the active plume model. As entrained material, the FOZO mantle is initially cooler than the HIMU mantle. How do we produce melting at 110 km depth (FOZO) right adjacent to melting at 140 km depth (HIMU)? As above, one solution is to appeal to a lower solidus temperature for the HIMU component (as in an eclogitic or pyroxenitic source scenario). Alternatively, and perhaps realistically, the temperature gradient in the fossil layering, inherited from the plume head, may be sub-parallel to a lherzolite solidus. The entrained material is clearly cooler than the plume interior. The actual thermal gradient will depend on radial distance from the plume axis, but is in the range of 2°/km (on wings) to > 10°/km (on axis) (Fametani and Richards, 1995). These gradients are approximately the same as the solidus slope for lherzolite in the 100-140 km depth range (3.4-4.5°C/km; Zhang and Herzberg, 1994). During subsequent Cenozoic tiffing, adiabatic decompression of this FOZO and HIMU layered 'package' could produce coeval melting over a substanfial depth range. With respect to lateral localization of the melting, we follow the model of King and Anderson (1995) and suggest that this melting may be a natural consequence of the lithospheric asymmetry left behind after breakup. Small-scale convection driven by differences in heat flux and lithospheric thickness (for example, at a rifted margin; Mutter et al., 1988), will drive an upwelling mantle flow from under the thick lithosphere to regions of thinner lithosphere. While extension is still required in this model, abnormally hot mantle (e.g., plume-type) is not. While the total amount of Cenozoic extension in the WARS is not large (Lawver and Gahagan, 1994), it may be sufficient to enhance small-scale convection if it is focused or localized in the regions of lithospheric asymmetry.
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Acknowledgements We commend AI Hofmann, the 1996 Goldschmidt medalist, and recognize with pleasure the decades of stimulation and liveliness he has brought to our science. We are most grateful to Peter Kelemen for his tutorials in MELTS and his unselfish sharing of his mind, his computer and his time. We are grateful to Mark Kurz for access to the VG-354 facility at WHOI. Very helpful and perceptive reviews were provided by Simon Turner, Andy Saunders and Malcolm Hole; we have hardly done justice to them in the limited turn-around time available. This work was supported by DPP-9117853 and DPP 9419094 to SRH and DPP-7727546 and DPP8020836 to WEL.
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CHEMICAL GEOLOGY fMC~DING
ELSEVIER
ISOTOPE GEOSCIENCE
Chemical Geology 139 (1997) 223-248
Hobbs Coast Cenozoic volcanism: Implications for the West Antarctic rift system Stanley R. Hart a,*, Jerzy Blusztajn a, Wesley E. LeMasurier b, David C. Rex c a Department of Geology and Geophysics, Woods Hole Oceanographic Institution, Woods Hole, MA 02543, USA b Department of Geology, University of Colorado, Denver, CO 80217, USA c Department of Earth Sciences, University of Leeds, Leeds, LS2 9JT, UK
Received 31 December 1996; accepted 8 January 1997
Abstract
Basaltic lavas were erupted from a 40-km-long lineament near the Hobbs Coast of Made Byrd Land, Antarctica, over the period from 11.7 m.y. to 2.3 m.y. ago. The lavas from the southernmost locality, Coleman Nunatak, are virtually constant in major, trace dement and isotopic composition over this entire age span. Their high FeO-low A120 3 character indicates melting of garnet peridotite at about 140 km depth. There is no evidence for the involvement of ancient continental lithosphere or MORB asthenosphere in the magmatism. Isotopically, the lavas show the highest 2°sPb/2°4pb ratios (up to 20.7) of any of the Cenozoic volcanism associated with the West Antarctic rift system (WARS). This HIMU isotopic signature is also clear in the trace element patterns, which closely mimic end-member HIMU basalts from the oceanic islands of Tubuai and Mangaia. From the other localities along the Hobbs Lineament, the earliest volcanism, which is coeval with that at Coleman Nunatak, is of shallower derivation ( ~ 110 km), and isotopically like the oceanic FOZO end-member (Z°tPb/2°4pb ~ 19.5). The trace-dement patterns are similar to those at Coleman, but less enriched in the most incompatible elements by a factor of two. Modeling of the trace element data is consistent with a uniform mantle source composition, depleted in major dements, but hydrous and mildly enriched in the incompatible and LREE. Inversion for the bulk distribution coefficients of the source mantle reveals a spidergram with a marked negative Ti anomaly and marked positive anomalies for K, Sr, Zr and Hf. From this modeling, the extent of melting at Coleman is inferred to be ~ 1.6%, as compared to ~ 3.2% during the earliest volcanism elsewhere on the lineament. With time, the volcanism from these other localities progresses to greater depth, becomes more HIMU in character, and lower in extent of melting (i.e., approaches the character of basalts from the Coleman locality). The FOZO component is prevalent as a mixing end-member in WARS volcanism from numerous other Made Byrd Land (MBL) and Northern Victoria Land (NVL) localities. It is also the main constituent of the three nearby oceanic plumes (Balleny, Scott, Peter I islands). The HIMU component is at best a minor constituent of these oceanic plumes, but is present at several other MBL and NVL localities, as well as in pieces of Zealandia which were adjacent to this coast of Antarctica prior to fragmentation of Gondwana. We propose that this HIMU mantle source was emplaced under Gondwana lithosphere prior to breakup, as a large weak plume head, with little or no accompanying volcanism. This 'fossil-plume' proto-lithosphere is now being sampled during WARS extension. Likely mechanisms for the volcanism relate either to small-scale
* Corresponding author. FAX: 508-457-2175; E-mail: [email protected]. 0009-2541/97/$17.00 ~9 1997 Elsevier Science B.V. All fights reserved. PH S0009-2541 (97)00037-5
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convection associated with strong basal topography of the lithosphere (such as that recorded by the Hobbs Lineament volcanism), or to emplacement of a new plume, which may in part be driving the extension. © 1997 Elsevier Science B.V. Keywords: volcanism;rifting; plumes; isotopes;trace elements
1. Introduction Cenozoic volcanism in Antarctica is largely confined to west Antarctica and ranges in age from ~ 30 m.y. to recent (Armstrong, 1978; LeMasurier and Rex, 1983; Smellie et al., 1988; LeMasurier and Thomson, 1990), though the preponderance of volcanism is less than 6 - 8 m.y. old. Aside from the subduction-related volcanic rocks of the northern Antarctic Peninsula and Scotia Arc, the remainder are alkaline basalts and related differentiated rocks; tholeiitic-series rocks are not common. While the tectonic setting of the subduction volcanic rocks is clear enough, the alkalic volcanics occur both as isolated 'hotspots' (Balleny Island, Scott Island, Peter I Island) and as an alignment of volcanic provinces along the edges of the West Antarctic rift system. As defined by LeMasurier (1990a), this rift system extends from the Ross Sea to the Bellingshausen Sea; the volcanic provinces along the Transantarctic Mountain front in Victoria Land lie on the south side of the rift, and the volcanic provinces in Marie Byrd Land (MBL), and extending to Jones Mountains, lie on the northern side of the rift (Fig. 1). While the bulk of the mainland volcanism is presumed to be related to rifting and extension (LeMasurier, 1978; Cooper and Davey, 1985; Behrendt and Cooper, 1991), the underlying causes are not well understood. Some researchers subscribe to variants of conventional active or passive rift magmatism models (Fitzgerald et al., 1986; Tessensohn and Wtmer, 1991). Others appeal to lithospheric effects associated with the continental plate overriding a mantle 'hot line' associated with earlier continental break-up (Smith and Livermore, 1991). Then there are a diversity of models involving some connection to mantle plume activity (LeMasurier and Rex, 1989; Behrendt et al., 1991, 1992; Kyle et al., 1992; Hole and LeMasurier, 1994). On a grander scale, this volcanism forms part of a major arcuate belt of alkaline volcanism that extends from southern South America (Stem et al., 1990) via the Antarctic Peninsula, Marie Byrd Land and Victo-
ria Land to New Zealand (Gonzalez-Ferran, 1972; Cooper et al., 1982). We believe that all of this volcanism shows a common geochemical and isotopic 'fingerprint'. While a full understanding of the WARS volcanism cannot be achieved without viewing it in this larger context, regional studies are important for elucidating the temporal and spatial evolution of the volcanism. We report here on detailed chemical studies of one key lineament in Marie Byrd Land, where the HIMU isotopic signature which characterizes this whole Cenozoic volcanic belt is most prominently displayed.
2. Regional setting The basaltic rocks of the Hobbs Coast region are part of a large alkaline basalt-trachyte volcanic province that lies within the West Antarctic rift system (Fig. 1), about 2000 km from the nearest active plate boundary. The province has been volcanically and tectonically active from about 25-30 m.y. to the present. It is centered over a large Neogene tectono-magmatic dome whose growth and concurrent volcanism are believed to be related to the inception of mantle plume activity around 30 m.y. (LeMasurier and Landis, 1996). Extension and block faulting, related to this doming, have produced basin and range topography in the Hobbs Coast area and throughout Marie Byrd Land. Basaltic rocks (basalt, basanite, tephrite, hawaiite) are the most widespread and voluminous of the volcanic rocks, and have been erupted almost continuously throughout Neogene time. They are found as basal sections up to several thousand meters thick, underlying large felsic shield volcanoes, and as isolated flows, hyaloclastite sheets, cinder cones and tuff cones that rest on basement rock (satellite centers), or on the flanks of the shield volcanoes (LeMasufier, 1990b). The basement rock in this region is part of a large Devonian to Cretaceous magmatic arc terrain that
S.R. Hart etal./Chemical Geology 139 (1997) 223-248 1135eW
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Fig. 1. Index maps. (At Map of Antarctica showing the position of map B (square), the position of the Marie Byrd Land dome (dotted line), and the southern boundary of the West Antarctic rift system (dashed line). Note that, for maps of this scale, the Greenwich meridian is oriented up by internatiolml convention, but that for maps (B) and (C), north is up. Abbreviations are: AP, Antarctic Peninsula; B, Balleny Islands; P, Peter I Island; S, Scott Island. (B) Made Byrd Land volcanic province, showing major central volcanoes and their inferred potential for future activity (circles). 'Satellite centers' are relatively small volcanic centers that consist of cinder cones, tuff cones, flows and hyaloelasfite sequences of predominantly basaltic compositions. (C) Reconnaissance geologic map of the area described in this study, modified from LeMasurier (1990c).
extends along the coast of West Antarctica from the Antarctic Peninsula to the Ross Sea, representing a prolonged interval of subduction prior to the 85 m.y. breakup of this sector of Gondwanaland. In the Hobbs Coast and its environs the basement is dominated by a 9 5 - 1 0 2 m.y. suite of A-type granitoids (Weaver et al., 1994), and less common talc-alkaline granitoids with U - P b ages of 110-322 m.y. (Mukasa,
1995). These rocks intrude early Paleozoic metaclastic rocks. Granitoids from this region have Nd model ages of 900-1200 m.y., suggesting the presence of a Proterozoic basement that is not exposed anywhere in Made Byrd Land (Pankhurst et al., 1995). Thus, the work done so far suggests a relatively juvenile lithosphere for this region. Previous trace element and Sr and Nd isotopic studies of the Cenozoic
Coleman Nunatak Coleman Nunatak Coleman Nunatak Coleman Nunatak
Coleman Nunatak
Coleman Nunatak
Coleman Nunatak
Cousins Rock
Patton Bluff Shibuya Peak
Kouperov Peak
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Holmes Bluff
35K 35E 35M 35J
46B
46,I
46D
36B
37A 38A
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48B
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1.33
1.644
1.468
1.487
1.482
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%K
2130 2311
3038 3040 3020 3022 3461 3480 3021 3.23 3366 3376
Ar Nr.
0.04262 0.04219
0.01593 0.01758 0.01714 0.01915 0.0654 0.0700 0.04550 0.04482 0.0342 0.0336
Vol. 4°Ar rad.
39.2 28.2
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8.17 + 0.33
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Age (M a)
Basalt flow on WAES (LeMasurier and Rex, 1983). Basalt flow in valley cut into WAES (LeMasurier and Rex, 1983).
Basalt flow on granite. (LeMasurier and Rex, 1983). Basalt clast in hyaloclastite, middle of exposed section (LeMasurier and Rex, 1983). Basalt flow on granite.
Dike in eroded cinder cone.
South cliff, basalt clast from base of hyaloclastite section.
Flow rock, base of south cliff.
~ Same horizon as 35E and 35M, different locality.
Top of section, tuff cone (LeMasurier and Rex, 1983). Top hyaloclastite underneath tuff cone (LeMasurier and Rex, 1983). Top hyaloclastite underneath tuff cone (LeMasurier and Rex, 1983). ~ Same horizon as 35E and 35M, different basalt clast.
Comments
Constants: ~8 = 4.962 × 10- I 0 y r - t ; E = 0.581 × 10- l o y r - t ; 40K = 0.01167 atom%; 4oAr/36At = 288.5. AH dated materials are whole rock. Analyses were done at the University of Leeds, following procedures described by LeMasurier and Rex (1989).
Locality
Sample Nr.
Table 1 K - A t data for Hobbs Coast Nunataks
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
basalts have found no evidence for contamination by this lithosphere (Futa and LeMasurier, 1983; Hole and LeMasurier, 1994).
3. Field and age relationships The samples described were collected from a north-south line of nunataks, extending from Coleman Nunatak, in the south, to Holmes Bluff, that project above the continental ice sheet just inland of the Hobbs Coast (Fig. 1C; Table 1). We will call this the Hobbs Lineament. Coleman Nunatak is flattopped, about 1.3 krn long, and composed of subhorizontaUy bedded basaltic hyaloclastite breccias with a total exposed thic~mess of ~ 150 m. This section is, in turn, overlain by a low tuff cone at the north end. Glacial striations on the highest points indicate that the whole nunarak has been overridden by ice. Locality 35 is from the vicinity of the tuff cone, and represents the highest stratigraphic levels at Coleman Nunatak. Locality 46 is at the south end of the Nunatak, at the top of a ~ 100 m cliff with limited access to the base. The cross-section exposed here contains multiple stratigraphic discordances, interbedded flow rock, and foreset-bedded hyaloclastite with redeposited tuff clasts. There is, therefore, some uncertainty about the stratigraphic relationships of the samples because of the possibility of redeposition of clasts. However, the ages are generally concordant with stratigraphic position; 46D and 46J are from the bottom of the section, 35K from the top, and the remainder in between (Table 1). All samples are from lava clasts or entire flows, interbedded in the hyaloclastite. The K - A r ages were obtained from glass--free samples. The ages suggest that this section formed by multiple, small-scale eruptions over a pe~riod of ~ 5 - 9 m.y., in marked contrast to the relatively large magma production rate at a shield volcano like Mount Murphy (Fig. 1B), where it appears that ~ 500 km 3 of basalt was erupted in < 1 m.y. (LeMasurier, 1990c). Field relationships at the remaining localities are described in Table 1 and Fig. 1C. At Patton Bluff and Holmes Bluff, basalt flows rest on granite that has been truncated by the very low relief West Antarctic erosion surface (WAES). Displacement of the WAES to different levels in adjacent nunataks
227
provides evidence for the effects of Neogene block faulting (LeMasurier and Landis, 1996). At Kouperov Peak and Holmes Bluff (sample 50A) basalt flows lie on the granite floor of valleys cut into the WAES. Samples 48B and 50A therefore satisfy a stratigraphic test of their ages, in that flow 50A, with an age of 6.27 m.y., fills a valley cut through an older flow (48B, 8.17 m.y.). To summarize, volcanism along the Hobbs Lineament started 11.7 m.y. ago at Coleman Nunatak; volcanism continued at Coleman until 2.34 m.y. ago. Coleman volcanism was joined at 9.97 m.y. at nearby Patton Bluff, then spread to the north and central parts of the lineament at 8.17 m.y. (Holmes Bluff), and 8.21 m.y. (Kouperov Peak).
4. Sample and analytical details The number associated with each sample is a locality number (see Table 1 and Fig. 1); the letters refer to different hand-specimens from a given locality. Megascopically, most of the samples are finegrained to aphanitic; the locality 46 samples are somewhat coarser-grained. Most are dense to only slightly vesicular; 35J and 35K are somewhat more vesicular. Phenocrysts of olivine and clinopyroxene are both usually present in small amounts (0-5% each); 35E and 35K are slightly richer in clinopyroxene (5-10%); plagioclase is present as a trace phenocryst in 37A, and 38A, and at about 5% in 47A. Eight of the samples are classified as basanites, using the LeBas et al. (1986) nomenclature (see Table 2). There are two hawaiites (but close to the basanite boundary), two alkali basalts and only one differentiated rock, a mugearite. All but one of the samples are nepheline-normative (48B has just a trace of nepheline; 37A shows 5% hypersthene in the norm). Except for the mugearite, the Mg numbers range from 49 to 61. All of the samples appear fresh and unweathered, though several contained thin oxidative finds which were removed before powdering. Though we did not do volatile analyses, various trace elements which are typically sensitive to weathering and alteration show no hint of such effects. For example, only one sample (38A) shows an anomalous R b / C s ratio (41), whereas the other twelve show values which cluster very tightly around a value of 112 ( + 8 % ,
0.702843 0.512873 20.637 15.715 39.732 43.55 3.95 15.00 12.85 0.221 7.02 10.21 4.55 1.74 0.90 6.29 0.520 32%
531 37.6 0.36 1045 294 98.4 36.5 17 264 52 98 20 24 94 64.3 123.4
87Sr/S6Sr 143Nd/144Nd 2°6pb/2°4Pb 2°7pb//204pb 2°spb/2°4pb SiO 2 TiO 2 Al203 FeO* MnO MgO CaO Na20 K20 P205 Na20+K20 Mg number Olivine, Mg number 73
Ba (ppm) Rb Cs Sr Zr Nb Y Sc V Ni Cr Ga Cn Zn La Ce
Rock type: Age(m.y.):
46D Coleman Nunatak Basanite 7.9(0.36)
Sample Nr.: Location:
533 36.7 0.35 1047 297 97.8 35.7 27 255 51 95 21 24 95 64.8 124.1
0.702809 0.512919 20.582 15.688 39.641 43.59 3.85 15.08 13.02 0.222 6.97 10.08 4.52 1.77 0.89 6.29 0.515 33%
46I Coleman Nunatak Basanite 11.7(1.0)
309 64 103 19 44 83 64.5 122.6
533 38.2 0.34 1076 301 97.3 36.4
0.702815 0.512883 20.619 15.727 39.733 44.32 3.92 14.95 13.06 0.12 7.15 10.08 3.84 1.75 0.81 5.59 0.520 33%
4613 Coleman Nunatak Basanite 3.17(0.13)
516 36.7 0.33 1028 293 95.0 35.8 29 262 53 95 20 29 99 64.4 122.6
0.702858 0.512908 20.652 15.746 39.778 43.37 4.00 14.91 13.12 0.218 7.23 10.20 4.35 1.71 0.89 6.06 0.522 33%
35E Coleman Nunatak Basanite 3.19(0.33)
508 36.5 0.36 1046 293 96.6 37.0 28 260 55 96 16 25 97 63.4 120.7
0.702830 0.512894 20.629 15.708 39.692 43.42 4.00 15.07 12.92 0.222 7.25 10.24 4.26 1.72 0.90 5.98 0.526 32%
35I Coleman Nunatak Basanite 2.95(0.12)
Table 2 Chemical data for volcanic rocks from the Hobbs lineament, Antarctica
313 67 102 18 43 91 65.4 124.2
529 39.1 0.37 1071 296 103.0 37.3
0.702807 0.512908 20.689 15.769 39.904 43.52 3.79 15.34 13.25 0.27 6.89 9.86 4.47 1.73 0.88 6.20 0.507 35%
35M Coleman Nunatak Basanite 2.34(0.11)
517 37.1 0.36 1040 295 96.5 36.9 19 256 51 94 15 22 97 64.1 123.3
0.702833 0.512900 20.641 15.711 39.729 43.77 3.91 15.05 12.92 0.219 6.96 10.09 4.41 1.77 0.89 6.18 0.516 33%
35K Coleman Nunatak Basanite 2.63(0.11)
117 12 12 18 29 100 68.2 128.3
489 35.7 0.30 953 344 94.8 35.2
0.702883 0.512892 20.450 15.689 39.614 50.17 2.31 16.62 12.50 0.21 3.30 6.44 5.30 2.01 1.13 7.31 0.343 50%
3613 Cousins Rock Mugearite 5.3(0.2)
237 192 329 19 63 83 35.5 68.7
291 22.5 0.19 652 222 52.2 29.6
0.702758 0.512961 19.502 15.643 38.882 47.01 2.69 15.20 11.73 0.08 9:42 9.43 2.58 1.28 0.58 3.86 0.614 19%
37A Patton Bluff Basalt 9.97(0.40)
235 99 211 18 66 98 46~9 90.4
375 24.8 0.61 829 234 70.4 29.7
0.702831 0.512911 20.195 15.639 39.242 46.52 2.84 15.14 12.28 0.18 7.25 9.63 3.92 1.44 0.79 5.36 0.539 29%
38A Shibuya Peak Hawalite 4.66(0.50)
232 51 82 20 47 90 43.9 86.7
368 23.7 0.21 762 252 62.6 33.4
0.702700 0.512952 19.863 15.627 39.004 46.74 2.78 17.17 11.56 0.22 5.64 9.26 4.27 1.61 0.77 5.88 0.491 32%
47A Kouperov Peak Basanite 8.21(0.33)
224 154 285 20 63 82 31.1 60.2
257 14.7 0.11 576 203 46.1 27.6
0.702535 0.513008 19.566 15.631 38.857 48.85 2.20 15.74 10.89 0.10 7.84 9.58 3.43 0.93 0.44 4.36 0.588 21%
48B Holmes Bluff Basalt 8.17(0.33)
227 113 220 17 57 84 48.2 87.5
463 27.6 0.24 900 215 70.0 28.6
0.702716 0.512945 20.032 15.688 39.339 46.12 2.45 15.83 12.18 0.21 7.65 9.41 4.17 1.14 0.85 5.31 0.554 27%
50A Holmes Bluff Hawalite 6.27(0.25)
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5.40 88.2 14.13 8.25 384 27.2 104.4 2.99 16.34 1.68 44.9 3.81 5.49 0.575 62.3
Ba/Nb Ba/Th Ba/Rb Ba/La K/Rb K/Ba Rb/Cs Zr/Nb Nb/Th Sm/Hf Ce/Pb Th/U La/Sm U/Pb Nb/U
5.45 89.9 14.54 8.23 401 27.6 104.7 3.04 16.49 1.70 44.9 3.68 5.48 0.583 60.7
15.18 58.3 11.83 3.54 9.24 1.35 6.87 1.25 3.14 0.44 2.62 0.40 5.76 5.93 6.94 1.61 2.76 5.48 88.0 13.97 8.26 381 27.3 112.2 3.09 16.06 1.65 50.0 3.76 5.56 0.657 60.5
14.99 57.3 11.60 3.51 9.13 1.34 6.78 1.24 3.17 0.44 2.59 0.39 5.89 6.06 7.04 1.61 2.45 5.43 90.1 14.08 8.01 387 27.5 111.1 3.08 16.58 1.71 47.7 3.60 5.41 0.619 59.8
15.17 57.7 11.90 3.57 9.50 1.36 7.08 1.30 3.24 0.44 2.63 0.41 5.59 5.57 6.97 1.59 2.57 5.36 87.6 13.97 8.02 391 28.1 101.5 3.03 16.65 1.70 44.0 3.74 5.39 0.566 62.3
14.86 58.1 11.76 3.57 9.62 1.34 7.10 1.27 3.16 0.44 2.69 0.42 5.65 5.80 6.91 1.55 2.74 5.14 87.7 13.53 8.09 368 27.2 105.7 2.87 17.08 1.75 54.9 3.61 5.54 0.739 61.7
14.48 57.9 11.80 3.76 9.65 1.38 7.40 1.41 3.50 0.44 2.66 0.40 5.62 6.03 6.76 1.67 2.26 5.36 86.5 13.92 8.07 396 28.4 103.2 3.06 16.13 1.68 59.6 3.71 5.65 0.778 59.9
15.26 58.3 11.34 3.49 9.19 1.34 6.78 1.23 3.15 0.45 2.64 0.39 5.66 5.98 6.77 1.61 2.07 5.16 73.5 13.71 7.17 468 61.2 118.9 3.63 14.26 1.38 51.1 3.91 6.24 0.677 55.8
15.52 57.7 10.92 3.27 8.50 1.23 6.55 1.19 3.10 0.45 2.75 0.43 5.86 6.65 7.90 1.70 2.51
Following the age, number in parentheses is the uncertainty, in m.y. 87Sr/86Sr and 143Nd/I'~Nd normalized to 0.71024 (SRM987) and 0.511847 (La Jolla). Majors recalculated to 100% volatile-free, all Fe as FeO, Mg number calculated with FeO = 0.9 FeO *. Olivine, Mg number 73 is the percent of equilibrium olivine added incrementally to bring Mg number to 73. Zr, So, V, Ni, Cr, Ga, Cu, Zn by XRF; all other traces by I C P / M S .
15.09 59.0 11.71 3.63 9.24 1.33 7.06 1.27 3.17 0.44 2.65 0.41 5.47 6.02 6.99 1.58 2.75
Pr Nd Sm Eu Gd Tb Dy Ho Er Tm Yb Lu Ta Th Hf U Pb 5.57 86.4 12.95 8.21 473 36.5 118.3 4.25 15.49 1.42 26.4 3.79 4.80 0.342 58.7
8.52 33.8 7.38 2.26 6.45 0.96 5.38 1.02 2.63 0.38 2.31 0.37 3.13 3.37 5.21 0.89 2.60 5.33 88.2 15.12 8.00 481 31.8 40.7 3.32 16.56 1.66 30.0 3.54 5.29 0.399 58.7
10.61 42.8 8.86 2.88 7.74 1.09 5.94 1.11 2.91 0.36 2.12 . 0.30 3.87 4.25 5.34 1.20 3.01
10.09 42.1 8.90 2.96 7.37 1.04 5.93 1.10 2.70 0.35 Z07 0.31 3.41 3.96 4.69 0.99 2.46 6.61 116.9 16.78 9.60 342 20.4 115.0 3.07 17.68 1.90 35.6 4.00 5.42 0.402 70.7
7.15 28.5 6.49 2.15 6.02 0.90 5.2i 1.02 2.74 0.36 2.17 0.32 2.52 2.80 4.57 0.78 2.15 5.57 91.8 17.48 8.27 523 29.9 133.6 4.40 16.46 1.42 28.0 3.59 4.79 0.363 59.1
10.56 41.8 8.91 2.99 7.81 1.13 6.50 1.22 3.18 0.42 2.50 0.37 3.57 4.09 5.69 1.13 3.58 5.88 90.0 15.55 8.38 566 36.4 112.7 4.03 15.30 1.57 24.2 3.62 4.93 0.316 55.4
;zt
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230
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
ltr). This compares well with the canonical value of 85-95 found for fresh ocean island basalts (Hofmann and White, 1983; Morris and Hart, 1983). The T h / U ratio of all the samples is constant at 3.72 ( + 4%, l tr); Th is quite robust to alteration, whereas U can be significantly 'mobile'. Sr, Nd and Pb isotopic analyses were carried out with conventional techniques (Hauri and Hart, 1993), using sample powders which had been leached in warm 6 N HC1 for 1 h. Sr and Nd data carry precisions of _ 0.003%, and are reported relative to 0.71024 (NBS 987) and 0.511847 (La Jolla). Pb is corrected ratio by ratio for fractionation, using NBS 981 and the values given by Todt et al. (1996). Major and trace elements were determined on unleached powders prepared in an agate shatterbox. Major elements and some trace elements (Ba, Sr, Cr, Ni, Nb, Ga, Y, Rb, Zr, Cu, Zn, V) were determined by XRF; the other trace elements (plus Ba, Sr, Rb, Nb, Y) were determined by ICP/MS at Washington State University. Much of the XRF data has been previously published (LeMasurier, 1990c). The chemical and isotopic data are given in Table 2, with the samples arranged in order of locality, from south to north along the Hobbs Lineament.
50 4948" 47" © 4645" ~2
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20
5. C o n d i t i o n s o f m e l t i n g a n d d i f f e r e n t i a t i o n
The major element data are shown in Figs. 2 and 3 in selected variation diagrams, both as raw data and as olivine-corrected data. Noteworthy in the raw data is the very tight clustering for samples from the two localities on Coleman Nunatak (46, 35; squares), despite their considerable range in age from 11.7 m.y. to 2.3 m.y. Typically, this Coleman group will lie outside (SiO 2) or at the end (Na20) of the data fields for the other localities. The other localities show more dispersed data arrays, with one extreme (low Na20) being the hypersthene-normative basalt from Patton Bluff (37A). The raw data for the mugearite (36B) is not shown, as it lies well away from the basalt data and would expand the plots excessively (it has only 3.3% MgO, and 50.2% SiO 2, for example). Comparison of the raw basalt data is hindered by the significant range in Mg number (49 to 61). For this reason, we have attempted a first-order fractiona-
2
MgO Fig. 2. Correlation of SiO 2, CaO and N a 2 0 with MgO, Hobbs Lineament volcanics. The raw data (Table 2) is enclosed in fields;
the other symbolsare for data which has been correctedto Mg 73 by incremental addition of equilibrium olivine (see text). The curves are the near-solidus garnet lherzolite melt compositions from the Herzberg and Zhang (1996) algorithms, labeled with pressure in GPa. Herzberg and Zhang (1996) do not give an algorithm for SiOz, but state that it is fairly constant at 45-47% over the pressure range 2.5-10 GPa. The same symbols will be maintained in all figures. Squares are Coleman series 35 and 46 samples, undifferentiated; X -- 38A; Y = 47A; dots = 50A and 48B; triangle = 37A. The mugearite is not plotted.
tion correction, on the assumption that only olivine has been a fractionating phase. Clearly this is a simplification, as clinopyroxene is commonly present as a phenocryst phase, along with olivine. We modeled possible Cpx fractionation effects by using V -
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
18 16-. ~
I18" 16"
12-
~..,~12-~ 5
10-
10"
86
8" ....
5
231
I' 10
•
6 15
MgO
20
.... 10
I .... I .... I .... 11 12 13 14 FeO
Fig. 3. Correlationof AI:~O3, MgO and FeO, HobbsLineamentvolcanics.The 'as is' data (Table 2) is enclosed in fields; the other symbols are for data which has been corrected to Mg 73 by incremental addition of equilibrium olivine (see text). The curves are the near-solidus garnet lherzolite melt compositionsfrom the Herzbergand Zhang (1996) algorithms, labeled with pressure in Gpa. Squares are Coleman series 35 and 46 samples, undifferentiated; X = 38A; Y= 47A; dots = 50A and 48B; triangle= 37A. The mugeariteis not plotted. C a O / N a 2 0 relationships. C a O / N a 2 0 is virtually independent of pressure of melting (Herzberg and Zhang, 1996), and of olivine fractionation, but is very sensitive to Cpx fractionation. V, with a C p x / m e l t partition coefficient of ~ 4 (Hart and Dunn, 1993; Jenner et al., 1993), was modeled with fractional crystallization. The data do not lie along a Cpx fractionation tJ~jectory, and there is at most a 3% Cpx fractionation effect between the average Coleman samples and the other basalts from the Hobbs Lineament. While we could in principal normalize the data for Cpx control using V contents, the effects are small enough to overlook. Normalization for olivine control was achieved by incrementally adding olivine back to each composition, while maintaiLning a constant F e / M g g o = 0.30, until the Mg number was 73. This is equivalent to producing melt,; in equilibrium with a mantle olivine of Fo90 composition. This is a reasonable mantle composition; other choices could be made, but the key point is to normalize all the samples to the same Mg number. The amount of olivine required to reach Mg 73 ranged from 32 to 35% for the Coleman locality, and from 19% (37A) to 32% (47A) for the other localities (the nominal value for the mugearite was 50%, though this data is not shown in Figs. 2 aud 3). Both the olivine correction and the Cpx fractionation issue were evaluated further by carrying out liquid-line-of-descent calculations using the MELTS program (Ghiorso and Sack, 1995), starting with a 4 GPa melt from Herzberg and
Zhang (1996), cooled with variable amounts of water and K 2 0 at pressures in the 0.1-0.5 GPa range. These calculations confirm the simple approach we have taken above. In most cases, the olivine correction decreases the scatter in the data, typically generating fight curvilinear arrays or compact clusters. As we will argue later, the relative homogeneity of isotopic and trace element ratios in all of these samples argues strongly for a homogeneous mantle source composition. The trends in the olivine-corrected data are thus unlikely to be related to differences in source composition. Based on the high abundances of incompatible trace elements, the extent of melting must also be very small for these basalts, so the compositional trends seen in Figs. 2 and 3 are not related to wide ranges in the degree of melting. It is now well accepted for MORB melting that a principal control on elements such as FeO and SiO 2 is pressure of melting (e.g., Kinzler and Grove, 1992; Hirose and Kushiro, 1993). The trends shown in Figs. 2 and 3 are consistent with such a pressure-dependent melting model, though the actual pressures are likely to be much higher than for the MORB case; this is based on the need for a large residual garnet component during melting, to be discussed in Section 7. Recently, Herzberg and Zhang (1996) have reported melting experiments on a peridotite with near-primitive composition over a large range of pressures in the garnet stability field. Using their algorithms for near-solidus melt compositions, we
232
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
have plotted on Figs. 2 and 3 the compositional trajectory for these melts, in the 2-5 GPa pressure range. While our data arrays are displaced from the Herzberg and Zhang curves, the general parallelism is quite stalking--see especially the AI203-MgOFeO data in Fig. 3. In part, the displacement of the arrays may be due to the fact that their source is a fertile peridotite, whereas we show below that the Hobbs Lineament mantle source is most probably fairly depleted in elements such as CaO and AI203. We also show that the Hobbs mantle source probably contains hydrous phases, in contrast to the anhydrous experiments of Herzberg and Zhang. Considering this source difference and the difficult nature of the experiments, we feel that the general correspondence with our data is adequate, and supports a model where the principal control on our compositions is pressure (i.e., depth of melting). For the Coleman locality, the pressure estimates derived from the Herzberg-Zhang curves are ~ 4.7 GPa (FeO), ~ 4.1 GPa (MgO), > 3.6 GPa (A1203), < 6.1 GPa (CaO), and ~ 4.1 GPa (Na20). For the lowest FeO sample from the other localities (48B), the comparative pressures would be 3.6, 3.3, 3.0, 5.5 and 3.5 GPa; the pressure range exhibited by our samples is thus about 1 GPa (based on the MgO and FeO parameters). Taken at face value, these considerations suggest that the Coleman locality samples represent melt-equilibration at about 140 km depth, compared to a depth of about 110 km for sample 48B, Holmes Bluff. It might also be noted that, whereas the Coleman samples appear to have maintained their 140 km melting depth over a significant time span, melting at Holmes Bluff appears to migrate deeper with time, from ~ 100 km at 8.2 m.y. to ~ 115 km at 6.3 m.y. While this may be stretching the use of the H - Z algorithms in absolute values, we feel that the sense of the trends is certainly correct (higher FeO is higher pressure, etc.), and that the differences in indicated pressure are robust estimates (comparable to a 30 km depth difference between Coleman and Holmes). For ease of discussion we will use the 110 km and 140 km depth estimates, though we recognize that the absolute values must be tentative due to the different source compositions in their experiment and in the Hobbs case. We will make use of these melting depths later, in talking about models for WARS rifting.
Finally, to evaluate the conditions under which these magmas fractionated in the crust, we carded out MELTS modeling (Ghiorso and Sack, 1995) on samples from several of the localities, cooling them at various pressures and H 2 0 contents to look for the conditions of multiple-saturation (olivine + cpx, or in the case of sample 47A, ol + cpx + plagioclase). In several cases (37A, 50A), the melts had to be virtually dry to achieve multiple saturation at crustal pressures. For dry melts, multiple saturation was achieved at 0.41 GPa (35M), 0.7 GPa (37A), 0.28 GPa (47A), 0.37 GPa (48B) and 0.71 GPa (50A). These all represent mid-to-deep crustal levels; apparently little or no differentiation occurred in shallow magma chambers, perhaps suggesting the lack of any magma chamber for the small eruptions which make up the Hobbs Lineament.
6. Isotopic signatures One of the unusual aspects of the heavy isotope data, shown in Fig. 4, is the large range in 2°6pb//2°4pb isotopic ratios (19.50-20.69) coupled with a very restricted range in 87Sr//86Sr (0.702540.70288) and 143Nd/144Nd (0.51287-0.51301). Much of the data fall in areas of isotopic space not previously populated with data; at first glance the isotope arrays appear to be aligned between HIMU and MORB mantles. The HIMU mantle component (or at least a mantle component with fairly radiogenic Pb) is quite ubiquitous in Antarctic Cenozoic volcanism, first identified in the Balleny Plume (Hart, 1988; Lanyon et al., 1993), and also noted in Marie Byrd Land (Futa and LeMasurier, 1983), Northern Victoria Land (Hart and Kyle, 1994; Rocholl et al., 1995) and Peter I Island (Hart et al., 1995). The HIMU signature is found as well in Gondwanan pieces which were once in proximity to the Pacific coast of Antarctica (e.g., Tasmania, Ewart et al., 1988; Lanyon et al., 1993; New Zealand, Barreiro and Cooper, 1987; S.R. Hart et al., unpubl, data). To date, the 2°6pb/2°4pb value of 20.69 reported here for Coleman Nunatak is the highest reported for any of these 'HIMU' localities (and was one of the motivations for this detailed study of the Coleman area). A possible HIMU-MORB mixing explanation for
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
40.5
233
St.HELI~IA
IA
j;
40.0 39.5 R
39.0
E
U
N
I
O
N
~
~
oO
38.5 38.0 ~
37.5
P/A RIDGE
37A 48B
47A 50A 38A
36B .
18.0
18.5
19,0
19.5
20.0
.
.
20.5
.7045
.
21.0
IB
.7040,
JONES~rms~
%
PAGALU~ t ~ PRINCEEDWARD
\ ~\
~
CANARIES CAMEROONS
o0 .7030
.7025.
~
~
~C~SXON BALLENY/$COTT
.7020
.... 18.0
P/ARIDGE , .... , .... 18.5 19.0
37A 48B 47A 50A 38A 36B , .... , .... , .... 19.5 20.0 2,0.5
2,1.0
206/204 Pb Fig. 4. 20spb/204pb_ s7Sr/86 S r - 206Pb/204pb isotope correlation plots for Hobbs Lineament volcanics, in comparison with other Antarctic localities (Balleny/Scott, Peter I, and Jones Mountains), and a nearby spreading ridge ( P / A = Pacific Antarctic). Selected oceanic basalt data are also shown for comparison (both Atlantic and Pacific N-MOP, B, and OIB). Data sources are as given in Halliday et al. (1992) and Hart et al. (1992, 1995). Squares are Coleman series 35 and 46 samples, undifferentiated; X = 38A; Y = 47A; dots = 50A and 48B; triangle = 37A; Z = 36B mugearite; sample numbers appear at bottom of each panel, directly underneath symbol.
the Hobbs array can be tested in three-isotope space, using the mantle tetrahedron plot of Hart et al. (1992); see Fig. 5A. Here again the Hobbs array projects neatly to MORB, and in particular to the MORB from the nearest segment of the Pacific/Antarctic Ridge (the projection is excellent from other viewing angles as well). This viewpoint also shows that the ,'u'ray does not exactly back-project to HIMU, but to a location slightly up the HIMU-EM 1 edge.
When compared with the data arrays for Scott/Balleny, Peter I, Jones Mountains, and Mt. Early/Sheridan Bluff, all show a tendency to converge in the south-central part of the tetrahedron, i.e., the other arrays are clearly not mixing toward MORB. This is the same characteristic shown for the ocean island basalt (OIB) data set as a whole (Hart et al., 1992), with the convergence zone labeled FOZO (Focus Zone). To examine this phenomenon more closely, we show in panel B of Fig. 5 an enlargement
234
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
Nd•6/4 ~ ~ / f r EM1
Early/Sherid~
P/A Ridge Hobbs
DMM I
HUDSON
HIMU
SIPLE
~ID
of the Hobbs array portion of the tetrahedron, with the addition of some isotopic data from other regions of Marie Byrd Land, for cases where we have more than one sample from a given locality (see Table 3 for the isotope data on these samples). There is a remarkable fan-shaped nature to these arrays, with a convergence or focus just at the left end of the Hobbs array. This suggests that the Hobbs array is not mixing to MORB, but to a more central component common to many other MBL arrays (and very close to the FOZO component of Hart et al., 1992; Hauri et al., 1994a). We note that there is no evidence from nearby ridges or ocean crust for this component, only MORB with normal 2°6pb/2°4pb ratios ( < 19.0) have been reported (Ferguson and Klein, 1993; Pyle et al., 1995). This FOZO component will figure prominently in our WARS model, to be discussed below.
FOSDICK
7. Trace element patterns
Fig. 5. Three-dimensional S7Sr/S6Sr, 143Nd/144Nd and 2°6pb/2°4pb isotope correlation plots. The end-member mantle compositions of Zindler and Hart (1986) define a tetrahedron (Hart et al., 1992), and the data are projected onto the DMMEM1-HIMU base plane of this tetrahedron. (A) View of whole mantle tetrahedron, showing the Hobbs Lineament data (Table 2) compared to arrays from Jones Mountains and Peter I Island (Hart et al., 1995), Scott/Balleny islands (Hart, 1988), Mount Early/Sheridan Bluff (S.R. Hart et al., unpubl, data) and MORB from the Pacific-Antarctic Ridge (P/A, Ferguson and Klein, 1993). This MORB locality is only 650 km northeast of Scott Island. Projections of the three orthogonal isotope axes are shown for reference; end-member compositions are as in Hart et al. (1992), except FOZO has been moved to slightly lower 143Nd/144Nd. (B) Enlargement of the south-central region of the tetrahedron comparing the Hobbs array with several other localities from Made Byrd Land (Table 3). The arrays converge on the FOZO region of Hart et al. (1992); all except Jones Mountains lie very close to the DMM-EM1-HIMU base plane (the Jones Mountains trend, panel A, intersects the base plane at FOZO).
The trace element data are shown in Fig. 6 as primitive-mantle-normalized spidergrams. Because all of the Coleman 35 and 46 locality samples are so similar, we show the average of all seven; for compadson, we show the lowest spidergram for the Hobbs group, 48B, Holmes Bluff. All the other samples lie intermediate to these two spidergrams. All show the typical negative K and Pb anomalies of OIB; these anomalies, coupled with the marked depletions in Cs and Rb, argue against any significant role for crustal contamination of the lavas (as of course do the very low 87Sr/86Sr isotope ratios). Hole and LeMasurier (1994) also provided compelling arguments against significant crustal or mantie lithospheric components in alkaline volcanics from Made Byrd Land and the Antarctic Peninsula. The principal difference between the bounding spidergrams in panel A of Fig. 6 is the more pronounced Z r - H f dip in the Coleman average, and the larger Pb anomaly for Coleman, suggestive of a 'buffering' of Pb in the Hobbs volcanies by a Pbcompatible residual phase. The overall nature of the spidergrams is consistent with melting to variable degrees of an approximately homogeneous source. As noted before, the Sr and Nd isotopes support this notion, in the sense that only small R b / S r or S m / N d variations are needed to
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
235
Table 3 Isotopic data for selected Marie Byrd Land localities Sample Nr.
Locality
87Sr/86 Sr
E-27-34c 66-D-91 66-D-30 W83-5C W83-1 62 85-32B 24-1-A 9-1-C 20-2B
Scott Island Fosdick Mountains Fosdick Mountains Mt, Siple Mt, Siple Mt. Mmphy Mt. Mmphy Hudson Mountains Hudson Mountains Hudson Mountains
143Nd / 144Nd
0.512989 0.512859 0.512948 0.512847 0.512792 0.512893 0.512839 0.512980 0.512859 0.512928
0.702849 0.703338 0.703106 0.703487 0.703723 0.702917 0.703165 0.702904 0.703366 0.703026
1000
~B, HOLMES•
,Iktlll
]
CsRbBa Th UNbTa K l.a Ce PbNd Sr Zr HfSmEuTi GdTbDyHo y Er Yb Lu
1000
'B'
[I
1 CqRb Ba Th U NbTa
il '1:
La Ce Pb Nd Sr Zr FLfSmEu Ti GdTbDyHo Y Er Yb Lu
ELEMENT
Fig. 6. Trace element spidergrams for Hobbs Lineament volcanic rocks. (A) The average of seven Coleman Nunatak samples is compared to the most depleted sample, 48B, Holmes Bluff. (B) The average of the Coleman Nunatak samples compared to the oceanic HIMU end-mem~r. The latter is based on seven basalts from Mangaia and Tubuaii (S.R. Hart and E.H. Hauri, unpubl. data). The Coleman average is as in panel A, but adjusted uniformly down by 21%, which is the amount of olivine which would need to be added to bring the Mg number of Coleman (51.8) up to that observed in the HIMU suite (67.5). Primitive upper mantle (PUM) normalizing factors are from McDonough and Sun (1995); for the refractory elements, these are ~ 10% higher than the values of Hart and Zindler (1986).
2°6pb / 204pb
2°7pb / 204pb
19.638 19.054 19.224 19.669 19.753 19.764 20.010 19.406 19.138 19.333
15.601 15.580 15.607 15.665 15.725 15.667 15.691 15.623 15.603 15.692
2°8pb / 204pb
:
39.197 38.489 38.595 39.251 39.494 39.195 39.463 38.840 38.669 38.984
generate the small differences observed in 87Sr/86Sr and raNd/144Nd. Similarly, while the observed Pb isotope range is large (e.g., 19.5-20.7), the time-integrated differences in U / P b required to produce this in a few billion years is only 25%. We may test the concept of a single mantle source, as well as derive information about the source composition and bulk partition coefficients, using the modeling techniques for partial melting pioneered by Treuil and Joron (1975) and Minster and All~gre (1978). Numerous refinements and successful applications of these techniques have been reported subsequently (Albar~de, 1983; Hofmann and Feigenson, 1983; Albar~de and Tamagnan, 1988; Maalee, 1994; Zou and Zindler, 1996). We will adopt the formuladon given by Sims and DePaolo (1997). For modal batch partial melting: C
1
Co = F + o ( 1 - F )
(1)
where C is the concentration of a trace element in a melt fraction F; CO is the initial source concentration, D is the bulk partition coefficient. When two trace elements A and B are plotted against each other as reciprocals ( x = I / A , y = I / B ) , the resulting correlation is linear, with a slope and intercept defined as: slope
CoA(1 --DB) CoB(1 _DA)
( D s --DA) intercept --- Con ( 1 - DA)
(2)
(3)
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
236
For non-modal batch melting, the array is still linear, but the slope and intercept contain an additional parameter P, which is the bulk partition coefficient of the phases in the proportion in which they are melting:
co (1 slope = Con(1 _ eg)
intercept =
(4)
DB(1 --PA) --Oh(1 - - e a ) CoB(1 - PA)
mildly incompatible element (Gd). For La and Nb, since D << 1, the slope, from Eq. (2) gives the source ratio of N b / L a (here = 1.383); the intercept from Eq. (3) is proportional to the difference in partition coefficient between Nb and La. Because the intercept is positive in this case, DLa > DNb (though the difference is quite small). For Th and Gd, again DTh << 1 but DGa may not be small, so the slope from Eq. (2) gives:
(5)
Co
0.2966 = ~oGd(1 -- DGd) For fractional melting, the equations are not linear, but may nevertheless by treated as linear in limited situations, as we discuss below. Two examples of this approach are illustrated, for the Hobbs data, in Fig. 7; panel A shows two fairly incompatible elements (La and Nb) and panel B shows one highly incompatible element (Th) and one 0.040 : 0.035 ".: 0.030 0.025 ~ 0.020' 0.0150.010' 0"005' ~.,.00181 :~ .00088 0.000
,*,,l.,,i
i,,.,i,.,,
~" ii i f , l , ,
.0(30 .005 .010 .015 . 0 2 0 .025
1/Nb 0.20-' 0.15'
.2966* .0264
0.10' 0.05'
•
and the intercept, from Eq. (3), gives: DGd
0.0756 = - -
Cod
Thus fixing one parameter, say Co~d, allows one to calculate DGd and C~. And from a Th-Nb plot, the additional source concentration of Nb could be determined (and from this, La from Fig. 7A). In principle, for modal melting, by fixing only a single source element concentration, complete spidergrams for both the source concentrations and bulk partition coefficients can be derived, subject only to uncertainties relating to the actual scatter of data in plots such as Fig. 7. From Eqs. (2) and (3), it is clear that the right-hand end of the spidergram is well constrained for both Co and D; the highly incompatible elements are well constrained for CO, but substantial error magnification can occur in the D's, due to the (D B - Da) term. For non-modal melting, some ambiguity arises for the most compatible elements, but by judicious choice of elements the rest of the spidergram is still well constrained. For example, for a highly incompatible element A (where DA and PA << 1), relative to a mildly incompatible element B (D B < 1), Eqs. (4) and (5) become:
•
Co (1 - e . ) 0.00
....
I,,,, I,,,, I,,'" l, ,, ,I, , , , | , , , , I , , , ' m '~'~ ~
.00 .05 .10 .15 .20 .25 .30 .35 .40 .45 1/Th
Fig. 7. Reciprocal L a - N b and G d - T h plots, with York (1966) two-error regression lines plotted through the Hobbs Lineament data. Input uncertainties given as + 5% for each element, with concentrations from Table 2 corrected for olivine fractionation (see text). Output errors for slopes and intercepts are absolute values, quoted at + 2 t r , and are independent of the input errors, to first order (Brooks et al., 1972). The mugearite 36B is not plotted.
slope
CoB
DB intercept = C°a While a global inversion of all the data could be performed, as done by Minster and All~gre (1978), it is difficult to properly weight the 'geologic scatter' on the input errors; the output errors can be signifi-
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
cantly impacted by this aspect (All~gre et al., 1983a,b). Instead, we have iteratively worked our way through the whole data set with the two-error regression treatment of York (1966), which generates error estimates for :slope and intercept which are dominated by the actual data scatter and are relatively insensitive to the input weighting factors (Brooks et al., 1972). This approach capitalizes on the best-constrained data arrays. The results of this process are listed in Table 4 and shown in Fig. 8,~ for both the source concentration (panel A, top, normalized to primitive upper mantle: PUM) and the bulk partition coefficients (panel B, bottom), tks mentioned above, all of the data floats on the choice of one C O value; due to the relative constancy of the heavy REE in peridotltes, we selected Dy (the heavier REEs in our data set, and Yb and Lu in particular, are somewhat decoupied from more incompatible elements and form poor data arrays in plots such as Fig. 7). First, the
237
C s R b k Th U NbTa K La CeNd $r Zr HfSmEuTi GdTbDyHo Y Er Yb Lu 1
Table 4 Source concentrationsand partitioncoefficients Element
C O(ppm)
Co, P U M
20" ( % )
Cs Rb Ba Th U Nb Ta K La Ce
0.00314 0.486 7.08 0.0768 0.0208 1.199 0.0616 258.1 0.864 1.703
0.150 0.gl0 1.(173 0.966 1.023 1.822 1.666 1.075 1.334 1.017
6.9 4.1 3.1 0.4 0.7 1.1 0.8 4.2 1.0 0.5
Pb
.
Nd
0.842
Sr Zr Hf Sm Eu Ti Gd Tb Dy Ho Y Er Yb Lu
.
19.12 8.94 0.1757 0.2163 0.0822 384.9 0.2544 0.0403 0.337 0.0756 1.642 0.255 0.1355 0.01582
.
.
DO
-0.00359 0.00133 0.00167 0.00126 0.00129 0.00032 -0.00143 0.0165 0.00183 0.00245
20" ( % )
26.1 52.9 40.6 6.8 7.4 81.0 8.6 10.1 15.8 6.0
.
0.673
-
0.00332
-
0.960 0.851 0.621 0.533 0.534 0.319 0.,168 0.407 0.5 0.:507 0.1382 0.582 0.'.307 0.234
2.5
0.00816 0.0250 0.0183 0.00879 0.01471 0.00702 0.0206 0.0244 0.0486 0.0644 0.0446 0.0813 0.0526 0.0366
8.2
2.2 2.7 2.3 4.6 2.1 5.8 -
4.2 13.2 8.2 18.0 4.6 3.4 1.9 2.4 20.7 -
C.sRbBa Th U NbTa K La C.eNdSr Zr HfSm EuTi C-dTbDyHo YEr Yb Lu ELEMENT F i g . 8. T r a c e e l e m e n t ( p a n e l A ) a n d b u l k p a r t i t i o n c o e f f i c i e n t (panel B) spidergrams for the Hobbs Lineament mantle source
(inversionbased on Dyo = 0.500 PUM; data is in Table 4). PUM normalizingfactors are from McDonoughand Sun (1995). Cs and Ta points are not plotted in panel B because the D values ended up being negative (see Table 4). No line is drawn connectingEr and Yb because of the poor correlationsbetween Yb and Lu and the rest of the spidergram.Dashed lines have been drawn between the various REE to highlight the anomalies of intervening elements. Error bars are not shown but are very small for the elements in panel A; some of the elementsin panel B have sizable uncertainties(Rb, Ba, Nb).
enriched nature of the Hobbs source mantle spidergram (Fig. 8A) is a robust feature independent of the choice of any C o (for example, L a / S m is virtually constant for Dy input choices ranging from 0.5 × PUM to 2 × PUM). Second, as has been known since the classic work of Frey and Green (1974), peridotites with enriched LREE patterns are typically depleted in heavy REE, and also less fertile (in CaO, A I 2 0 3, etc.) than PUM. Utilizing the peridotite
238
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
database of McDonough (1990) (195 spinel and 57 garnet peridotite xenoliths), for samples with L a / S m > 2.5 × PUM ( L a / S m of the source in Fig. 8A is 2.5), only a very few samples (5) have heavy REE > 1 × PUM, and the average is about 0.5 × PUM. For the 68 massif peridotites in the McDonough data base, none have L a / S m > 2 × PUM; the average heavy REE content is ,-, 0.62 × PUM. We believe that our choice of Dy = 0.5 × PUM for Fig. 8 is unlikely to be uncertain by more than ___50%. This uncertainty will translate almost directly into a + 50% variation in the source spidergram and the bulk partition coefficient spidergram. The various anomalies on the spidergrams will not be materially affected. Most of the features of the Fig. 8 spidergrams appear reasonable. First, the source pattern is consistent with a mantle which was once depleted (positive slope on left end), as suggested by the Sr and Nd isotopes, but has subsequently been enriched, to account for the high concentrations in the Ba to La portion of the spidergrams of the erupted basalts (Fig. 6). The bulk partition coefficient spidergram is very steep, as would be consistent with melting in the garnet facies (depths of melting inferred earlier to be ~ 110-140 km). There is a marked negative Ti anomaly, and marked positive Z r - H f anomalies in the partition coefficient spidergram, again, typical signatures of garnet melting (Hauri et al., 1994b; Salters and Longhi, 1997). The absolute values of the heavy REE partition coefficients are quite low, as would be expected for a depleted (harzburgitic) source, with low modal garnet and Cpx abundances. Using the garnet/melt Er partition coefficients of Salters and Longhi (1997) or Hauri et al. (1994b), which have values of 3.5-4.0, no more than ~ 2% garnet can be present during melting. The negative source anomalies at Ti and Y are compensated by negative partition coefficient anomalies, resulting in almost anomaly-free Ti and Y in the melts (see Fig. 6). The positive Zr-Hf anomalies in the source are over-compensated by the large positive partition coefficient anomalies for Zr and Hf, leading to mild negative anomalies in the erupted basalts. The negative K anomaly in the source is exacerbated by a large positive partition coefficient anomaly for K, leading to melts with marked negative K anomalies. A positive source anomaly for Sr is largely offset by
a positive partition coefficient anomaly, leading to melts with only a minor positive Sr anomaly. There are no Eu anomalies in either the source or bulk partition coefficients, and there are no Eu anomalies in the melts (consistent with high-pressure melting). There is a slight bowl-shape to the REE source spidergram (ignoring Lu and Yb) as is frequently found in xenolith suites where an enrichment event is superimposed on a depleted source. The REE slope of the partition coefficient spidergram is not as steep as would be expected for pure garnet control, so it is likely that some Cpx is present as well. Cpx cannot be the dominant phase melting, however, as it carries a negative Zr partitioning anomaly (Hart and Dunn, 1993), which will tend to cancel the positive Zr anomaly in garnet (and in Fig. 8B). The positive K anomaly is indicative of minor amounts of phlogopite or amphibole; the lack of enhanced Rb and Ba partition coefficients argues for amphibole as opposed to phlogopite. It would require 1-2% amphibole to generate the K anomaly in Fig. 8B, for the range of amphibole/melt partition coefficients cited by Green (1994). The origin of the mild positive Sr anomaly in Fig. 8B is less clear, though it may also be attributable to amphibole (Green, 1994). The negative Y anomaly is unexplained, as neither garnet, Cpx nor amphibole appear to fractionate Y from the heavy REE (Hart and Dunn, 1993; Green, 1994; Salters and Longhi, 1997). We ran a series of numerical tests with an aggregated fractional melting model, using the source composition and bulk partition coefficients of Table 4, to see how closely the recovered slopes and intercepts agreed with those derived by assuming a simple batch modal melting model. The data are no longer linear on 1 / A - 1 / B plots, but for our particular data set the degree of curvature is very small and would not be noticeable given the scatter of data on plots such as Fig. 7. A linear regression (of the non-linear data) produced intercept values which were generally very close (within 10%) to those calculated using Eq. (3). The recovered slopes, however, typically underestimated slopes calculated from Eq. (2) by 5-40%. Thus, the bulk partition coefficient spidergram (Fig. 8B) will be virtually unaffected, were the actual melting process for the Hobbs Lineament to be fractional as opposed to batch melting. The general appearance of the source mantle
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
spidergram (Fig. 8A) will also be little changed, though it will be rotated clockwise (about the fixed Dy point), with the highly incompatible elements becoming more enriched by up to 40%. Given the data in Fig. 8, it is straightforward to calculate F, the melt fraction, for any of the Hobbs Lineament lavas. For the average Coleman Nunatak lavas, after olivine-correction, F--1.6%; for the lower bounding spidergram, sample 48B (Fig. 6), F = 3.2%. These melt fractions are directly dependent on our estimate for DY0; for example, if Dy 0 were as low as 0.25 × PUM, then the melt fractions would become 0.8% and 1.6%, respectively. In any event, the two bounding spidergrams in Fig. 6 can be derived from a homogeneous source with a melt fraction for 48B that is twice that for Coleman. The HIMU 'end' of the Hobbs array, represented by the average Coleman basalt in Fig. 6A, may be compared with the spidergram for 'classic' HIMU as defined for OIB. This is done in Fig. 6B, where the HIMU pattern is that of seven basalts from the islands of Mangaia and Tubuaii; the Coleman data have been adjusted down by 21%, to account for the difference in observed Mg number (Mg 51.8 for Coleman, Mg 67.5 for the HIMU average). The patterns are amazingly similar, and more or less parallel throughout; both have the negative K and Pb anomalies typical of OIBs in general, both have mild Zr-Hf anomalies arid again the Pb anomaly looks 'pinned' as it was i~LFig. 6A. The only significant difference seems to be in the Ba to Nb region, where the HIMU pattern is relatively smooth in contrast to the Coleman pattern which has a clear depletion of Th and U relative to Ba and Nb. This leads to a high B a / T h ratio in Coleman (av. 88.3) relative to HIMU (52.4). The exact significance of this difference is unclear. At face value, Ba and Th partitioning data for garnet and cpx suggests a much lower K o for B a / T h in cpx than in garnet (Hart and Dunn, 1993; Hauri et al., 1994b), so a smaller residual garnet component during Coleman melting would tend to lead to higher B a / T h ratios (note that Beattie, 1993 does have garnet B a / T h < Cpx Ba/Th). However, the Coleman and HIMU spidergrams are very nearly parallel throughout the whole REE range, suggesting that the modal gamet/cpx ratio during melting was very similar for both HI/vlU and Coleman. It is possible
239
that one can appeal to minor phases to fine-tune these elements; for example, Class and Goldstein (1997) have proposed that phlogopite and amphibole from mantle lithosphere may modulate trace element patterns in some OIB. We have suggested above that amphibole is a likely residual phase during Hobbs Lineament melting. Ba is compatible in phlogopite; residual phlogopite in the HIMU source could lead to the low B a / T h in HIMU relative to Coleman. However, K and Rb are also compatible in phlogopite, and it is not clear the B a / T h can be fractionated without upsetting Ba/Rb, etc.
8. Elemental and isotopic correlations First, as implicit in the previous section, many of the incompatible trace elements in the Hobbs suite show very good correlations with each other. These trace elements are found also to be well correlated with major elements such as FeO, particularly after the olivine normalization to Mg 73. There are also clear correlations between major and trace elements and isotopes (e.g., F e o - l g a N d / 144Nd, 2°apb/2°4pb-Nb, etc.). A few of these various correlations are shown in Fig. 9. Panels A and E show that trace element concentrations (e.g., Nb and Ce), which are argued above to be controlled by degree of melting, are correlated with major elements such as FeO and MgO. Since these major elements were argued in Section 5 to be pressure (depth) controlled, this leads to a correlation between extent of melting and depth, with the lowest extents of melting at the greatest depth (Coleman group, F ~ 1.6%, Z ~ 140 km). Panels C, D and F show that the St, Nd and Pb isotopes are also related to major elements, and therefore depth. In this case, the HIMU (Coleman) end of the isotope array (see Fig. 5) is correlated with the greatest depth of melting, and the FOZO (Holmes) end of the isotope array is from the shallower depths of melting. It might be noted that, of the other MBL localities plotted on Fig. 5B which show trends toward FOZO, the Balleny/Scott, Murphy and Siple arrays also show decreasing FeO (shallower depths) in the FOZO direction; Hudson and Fosdick show little or no change in FeO along the arrays.
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Finally, Fig. 9B demonstrates one of the isotopetrace element correlations which is implicit from the above correlations: the HIMU (Coleman) component represents the highest trace element concentrations, i.e., the lowest degrees of melting. This lower degree of melting at Coleman leads to slightly elevated R b / S r and N d / S m ratios (16% and 11%, respectively), which then produce Sr and Nd pseudo-iso-
chron correlations (87Sr/86Sr versus Rb/Sr, etc.) for the Hobbs Lineament volcanics (since Coleman has higher 87Sr/a6Sr and lower 143Nd/144Nd than the rest of the Hobbs localities). There are good pseudo-isochrons for the U / P b and T h / P b systems also, though in this case normal melting processes cannot produce the large range observed for U / P b and T h / P b (both a factor of 2.3). If Pb is strongly
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
buffered as a compatible element by some trace phase (sulfide?), this large range in parent/daughter ratio could be produced from a homogeneous source. Alternatively, and we feel, more likely, the Pb concentration in the HIMU and FOZO sources is different by a factor of ~Iwo to begin with (and is the fundamental reason for the large Pb isotope differences between the HIMU and FOZO sources). During the latest melting event, Pb has behaved in a normal magmatic way (i.e., like Ce), and the original factor of two deficit in Pb content of the HIMU source has been compensated for by the lower degree of melting. The result is comparable Pb contents in all the Hobbs volcanics. In this case, the U / P b and T h / P b pseudo-isochrons are in fact real mantleisochrons, and their approximate age of 230 m.y. may relate to the time of Pb depletion in the Coleman (HIMU) end-member. Thus the HIMU signature is not related to a U + Th enrichment process as frequently asserted, but rather to a Pb depletion process. Peucker-Ehrenbrink et al. (1994) have argued that a significant fraction of the Pb in the ocean crust is transferred into metalliferous sediments during ridge crest hydrothermal processes. If HIMU mantle is recycled ocean crust, as originally postulated by Hofmann and White (1982), then the Pb depletion we hypothesize at Coleman may relate to such a ridge crest process. While speculative, the pseudo-isochron age of 230 m.y. would point to fairly recent crust recycling, then, and would seem to obviate a deep-mantle transport cycle for this particular HIMU source. In addition to these numerous major-trace-isotope correlations, there are also important correlations with age of eruption of the volcanics; some of these are shown in Fig. 10. In all cases, since the basalts from Coleman show virtually no chemical change over their 9 m.y. eruption span, horizontal arrays are formed for the Coleman suite. In contrast, the rest of the Hobbs group show decreases in A1203 (and increases in FeO, MgO), increases in Nb (and other LIL-elements), and increases in 2°6pb/:°4Pb with younging of eruption age. In other words, Coleman erupts deep, low-degree HIMU-type melts through time, whereas volcanism elsewhere on the lineament starts shallow, with higher-degree FOZO-type melts. Through time, the melting becomes deeper and smaller in degree, thus approaching the Coleman
241
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suite in all aspects. To summarize these relationships in terms of our estimates for depth and extent of melting: (1) Coleman group: 1.6% melting at 140 km depth; HIMU in character; constant from 11.7 m.y. to 2.3 m.y.
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S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
(2) Cousins to Holmes group: starts at 10 m.y. with 3.2% melting at 110 km depth; FOZO in character; through time the melting gets deeper and the HIMU component increases, F decreases.
9. Implications for the WARS One of the central issues regarding the West Antarctic rift system is the role, if any, played by mantle plumes. The geochemical case for plumes rests on the chemical similarity between WARS volcanism and that of oceanic islands, where plumes are the reigning paradigm. This similarity has been argued in numerous papers (Kyle et al., 1992; Lanyon et al., 1993; Hole and LeMasurier, 1994; Hart et al., 1995), and the essence can be seen in the trace element patterns (e.g., Fig. 6) and isotopic signatures (e.g., Fig. 4), especially for the HIMU end of WARS volcanism. There are several garden-variety oceanic plumes in the WARS vicinity also, with appropriate chemical signatures to serve as analogues for some of the WARS volcanism. The Balleny Islands (see location B, Fig. 1A) have been in most plume catalogues for years (Morgan, 1972; Vogt, 1981), and, as shown in Figs. 4 and 5, the Balleny volcanics have isotopic characteristics appropriate for the 'FOZO'end of the Hobbs Lineament volcanism, as well as other volcanism in Marie Byrd Land (see Fig. 5) and Northern Victoria Land (see fig. 1 in Hart and Kyle, 1994). Scott Island (location S, Fig. 1A) has never been formally designated a plume, but its intraoceanic setting, young volcanism (WiSrner and Orsi, 1992), gravity signature (Marks and McAdoo, 1992), and striking isotopic similarity to the Balleny plume strongly implicate a plume attribution. And finally, Hart et al. (1995) argued for a plume origin for the young volcanism on Peter I Island (see location P, Fig. 1A). Note, however, that while these 'plumes' provide a FOZO-type mantle component, they do not strictly provide evidence for the HIMU component as seen at Coleman Nunatak (see Fig. 5). Aside from the geochemical case, a variety of strong arguments have been made for a large plume underlying Made Byrd Land. These include the doming and associated patterns o f volcanism (LeMasurier and Rex, 1989; Kyle et al., 1992; LeMasurier and Landis, 1996), the association with numerous geo-
physical manifestations of rifting (Behrendt et al., 1992), and the well-defined relationship between domal uplift of MBL and the volcanic history over the past 30 m.y. (LeMasurier and Landis, 1996). All of this seems quite persuasive as long as a single plume source is not argued for all of the Cenozoic volcanism in the WARS. As shown by Hart and Kyle (1994) and Rocholl et al. (1995), the volcanism in Northern Victoria Land exhibits a significant HIMU component, and to include this as part of the same plume as MBL requires a plume more than 3000 km in dimension (see fig. 2 in Behrendt et al., 1992). More importantly, as noted by Weaver et al. (1994) and by Blusztajn and Hart (1995), Mesozoic and Cenozoic HIMU-type volcanism is also present in Tasmania, New Zealand and the Campbell Plateau. These localities clearly cannot be embraced by a single plume presently centered under MBL. Short of postulating separate HIMU-type plumes for each area, we note that these areas were all contiguous prior to the breakup of Gondwana (see reconstruction in Lawver et al., 1992). As one resolution to an awkward multi-plume scenario, we suggest that a single large HIMU plume head arrived prior to breakup and underplated the lithosphere throughout this broad area. It is tempting to associate breakup with the arrival of this plume head, but as noted by Behrendt et al. (1992) and LeMasurier and Landis (1996), there is no large-scale volcanism or uplift related to the breakup event at about 85 m.y. The major volcanic event to which a plume head could be related is the Ferrar Group Large Igneous Province (Brooks and Hart, 1978), which, at ~ 185 m.y. (Encarnaci6n et al., 1996), significantly preceded NZ-MBL breakup (Hergt et al., 1991). Alternatively, because the area to be underplated by a single plume head is still very large ( > 5000 km), one could postulate a plume head which had thermally entrained such large fractions of mantle on the ascent that its thermal anomaly (and buoyancy flux) was highly dispersed, and in fact insufficient to cause plume melting, especially if its rise was terminated at depth by a thick Gondwanan continental lithosphere. Further flattening and radial expansion would provide a young proto-lithosphere under a vast region which could, after a suitable 'incubation' period (Kent et al., 1992), provide a fossil-HIMU mantle
S.R. Hart et al. / Chemical Geology 139 (1997) 223-248
source during subsequent rifting events (Halliday et al., 1990; Stein and Hofmann, 1992). This underplating/rifting scenario, origirtally proposed by Brooks et al. (1976) as a general phenomenon, can also be accompanied by metasomatism and enrichment of overlying mantle. (We concluded earlier that the Hobbs mantle source is both enriched and hydrous.) Thus the original plume need not have been HIMU in character but could have grown in the HIMU character subsequent to underplating, as postulated for the Cameroons by Halliday et al. (1990) and for the Antarctic Peninsula alkalic basalts by Hole et al. (1993). By postulating that the HIMU mantle component was emplaced prior to breakup, we can focus back on the Cenozoic volcanism in the MBL region. While the high 2°6pb//2°4pb ratios make the HIMU component particularly conspicuous, the FOZO component is probably more ubiquitous throughout the WARS, and potentially more important in understanding the evolution of WARS volcanism. Part of this is a nomenclature problem; in the past, volcanics with 2°6pb//2°4pb ratios of 19.5 were identified with a HIMU component. Based on the evidence of Fig. 5, we propose that only 2°6pb//2°4pb ratios greater than 20.5 be assigned HIMU 'status' (the oceanic HIMU end-member has a 2°6pb/204pb of 21.8; Hauri and Hart, 1993). The great bulk of Cenozoic volcanism in the west Antarctic region in fact has 2°rpb//2°4pb ratios in the 19.3-19.8 range, and this is more appropriately identified with the FOZO end-member. This is the dominant component in the three 'real' plumes in the area (Scott, Balleny and Peter I), and is therefore central to the issue of whether the WARS volcanism is plume-related, as argued by Kyle et al. (1992), Behrendt et al. (1992), Hole and LeMasurier, 1994, and LeMasurier and Landis (1996). Based on the isotopic arrays in Fig. 5, it would appear that much of the MBL volcanism contains a FOZO component, and therefore is consistent with some 'plume-connection'. For help in understanding this connection, we return to the evidence derived from the Hobbs Lineament volcanism: melting of the FOZO mantle is shallower than that of the HIMU mantle (110 km versus 140 km), represents a larger degree of melting (3.2% versus 1.6%), and migrates to the deeper
243
HIMU regions during the period from 10 m.y. to 5 m.y. Note, however, that these two mantle sources are erupting simultaneously during the 10-12 m.y. period (Fig. 10), from almost contiguous locations on the Hobbs Lineament (Coleman Nunatak and Patton Bluff are only l0 km apart). FOZO eruptions then commence at the far northern end of the lineament at 8.2 m.y. (Holmes Bluff, 48A). If a single large plume underlies all of MBL, how do we explain coeval volcanism at contiguous localities from different mantle sources, of very disparate depths? One solution would be to assert a heterogeneous two-component plume, with different solidus temperatures for the two components. For example, veined-mantle models (Wood, 1979; Zindler et al., 1984; All~gre et al., 1984; Hirschmann and Stolper, 1996) propose that pyroxenite or eclogite layers in a lherzolite might begin melting 100°-150°C lower in temperature than the lherzolite host; this would translate to a 25-40 km depth difference (for a difference in solidus slope and adiabat slope of 4°/km). If the pyroxenite layers are HIMU in character, and the lherzolite is FOZO, this would plausibly explain the deeper HIMU melts at Coleman relative to the shallower FOZO melts at Patton Bluff. With time, the melting gets deeper and the FOZO component becomes less dominant all along the lineament, perhaps because the upper surface of the plume becomes cooled upon impingement with the lithosphere. In this scenario, however, part of the logic unravels, as the depth differences we calculated were based on major element constraints from lherzolite melting experiments (Herzberg and Zhang, 1996). Similar melting experiments on marie (i.e., eclogitic) compositions (Yasuda and Fujii, 1994) show little or no dependence of melt composition on pressure, in the range 3-14 GPa. Also, the Yasuda and Fujii melt compositions are most similar to the FOZO compositions from Patton Bluff (37A) and Holmes Bluff (48B), suggesting that the melts of HIMU affinity (Coleman) derive from lherzolite, not pyroxenite. This argument is not firm, however, because highpressure melt compositions in mafic systems are quite dependent on the composition of the starting material. Yasuda and Fujii (1994) used an N-MORB starting composition, which at pressure is a quartzeclogite; a lower-silica starting material might conceivably lead to melt compositions more like the
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Coleman basalts. Also, in this model, it becomes difficult to explain our apparent success in Section 7 in modeling the full trace element data set in terms of a single uniform source. Surely the trace element partitioning from lherzolite and pyroxenite sources will look radically different! In any event, this heterogeneous plume scenario resurrects the implausibility problem of having HIMU-like volcanism occurring in many areas now well removed from one another (Tasmania-MBL). Alternatively, in a plume scenario, one could resort to major 'topography' at the base of the lithosphere (Saunders et al., 1992), to fit the constraints posed by the Hobbs Lineament data. If WARS rifting removed a significant thickness of a fossil HIMU lithosphere (30 km layer over a lateral distance of 10 km), then a FOZO plume could rise to ~ 110 km and begin melting, whereas the deeper adjacent lithosphere would melt to a lower degree during minor adiabatic uplifts associated with the rifting. Once on the solidus, the small extents of melting proposed here (1.6-3.2%) can be produced by only 3 to 5 km of adiabatic decompression (approximately 0.6% melt per km of uplift, using the preferred melting parameters of Hart, 1993). Finally, we explore a scenario in which no currently active plumes are needed. This accepts the relative depths of the FOZO and HIMU mantles (110 km versus 140 km) and locates both as part of a layered fossil plume lithosphere, a HIMU layer underlying a FOZO layer. This layering becomes a natural consequence of a plume-head model, if we adopt the Hart et al. (1992) explanation for the FOZO component. They showed that many oceanic plumes form elongated arrays in 3-D isotopic space, which converged on a region named FOZO (Focus Zone). FOZO was interpreted as lower mantle which was thermally and viscously entrained (coupled) to a plume arising from the core-mantle boundary layer. The dynamics of this entrainment were treated for steady-state plume tails by Hauri et al. (1994a); the case for plume heads may be similar (Griffiths and Campbell, 1990), in that the outermost sheath of the plume will be the entrained material. In the Hart et al. (1992) model, this entrained material is identified with FOZO. Therefore, a plume head comprised initially of HIMU mantle will arrive at the base of the lithosphere accompanied by a sheath of FOZO
attached to it; on underplafing and spreading, this proto-lithosphere will be a FOZO layer underlain by a HIMU layer. We do not envision that any melting is necessary during this underplafing event. If melting did occur, it would be dominated by the HIMU plume component, as Farnetani and Richards (1995) have demonstrated that plume head melts will be dominated by the original plume material, not the entrained material. This model is then beset by the same melting problem as the active plume model. As entrained material, the FOZO mantle is initially cooler than the HIMU mantle. How do we produce melting at 110 km depth (FOZO) right adjacent to melting at 140 km depth (HIMU)? As above, one solution is to appeal to a lower solidus temperature for the HIMU component (as in an eclogitic or pyroxenitic source scenario). Alternatively, and perhaps realistically, the temperature gradient in the fossil layering, inherited from the plume head, may be sub-parallel to a lherzolite solidus. The entrained material is clearly cooler than the plume interior. The actual thermal gradient will depend on radial distance from the plume axis, but is in the range of 2°/km (on wings) to > 10°/km (on axis) (Fametani and Richards, 1995). These gradients are approximately the same as the solidus slope for lherzolite in the 100-140 km depth range (3.4-4.5°C/km; Zhang and Herzberg, 1994). During subsequent Cenozoic tiffing, adiabatic decompression of this FOZO and HIMU layered 'package' could produce coeval melting over a substanfial depth range. With respect to lateral localization of the melting, we follow the model of King and Anderson (1995) and suggest that this melting may be a natural consequence of the lithospheric asymmetry left behind after breakup. Small-scale convection driven by differences in heat flux and lithospheric thickness (for example, at a rifted margin; Mutter et al., 1988), will drive an upwelling mantle flow from under the thick lithosphere to regions of thinner lithosphere. While extension is still required in this model, abnormally hot mantle (e.g., plume-type) is not. While the total amount of Cenozoic extension in the WARS is not large (Lawver and Gahagan, 1994), it may be sufficient to enhance small-scale convection if it is focused or localized in the regions of lithospheric asymmetry.
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Acknowledgements We commend AI Hofmann, the 1996 Goldschmidt medalist, and recognize with pleasure the decades of stimulation and liveliness he has brought to our science. We are most grateful to Peter Kelemen for his tutorials in MELTS and his unselfish sharing of his mind, his computer and his time. We are grateful to Mark Kurz for access to the VG-354 facility at WHOI. Very helpful and perceptive reviews were provided by Simon Turner, Andy Saunders and Malcolm Hole; we have hardly done justice to them in the limited turn-around time available. This work was supported by DPP-9117853 and DPP 9419094 to SRH and DPP-7727546 and DPP8020836 to WEL.
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