Abstract

Xenoliths from the southern flank of Mauna Kea volcano form two broad categories. (1) Ultramafic: porphyroclastic dunite, wehrlite, and olivine clinopyroxenite (Fo89.4–83.6, clinopyroxene mg-number 90.3–86.3, spinel mg-number 57–42, spinel cr-number 7–52, no palgioclase); and granular wehrlite and olivine clinopyroxenite (Fo83–76) with plagioclase (An84–69) ± orthopyroxene, and Cr-magnetite. (2) Gabbroic: granular gabbro, gabbronorite, and troctolite composed of olivine + clinopyroxene frameworks (Fo82–74, mg-number 85–79) enclosing plagioclase (∼An79–69) ±orthopyroxene, and Fe–Ti oxides; and plagioclase (<An77) forming frameworks for, and fine-grained mosaics with, evolved olivine (Fo75–61), clinopyroxene±orthopyroxene, and Fe–Ti oxides. Most xenoliths are petrographically uniform, but some manifest modal, phase, cryptic, or grain-size layering, and some are composites of two rock types. Whole-xenolith incompatible elements are ‘depleted’, and there are positive Eu anomalies; 87Sr/86Sr is 0.70360, and mineral δ18O is 4.05–5.62. Porphyroclastic ultramafic xenoliths are gravity-settled and in situ cumulates from reservoir bottoms. Plagioclase-bearing xenoliths represent modal, phase, and cryptic layering (e.g. wehrlite to gabbronorite) in reservoir-margin solidification zones superimposed with small-scale (centimeter) modal, cryptic, phase, and grain-size layering. Mineral compositions point to tholeiitic parentage for most xenoliths, but alkalic for some (e.g. clinopyroxene (Al2O3 >4 wt %). These Mauna Kea xenoliths are plutonic complements to postshield lavas (Hamakua volcanics), and they identify that stage of volcano development with 15–5 wt % MgO magmas that underwent processes intrinsic to mafic-layered intrusions; e.g. in situ and gravity-settled crystallization, extensive differentiation, varieties of layering, mobilizations of late-stage, evolved liquids, compaction and connective disturbances in reservoirs.

Introduction

Ultramafic and gabbroic xenoliths in cinder-cone ejecta and lavas on Mauna Kea volcano, Hawaii, offer opportunities to study crystallization environments and processes for magmas that constructed the Hawaiian Islands. Jackson et al., (1982) first provided a reconnaissance study of the xenoliths, and Fodor & Vandermeyden, (1988) later presented an account of mineral and whole-rock compositions of xenoliths at the summit cone. This expansion of Mauna Kea studies evaluates ultramafic and gabbroic xenoliths at three neighboring cinder cones on Mauna Kea's southern flank (Fig. 1). It reveals modal assemblages, textures, and mineral compositions more diverse than previously recognized, and phase, modal, rhythmic, and grain-size layerings—all characteristics of grand-scale layered intrusive bodies. Our study is an effort to elucidate the internal magmatic workings of Hawaiian volcanoes, as well as to understand how plutonic assemblages relate to the stages of Hawaiian magmatism. The results augment our understanding of magma-reservoir crystallization down to a centimeter-size scale and complement current knowledge about Hawaiian magma systems available from seismic and lava-composition studies.

Fig. 1.

Map showing the five shield volcanoes that make up the island of Hawaii, and the xenolith location on the southern flank of Mauna Kea volcano. The enlargement shows the topographic relief of the three cones, A, B, and C, sampled for xenoliths and their specific locations on the volcano.

Fig. 1.

Map showing the five shield volcanoes that make up the island of Hawaii, and the xenolith location on the southern flank of Mauna Kea volcano. The enlargement shows the topographic relief of the three cones, A, B, and C, sampled for xenoliths and their specific locations on the volcano.

Study Area

Mauna Kea is one of five volcanoes on the island of Hawaii and reaches an altitude of 4230 m (Fig. 1). Its lavas are categorized as shield-building and postshield (Hamakua and Laupahoehoe volcanics) (Frey et al., 1990, 1991; Wolfe et al., 1997). The youngest postshield substage (hawaiitic pyroclastics and lavas of the Laupahoehoe Volcanics, ∼66–4 ka) is pertinent here because it erupted xenoliths at dozens of sites located between ∼2800 m elevation and the volcano summit (e.g. Jackson et al., 1982). The sites we sampled are cinder cones on the southern flank of the volcano at ∼2900 m (Fig. 1). One cone is called Puu Kalepeamoa and another cone ∼1 km north is called Kilohana; we refer to them as cone A and cone C, respectively (Fig. 1). Each is ∼500 m across at its base and ∼125 m high. Nestled between them is the rim and shallow crater (∼25 m relief) of a third (unnamed) cone that we call B (Fig. 1).

Xenolith Rock Types

Sixty-eight studied xenoliths range from ∼4 to 30 cm. They are subrounded to angular, and they occur on the flanks and in the craters of the cones as individual blocks or with thin lava rinds; some form cores of volcanic bombs. The rock types are dunite, wehrlite, olivine clinopyroxenite, gabbro, gabbronorite, and troctolite (Table 1a and b). We broadly categorize them as ultramafic and gabbroic, where ultramafic has <10 vol. % plagioclase. We distinguish gabbronorite from gabbro by its interstitial and granular orthopyroxene. All rock types occur at cone C, and the proportions approximate those noted by Jackson et al., (1982) for this location. Cone A yielded all but troctolite, and contains a preponderance of gabbronorite. Dunite, wehrlite, gabbro, and gabbronorite are at cone B. Table 1(a and b) lists the rock types.

Table 1a

Ultramafic xenoliths from the southern flank of Mauna Kea volcano, Hawaii: petrography, mineral endmembers, and layering

 Rock type Texture Fo cpx-mg-no. An opx-mg-no. Mode (vol. %) Layering 
       
 
 
       ol cpx Cr-oxide pl opx  
Cone A             
A19 dun porphyroclastic 87.8 89.6   94.9 1.6 3.5    
A11 dun porphyroclastic 87.6 89.4   92.6 3.1 4.3    
A16 dun porphyroclastic 86.5 89.4   97.8 0.2    
A6 dun porphyroclastic 86.5 88.7   97.8 0.4 1.8    
A27d dun porphyroclastic 85.2 89   92.8 1.3 5.9   modal 
A27w wehr porphyroclastic 84.9 88.1   53.6 45.1 1.3   modal 
A45 wehr anhedral gran/poik 84.7 88   74.2 24.4 1.4    
A17 wehr porphyroclastic 84.7 86.6   65.5 34 0.5    
A18 dun porphyroclastic 76.8 80.9   92.5 0.2 7.3    
A28 wehr anhedral gran/mosaic 76 81.9 70.4  43.4 51.8 1.3 3.2   
Cone B             
B22 dun porphyroclastic 89.4 90.3   98.5 0.3 1.2    
B12 dun porphyroclastic 84.4 86.9   97.3 0.7    
B1 wehr anhedral gran 81.7 84.7 72.3  83.8 14.6 1.6 tr   
Cone C             
C27 dun porphyroclastic 87.8 89.9   97.6 0.4    
C61 dun porphyroclastic 87.8 89.8   97.8 0.5 1.7    
C70 dun porphyroclastic 87.8 89.8   92.7 2.9 4.4    
C47 dun porphyroclastic 87.6 89.7   97.6 0.4    
C46 wehr porphyroclastic 86.4 88.1   64.6 33.9 1.5    
C28 dun porphyroclastic 86.1 89.8   98 0.2 1.8    
C29 wehr porphyroclastic 84.6 87.6   85.5 12.5    
C25d dun porphyroclastic 84.3 86.8   96.5 0.1 3.4   modal 
C25c ol cpxn't porphyroclastic 83.6 86.3   15.4 84.6    modal 
C99 wehr porphyroclastic 83.7 85.2   73.8 22.6 3.6    
C78 ol cpxn't anhedral gran 82.7 86 68.5  13.9 86  0.2   
C30 wehr poikilitic 82 85.8 70.8  67.5 29.3 1.7 1.5   
C9 ol cpxn't anhedral gran/mosaic 81.4 85 78.5 82.2 34.7 61.4 0.6 2.9 0.4  
C12 ol cpxn't anhedral gran/mosaic 81.3 85.2 83.7 82.9 33.8 65.4 0.1 0.6 0.1  
C75 ol cpxn't anhedral gran/mosaic 81 84   29.5 70.4 0.1   phase 
 gab nor* anhedral gran 80.4 82.7 65.7 82.2 1.5 37.7 0.1 48.1 12.7 phase 
C20 wehr anhedral gran/mosaic 80.9 85.3 84.6 82.9 58.4 39.3 0.2 1.8 0.2  
C52 wehr poikilitic 80.1 84.3 83.1  43.3 56.2 0.1 0.4   
C21 ol cpxn't anhedral gran/mosaic 77.3 82.1 64.4 79.4 18.2 81.4 0.1 0.2 0.1  
C77 wehr anhedral gran/mosaic 77.2 81.6 67.6 79.3 55.4 40.6 0.4 3.5 0.1  
 Rock type Texture Fo cpx-mg-no. An opx-mg-no. Mode (vol. %) Layering 
       
 
 
       ol cpx Cr-oxide pl opx  
Cone A             
A19 dun porphyroclastic 87.8 89.6   94.9 1.6 3.5    
A11 dun porphyroclastic 87.6 89.4   92.6 3.1 4.3    
A16 dun porphyroclastic 86.5 89.4   97.8 0.2    
A6 dun porphyroclastic 86.5 88.7   97.8 0.4 1.8    
A27d dun porphyroclastic 85.2 89   92.8 1.3 5.9   modal 
A27w wehr porphyroclastic 84.9 88.1   53.6 45.1 1.3   modal 
A45 wehr anhedral gran/poik 84.7 88   74.2 24.4 1.4    
A17 wehr porphyroclastic 84.7 86.6   65.5 34 0.5    
A18 dun porphyroclastic 76.8 80.9   92.5 0.2 7.3    
A28 wehr anhedral gran/mosaic 76 81.9 70.4  43.4 51.8 1.3 3.2   
Cone B             
B22 dun porphyroclastic 89.4 90.3   98.5 0.3 1.2    
B12 dun porphyroclastic 84.4 86.9   97.3 0.7    
B1 wehr anhedral gran 81.7 84.7 72.3  83.8 14.6 1.6 tr   
Cone C             
C27 dun porphyroclastic 87.8 89.9   97.6 0.4    
C61 dun porphyroclastic 87.8 89.8   97.8 0.5 1.7    
C70 dun porphyroclastic 87.8 89.8   92.7 2.9 4.4    
C47 dun porphyroclastic 87.6 89.7   97.6 0.4    
C46 wehr porphyroclastic 86.4 88.1   64.6 33.9 1.5    
C28 dun porphyroclastic 86.1 89.8   98 0.2 1.8    
C29 wehr porphyroclastic 84.6 87.6   85.5 12.5    
C25d dun porphyroclastic 84.3 86.8   96.5 0.1 3.4   modal 
C25c ol cpxn't porphyroclastic 83.6 86.3   15.4 84.6    modal 
C99 wehr porphyroclastic 83.7 85.2   73.8 22.6 3.6    
C78 ol cpxn't anhedral gran 82.7 86 68.5  13.9 86  0.2   
C30 wehr poikilitic 82 85.8 70.8  67.5 29.3 1.7 1.5   
C9 ol cpxn't anhedral gran/mosaic 81.4 85 78.5 82.2 34.7 61.4 0.6 2.9 0.4  
C12 ol cpxn't anhedral gran/mosaic 81.3 85.2 83.7 82.9 33.8 65.4 0.1 0.6 0.1  
C75 ol cpxn't anhedral gran/mosaic 81 84   29.5 70.4 0.1   phase 
 gab nor* anhedral gran 80.4 82.7 65.7 82.2 1.5 37.7 0.1 48.1 12.7 phase 
C20 wehr anhedral gran/mosaic 80.9 85.3 84.6 82.9 58.4 39.3 0.2 1.8 0.2  
C52 wehr poikilitic 80.1 84.3 83.1  43.3 56.2 0.1 0.4   
C21 ol cpxn't anhedral gran/mosaic 77.3 82.1 64.4 79.4 18.2 81.4 0.1 0.2 0.1  
C77 wehr anhedral gran/mosaic 77.2 81.6 67.6 79.3 55.4 40.6 0.4 3.5 0.1  

Modes are based on 1500–2000 counts per thin section. dun, dunite; wehr, wehrlite; ol cpxn't, olivine clinopyroxenite; gab nor, gabbronorite (see footnote*).

*

C75 is phase layered, where one layer is gabbronorite and therefore not ultramafic, but included in this table none the less because the xenolith is a composite.

Anhedral granular: mosaic refers to an overprint of polygonal grains in triple-point junctures.

Lower-case d, w, and c in sample numbers respectively refer to dunite, wehrlite, and olivine clinopyroxenite portions of composite xenoliths. ‘Modal’ and ‘phase’ refer to the type of layering in composite xenoliths.

Table 1b

Gabbroic xenoliths from the southern flank of Mauna Kea volcano, Hawaii: petrography, mineral endmembers, and layering

 Rock type General texture Fo cpx mg- no. An opx mg- no. pl occurrence Modes (vol. %) Layering 
        
 
 
        ol cpx pl opx opaque amph  
Cone A 
A2 gab nor anhedral gran 80.7 84.9 78.3 82.7 interstit 33.6 45.4 20.7 0.2 0.2   
A14 gab nor anhedral gran/mosaic 79.7 83.9 76.2 82.1 interstit 43.3 44.6 10.4 0.3 1.4   
A21 gab nor anhedral gran 78 82.6 72.2 78.1 interstit 14.1 62.9 22.7 0.1 0.3   
A22 gab (ol) anhedral gran 77 81.6 76.9  interstit 28 36.2 35.7  0.1  modal–grain 
              size–cryptic 
 gab (pl) fine-grained mosaic 75.2 80.2 75.1  gran/mosaic 12.6 21.4 65.9  0.1   
A12 gab nor anhedral gran/mosaic 76.9 82.6 70.9 79.5 interstit 15.7 53.5 17.6 12.8 0.3 0.1  
A7 gab nor anhedral gran 76.9 81.5 75.3 78.2 gran/mosaic 10.3 55.6 33.2 0.2 0.6 0.1  
A13 gab nor anhedral gran 76.8 81.1 72.3 78.8 interstit 15.6 50.3 31.7 1.7 0.3 0.4 phase 
 gab nor anhedral gran ′  ′ ′ interstit  57.5 26.9 15.2 0.4   
A5 gab nor anhedral gran 76.8 81.2 59.8 78.2 gran/mosaic 35.2 32.3 24.8 5.7   
A3 gab nor anhedral gran 73.6 79.6 65.1 77.5 gran/mosaic 11 58 24.4 0.5 6.1   
A9 gab (ol) anhedral gran/mosaic 73.4 79.3 66.5  interstit 52.5 32.5 11.9   modal–grain 
              size–cryptic 
 gab (pl) fine-grained mosaic 71.8 78.5 69.1  gran/mosaic 10.4 35 52.1  2.3   
A10 gab nor anhedral gran 72.4 80.3 66.6 77.3 gran 27.9 23.3 30.2 15.8 1.8  
A15 gab anhedral gran/mosaic 72 77.8 70.6  gran/mosaic 12.8 25.6 56.8  4.8   
A2-2 gab nor fine-grained mosaic 69.4 77.8 60.1 73.8 gran/mosaic 4.5 38.1 47.9 2.6 6.8   
Cone B 
B8 gab anhedral gran altered 83.5 80.1  interstit 62.8 14.9 21.7  0.5   
B5 gab nor anhedral gran  79.1 62.8 74.4 gran/mosaic  44.3 46.4 0.1 9.1   
B4 gab nor anhedral gran 68.3 78 63.1 73.7 gran/mosaic 30 65.7 0.2 1.1   
B3 gab nor fine-grained mosaic  76.5 61.4 71.8 gran/mosaic  36.9 43.9 4.3 14.9   
B2 gab nor anhedral gran  77 63 73.3 gran/mosaic  38.7 47.5 13.4 0.4   
Cone C 
C7 gab anhedral gran/mosaic 81.8 84.5 79.1  interstit 67.3 19.5 11.5  1.5   
C17 gab nor anhedral gran/mosaic 80.9 83.4 68.5 82.7 interstit 70.2 10.4 14.9 4.3 0.2   
C8 gab anhedral gran/mosaic 77.4 80.8 69.1  interstit 63.4 19.8 14.1  2.7   
C22 gab anhedral gran/mosaic 75.9 79 75.3  interst/gran 11.1 42.7 32.5  12.9 0.8 variable cpx–ol–opx props 
C15 gab anhedral gran/mosaic 75.8 83.4 69.1  interstit 67.3 13 17.5  1.8 0.4  
C50 gab nor anhedral gran/mosaic 75.6 81.6 71.4 78.4 interst/gran 8.7 28.7 48.4 8.1 6.1  modal 
C18 gab anhedral gran/mosaic 74.8 78.9 66.7  gran/mosaic 1.1 35.9 60.4  0.6 modal 
C3 gab nor anhedral gran/mosaic 74.7 79.3 77.3 73.8 gran/mosaic 12.6 21.6 65.1  0.5 0.1  
C59 gab anhedral gran 74.4 81.2 73.5  gran 21 50.4 24.3  1.3 2.9  
C16 gab anhedral gran/mosaic 73.9 80 62  gran/mosaic 10.2 33.8 54.4  1.1   
C49 troc anhedral gran 73.8 81.2 64.1  gran 18.4 0.2 59.1  22.2   
C10c gab anhedral gran 73.4 79 69.8  gran/mosaic 14.7 15.8 69.4  0.1  grain size 
C10f gab fine-grained mosaic 73.6 79.1 71.4  gran/mosaic 16.4 39.3 42.9  0.9 0.5 grain size 
C76 gab nor fine-grained mosaic 72.3 81.3 66.3 74 gran/mosaic 14.7 60.7 21.4 0.5 2.7   
C6 troc anhedral gran/mosaic 72.3 79.6 67.1  gran/mosaic 27.3 2.7 68.2  1.8   
C14 troc anhedral gran/mosaic 72.3 80.9 62.1  gran/mosaic 30.3 1.1 67.6  0.9   
C2 gab nor fine-grained mosaic 71.9 78.7 63.4 75.2 gran/mosaic 5.8 35.1 50.3 0.1 8.7   
C5 gab nor fine-grained mosaic 71.7 77.3 64 72.3 gran/mosaic 8.7 36.5 48.4 0.2 6.1 0.1  
C4 gab nor fine-grained mosaic 70.6 78.7 63.2 71.7 gran/mosaic 0.4 30.2 55.2 0.2 14   
C1c gab anhedral gran/mosaic idd 79.1 65.8  gran/mosaic 11.1 31.1 55.6  2.2  grain size 
C1f gab fine-grained mosaic idd 78.1 71.9  gran/mosaic 13.1 35.5 45.1  6.3  grain size 
C19c gab nor anhedral gran/mosaic 62–66 78.4 62.9 69.8 gran/mosaic 38 45.9 0.1 13  grain size 
C19f gab nor fine-grained mosaic 61 77.9 64.7 68.7 gran/mosaic 0.1 48 45.8 0.4 5.4 0.3 grain size 
 Rock type General texture Fo cpx mg- no. An opx mg- no. pl occurrence Modes (vol. %) Layering 
        
 
 
        ol cpx pl opx opaque amph  
Cone A 
A2 gab nor anhedral gran 80.7 84.9 78.3 82.7 interstit 33.6 45.4 20.7 0.2 0.2   
A14 gab nor anhedral gran/mosaic 79.7 83.9 76.2 82.1 interstit 43.3 44.6 10.4 0.3 1.4   
A21 gab nor anhedral gran 78 82.6 72.2 78.1 interstit 14.1 62.9 22.7 0.1 0.3   
A22 gab (ol) anhedral gran 77 81.6 76.9  interstit 28 36.2 35.7  0.1  modal–grain 
              size–cryptic 
 gab (pl) fine-grained mosaic 75.2 80.2 75.1  gran/mosaic 12.6 21.4 65.9  0.1   
A12 gab nor anhedral gran/mosaic 76.9 82.6 70.9 79.5 interstit 15.7 53.5 17.6 12.8 0.3 0.1  
A7 gab nor anhedral gran 76.9 81.5 75.3 78.2 gran/mosaic 10.3 55.6 33.2 0.2 0.6 0.1  
A13 gab nor anhedral gran 76.8 81.1 72.3 78.8 interstit 15.6 50.3 31.7 1.7 0.3 0.4 phase 
 gab nor anhedral gran ′  ′ ′ interstit  57.5 26.9 15.2 0.4   
A5 gab nor anhedral gran 76.8 81.2 59.8 78.2 gran/mosaic 35.2 32.3 24.8 5.7   
A3 gab nor anhedral gran 73.6 79.6 65.1 77.5 gran/mosaic 11 58 24.4 0.5 6.1   
A9 gab (ol) anhedral gran/mosaic 73.4 79.3 66.5  interstit 52.5 32.5 11.9   modal–grain 
              size–cryptic 
 gab (pl) fine-grained mosaic 71.8 78.5 69.1  gran/mosaic 10.4 35 52.1  2.3   
A10 gab nor anhedral gran 72.4 80.3 66.6 77.3 gran 27.9 23.3 30.2 15.8 1.8  
A15 gab anhedral gran/mosaic 72 77.8 70.6  gran/mosaic 12.8 25.6 56.8  4.8   
A2-2 gab nor fine-grained mosaic 69.4 77.8 60.1 73.8 gran/mosaic 4.5 38.1 47.9 2.6 6.8   
Cone B 
B8 gab anhedral gran altered 83.5 80.1  interstit 62.8 14.9 21.7  0.5   
B5 gab nor anhedral gran  79.1 62.8 74.4 gran/mosaic  44.3 46.4 0.1 9.1   
B4 gab nor anhedral gran 68.3 78 63.1 73.7 gran/mosaic 30 65.7 0.2 1.1   
B3 gab nor fine-grained mosaic  76.5 61.4 71.8 gran/mosaic  36.9 43.9 4.3 14.9   
B2 gab nor anhedral gran  77 63 73.3 gran/mosaic  38.7 47.5 13.4 0.4   
Cone C 
C7 gab anhedral gran/mosaic 81.8 84.5 79.1  interstit 67.3 19.5 11.5  1.5   
C17 gab nor anhedral gran/mosaic 80.9 83.4 68.5 82.7 interstit 70.2 10.4 14.9 4.3 0.2   
C8 gab anhedral gran/mosaic 77.4 80.8 69.1  interstit 63.4 19.8 14.1  2.7   
C22 gab anhedral gran/mosaic 75.9 79 75.3  interst/gran 11.1 42.7 32.5  12.9 0.8 variable cpx–ol–opx props 
C15 gab anhedral gran/mosaic 75.8 83.4 69.1  interstit 67.3 13 17.5  1.8 0.4  
C50 gab nor anhedral gran/mosaic 75.6 81.6 71.4 78.4 interst/gran 8.7 28.7 48.4 8.1 6.1  modal 
C18 gab anhedral gran/mosaic 74.8 78.9 66.7  gran/mosaic 1.1 35.9 60.4  0.6 modal 
C3 gab nor anhedral gran/mosaic 74.7 79.3 77.3 73.8 gran/mosaic 12.6 21.6 65.1  0.5 0.1  
C59 gab anhedral gran 74.4 81.2 73.5  gran 21 50.4 24.3  1.3 2.9  
C16 gab anhedral gran/mosaic 73.9 80 62  gran/mosaic 10.2 33.8 54.4  1.1   
C49 troc anhedral gran 73.8 81.2 64.1  gran 18.4 0.2 59.1  22.2   
C10c gab anhedral gran 73.4 79 69.8  gran/mosaic 14.7 15.8 69.4  0.1  grain size 
C10f gab fine-grained mosaic 73.6 79.1 71.4  gran/mosaic 16.4 39.3 42.9  0.9 0.5 grain size 
C76 gab nor fine-grained mosaic 72.3 81.3 66.3 74 gran/mosaic 14.7 60.7 21.4 0.5 2.7   
C6 troc anhedral gran/mosaic 72.3 79.6 67.1  gran/mosaic 27.3 2.7 68.2  1.8   
C14 troc anhedral gran/mosaic 72.3 80.9 62.1  gran/mosaic 30.3 1.1 67.6  0.9   
C2 gab nor fine-grained mosaic 71.9 78.7 63.4 75.2 gran/mosaic 5.8 35.1 50.3 0.1 8.7   
C5 gab nor fine-grained mosaic 71.7 77.3 64 72.3 gran/mosaic 8.7 36.5 48.4 0.2 6.1 0.1  
C4 gab nor fine-grained mosaic 70.6 78.7 63.2 71.7 gran/mosaic 0.4 30.2 55.2 0.2 14   
C1c gab anhedral gran/mosaic idd 79.1 65.8  gran/mosaic 11.1 31.1 55.6  2.2  grain size 
C1f gab fine-grained mosaic idd 78.1 71.9  gran/mosaic 13.1 35.5 45.1  6.3  grain size 
C19c gab nor anhedral gran/mosaic 62–66 78.4 62.9 69.8 gran/mosaic 38 45.9 0.1 13  grain size 
C19f gab nor fine-grained mosaic 61 77.9 64.7 68.7 gran/mosaic 0.1 48 45.8 0.4 5.4 0.3 grain size 

Modes are based on 1500–2000 counts per thin section. gab, gabrro; gab nor, gabbronorite; troc, troctolite; (ol) and (pl) refer to olivine-rich and plagioclase-rich portions of modally layered gabbroic xenoliths. Anhedral granular/mosaic refers to an overprint of polygonal grains in triple-point junctures. ‘idd’ refers to olivine altered to iddingsite. ‘pl occurrence’ refers to the dominant way in which plagioclase occurs: interstitial to ol + cpx frameworks, granularly, and in mosaics of polygonal grains in triple-point relationships. ‘Layering’ refers to types of layering observed in xenolith; this column includes reference to some variable clinopyroxene–olivine–plagioclase proportions, or varying proportions of these phases across a cut surface of the sample. Lower-case c, and f in sample numbers refer to coarse and fine portions of grain-size layered gabbroic xenoliths.

Petrography

Uniform and nonuniform

Most xenoliths have uniform modes, textures, and grain sizes, but some vary in these features. Plagioclase-bearing ultramafic xenoliths, for instance, have modal olivine/clinopyroxene ratios that change across a few centimeters distance, and plagioclase occurs concentrated in areas, or ‘pools’, that may be a few millimeters thick (Fig. 2a). Depending on the cut portion of a xenolith examined, then, an overall olivine clinopyroxenitic or wehrlitic xenolith may locally (1–2 cm2) be dunitic or gabbroic. Similarly, distributions of mafic minerals in plagioclase-rich xenoliths may be nonuniform across a surface of, say, 10 cm2 (Fig. 2b). Accordingly, single thin sections for some xenoliths provide only good approximations of modal percentages.

Fig. 2.

(a) Wehrlite (C20) and olivine clinopyroxenite (C12) contain plagioclase interstitial to olivine+clinopyroxene frameworks, but plagioclase is sometimes concentrated (arrows), particularly in xenolith C20 as a layer several millimeters thick. Throughout, the olivine and clinopyroxene are in varying amounts and modal olivine/clinopyroxene ratios accordingly create local dunitic areas. (b) Plagioclase-free olivine clinopyroxenite C75 is ‘capped’ by gabbronorite to create phase layering; plagioclase-rich gabbro C3 and troctolite C14 have uneven distributions of mafic minerals (arrows show depletions), creating areas unrepresentative of overall modes. (c) A pyroxenitic layer in gabbronorite C50 and anorthositic and clinopyroxenitic layers in gabbro C18. (d) Modal layering creating dunite–wehrlite and dunite–clinopyroxenite composite xenoliths. (e) Small-amplitude folds, tapered layers, and faulted layers (arrow) are evident in some gabbroic xenoliths of cone C (these two not examined quantitatively). (f) Gabbros A22 and A9 are modally layered with olivine-rich and plagioclase-rich portions; there are also grain-size distinctions and mineral composition distinctions (cryptic layering) between the two types of portions in each gabbro. Gabbronorite A13 has phase layering, where olivine is absent from the lower zone as shown. (g) Fine- and coarse-grained composite gabbros (grain-size layering); the fine-grained portion of C10 is foliated. (h) A photomicrograph (plane-polarized light) showing the interface between the fine and coarse portions of composite C1 in (g).

Fig. 2.

(a) Wehrlite (C20) and olivine clinopyroxenite (C12) contain plagioclase interstitial to olivine+clinopyroxene frameworks, but plagioclase is sometimes concentrated (arrows), particularly in xenolith C20 as a layer several millimeters thick. Throughout, the olivine and clinopyroxene are in varying amounts and modal olivine/clinopyroxene ratios accordingly create local dunitic areas. (b) Plagioclase-free olivine clinopyroxenite C75 is ‘capped’ by gabbronorite to create phase layering; plagioclase-rich gabbro C3 and troctolite C14 have uneven distributions of mafic minerals (arrows show depletions), creating areas unrepresentative of overall modes. (c) A pyroxenitic layer in gabbronorite C50 and anorthositic and clinopyroxenitic layers in gabbro C18. (d) Modal layering creating dunite–wehrlite and dunite–clinopyroxenite composite xenoliths. (e) Small-amplitude folds, tapered layers, and faulted layers (arrow) are evident in some gabbroic xenoliths of cone C (these two not examined quantitatively). (f) Gabbros A22 and A9 are modally layered with olivine-rich and plagioclase-rich portions; there are also grain-size distinctions and mineral composition distinctions (cryptic layering) between the two types of portions in each gabbro. Gabbronorite A13 has phase layering, where olivine is absent from the lower zone as shown. (g) Fine- and coarse-grained composite gabbros (grain-size layering); the fine-grained portion of C10 is foliated. (h) A photomicrograph (plane-polarized light) showing the interface between the fine and coarse portions of composite C1 in (g).

Nonuniformity also occurs as layers, ∼0.3–3 cm, of concentrated plagioclase, or olivine, or pyroxene to form local leuco-gabbro, anorthosite, troctolite, wehrlite, or pyroxenite (Fig. 2c). Modal layering occurs where dunites are in composite xenoliths with wehrlite or olivine clinopyroxenite (Fig. 2d), and there is phase layering in a composite of gabbronorite and plagioclase-free olivine clinopyroxenite (Fig. 2b), and in a gabbronorite with an olivine-free zone (Fig. 2f). Some xenoliths have tapering, contorted layers that describe small-amplitude folds, and some xenoliths display small-scale faults (Fig. 2e).

Grain-size layering creates composite gabbroic xenoliths that are uniform with respect to modes but have fine-grained (∼0.5–1 mm) portions separated from coarse-grained (∼1–3 mm) portions by interfaces (Fig. 2g and h). Sometimes modal (and cryptic) layering attends grain-size layering (olivine-rich portions, 1–5 mm; plagioclase-rich portions, 0.5–1 mm; Fig. 2f). Table 1(a and b) identifies the layered xenoliths.

Rock types and textures

Dunite

Dunites are porphyroclastic (Fig. 3a) and have large (3–5 mm) strained, or kink-banded, olivines within smaller, <2 mm, equigranular, polygonal olivine. Cr-spinel occurs within and among olivine grains, and clinopyroxene is interstitial; both phases are generally <5 vol. %. No plagioclase was observed in dunite, but one sample, A18, contains a titaniferous magnetite vein with apatite.

Fig. 3.

(a) Large kink-banded olivine adjacent to a mosaic of smaller olivine grains in porphyroclastic dunite C70; crossed-nicols. (b) Porphyroclastic wehrlite A17 with large olivine and clinopyroxene in a mosaic of smaller olivine; crossed nicols. (c) Anhedral–granular texture of wehrlite C77 with interstitial plagioclase (one area labeled); plane-polarized light. (d) Gabbro C8 is a framework of olivine+clinopyroxene with ∼14 vol. % interstitial plagioclase; plane-polarized light. (e) Gabbronorite A10 is composed of about one-third plagioclase, too large a volume to any longer be interstitial or intergranular within an olivine+clinopyroxene framework; plane-polarized light. (f) Gabbro C16 is over one-half plagioclase, which creates a framework for olivine and clinopyroxene; plane-polarized light. (g) Fine-grained mosaic-textured gabbronorite C5; crossed-nicols. (h) Chadacrysts of olivine in an oikocryst of clinopyroxene in wehrlite C30.

Fig. 3.

(a) Large kink-banded olivine adjacent to a mosaic of smaller olivine grains in porphyroclastic dunite C70; crossed-nicols. (b) Porphyroclastic wehrlite A17 with large olivine and clinopyroxene in a mosaic of smaller olivine; crossed nicols. (c) Anhedral–granular texture of wehrlite C77 with interstitial plagioclase (one area labeled); plane-polarized light. (d) Gabbro C8 is a framework of olivine+clinopyroxene with ∼14 vol. % interstitial plagioclase; plane-polarized light. (e) Gabbronorite A10 is composed of about one-third plagioclase, too large a volume to any longer be interstitial or intergranular within an olivine+clinopyroxene framework; plane-polarized light. (f) Gabbro C16 is over one-half plagioclase, which creates a framework for olivine and clinopyroxene; plane-polarized light. (g) Fine-grained mosaic-textured gabbronorite C5; crossed-nicols. (h) Chadacrysts of olivine in an oikocryst of clinopyroxene in wehrlite C30.

Wehrlite and olivine clinopyroxenite

Wehrlites and olivine clinopyroxenites without plagioclase are almost all porphyroclastic with olivine, commonly strained, and clinopyroxene, 3–5 mm, within mosaics of polygonal olivine and clinopyroxene (Fig. 3b); one plagioclase-free wehrlite is granular. Cr-spinel occurs within and between the silicate grains.

These rock types with interstitial plagioclase (∼0.2–3.5 vol. %) are mainly anhedral granular, but some wehrlites are poikilitic. The granular samples contain areas where olivine and clinopyroxene are in mosaics of polygonal grains, and some have interstitial orthopyroxene. Strained olivine grains are present but not common. Cr-magnetite is the opaque phase.

Gabbro, gabbronorite, and troctolite

Coarse anhedral–granular. Most gabbroic xenoliths have anhedral–granular textures (Fig. 3c-f) with grains ∼0.5–4 mm, but some ∼5–12 mm. The granularity of most of these xenoliths, which is usually manifested as interlocking anhedral (occasionally subhedral) olivine, clinopyroxene, and plagioclase, is overprinted to some extent by mosaics of equigranular ∼0.1–3 mm grains that share polygonal grain boundaries and triple-point junctures (Fig. 3c-f). Strained olivine grains are rare. Orthopyroxene is usually intergranular, <0.5 mm, but sometimes large, 2–4 mm, and enclosing other phases. Generally, Cr–Fe–Ti oxides are <5 vol. % (one troctolite has ∼22 vol. % ilmenite) and occur interstitially and as inclusions in pyroxene and interstitial plagioclase. Some gabbroic xenoliths have amphibole intimately associated with clinopyroxene, and it is probably an alteration product thereof.

We further evaluate coarse, gabbroic xenoliths in terms of plagioclase abundances (∼10–68 vol. %; Table 1b) and the petrographic aspects that attend those modal variations. Specifically, one variety has up to ∼30–35 vol. % plagioclase that is intergranular within olivine+clinopyroxene frameworks; this intra-framework plagioclase often consists of mosaics of polygonal grains (Fig. 3c and d). Another variety occurs when plagioclase exceeds ∼35 vol. % and its coexisting olivine and clinopyroxene are too dispersed to constitute frameworks to contain the plagioclase (Fig. 3e). Ultimately, where plagioclase is >50 vol. %, as in troctolites, it forms frameworks of polygonal grains enclosing olivine, clinopyroxene, and Fe–Ti oxide (Fig. 3f).

Fine-grained mosaic. Gabbro and gabbronorite xenoliths also occur as fine-grained equigranular mosaics of anhedral and polygonal 0.2–1.0-mm plagioclase, clinopyroxene, olivine, and Fe–Ti oxides (Fig. 3g). They generally have >40 vol. % plagioclase. Textures suggest undercooling and co-crystallization of silicate phases; Fe–Ti oxides, however (5–10 vol. %), are largely interstitial and probably late-forming. Some fine-grained xenoliths are foliated owing to <1 mm layers of concentrated clinopyroxene, olivine, or plagioclase (Fig. 2g), and some are portions of composite fine- and coarse-grained gabbroic xenoliths (grain-size layering; Fig. 2g and h).

Recrystallization and unmixing features

Polygonal olivine, clinopyroxene, and/or plagioclase in nearly all xenoliths (Fig. 3c-f) suggest annealing or recrystallization histories. These grain boundaries can also be described as adcumulus and forming when interstitial liquids contributed to crystal growth and mutual interference boundaries (e.g. Hunter, 1987). Some gabbroic xenoliths have undergone subsolidus unmixing to create exsolution lamellae of orthopyroxene or Fe–Ti oxide.

Weathering or alteration

There are three types of secondary modifications. (1) Olivine is oxidized to ‘iddingsite’ in some cone-A and cone-C xenoliths, and in most cone-B xenoliths. (2) Extreme alteration, perhaps deuteric, in cone-B xenoliths blackened some olivine; in these samples, olivine grains can be entirely opaque or contain seemingly fresh areas mixed in with opaque areas. (3) Amphibole is an alteration or reaction phase within some clinopyroxene grains in cone-A and cone-C gabbroic xenoliths.

Analytical Techniques

We determined mineral compositions with ARL-EMX and ARL-SEMQ electron microprobes at North Carolina State University using olivine, clinopyroxene, orthopyroxene, plagioclase, microcline, spinel, and ilmenite provided by the Smithsonian Institution and Ni-doped diopside as reference minerals. We applied the matrix corrections of Bence & Albee, (1968) for EMX analyses and phi–rho–Z for SEMQ analyses. Our data tables present average values for 8–15 analytical points on single grains that we assessed as representative of a sample after analyzing several of its grains.

Whole-rock samples of selected xenoliths were analyzed in duplicate for major- and trace-element abundances (except REE) on a Phillips 1410 X-ray fluorescence spectrometer at NCSU; procedures have been described by Fodor et al., (1992). Major elements, except K, Na, and P, were analyzed using glass disks; trace elements and K, Na, and P were determined on pressed powder. Reference samples for calibration curves include standards of the US, Canadian, South African, and Japanese geological surveys. Neutron activation analyses for rare-earth element abundances on whole-rocks, clinopyroxene, and plagioclase were performed at the Oregon State University Radiation Center. Before those analyses, we cleaned the minerals in warm 10% HCl overnight. Sr and Pb isotope compositions for xenoliths and clinopyroxene grains are from analyses we did at the University of North Carolina isotope facility, using a VG Sector 54 mass spectrometer. The 87Sr/86Sr ratio for SRM987 is 0.710254±15 (2σ), and Pb is corrected for mass discrimination by a factor of 0.15%/a.m.u. Krueger Enterprises provided the oxygen isotope analyses for mineral separates.

Mineral Compositions

Olivine

Ultramafic xenoliths without plagioclase have compositionally primitive olivine, ∼Fo89.4–83.6, except dunite A18 with Fo76.8 (Table 2; Fig. 4). Plagioclase-bearing wehrlite and olivine clinopyroxenite have olivine (Table 3) that is slightly more evolved, Fo82.7–76, and gabbroic xenoliths have even more evolved olivine, Fo81–61. The fine-grained variety invariably has the most evolved olivine, <Fo73 (Table 1b). Compositional zoning in any of the xenoliths is generally <3 mol %. The Fo range we observe in the xenoliths collectively has also been observed for olivine phenocrysts in Mauna Kea lavas (Frey et al., 1990, 1991; Yang et al., 1994; Baker et al., 1996).

Table 2

Olivine, pyroxene, plagioclase, and oxide compositions (in wt %) for representative ultramafic xenoliths of Mauna Kea volcano

 Dunite 
 B22 C27 A19 A11 A6 A16 B12 A18 
 
 

 

 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp sp ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx Cr-mt Cr-mt Ti-mt 
SiO2 39.9 51.9  40 53  39.6 51.6   39.5 51.6  39.8 51.9  39.6 52.8  39 52.6  38 52.2    
TiO2  0.76 2.4  0.55  0.7 2.2 1.9  0.88 1.8  0.75 2.5  0.47 1.9  0.6 1.8  0.36 3.4 2.2 14.7 
Al2O3  15.2  2.5 13.2  2.8 17.4 15.5  4.2 16.4  3.6 14.2  2.2 12.4  2.9 13.7  2.7 9.7 10 5.8 
Cr2O3  1.2 39.3  0.93 40.9  0.97 39.9 42  0.71 41.5  0.91 37.5  41.4  0.87 32.5  0.8 21.2 24.5 3.3 
FeO 10.3 3.2 29.1 12 3.4 29.5 11.8 3.6 25.4 26.1 11.9 3.5 26.7 13 3.7 32.9 13 3.5 32 15 4.5 40.9 21.5 6.7 55.2 51 67.1 
MnO 0.14 0.06 0.31 0.09 0.04 0.26 0.16 0.08 0.22 0.23 0.14 0.1 0.2 0.14 0.1 0.28 0.15 0.1 0.24 0.22 0.09 0.29 0.41 0.33 0.46 0.42 0.51 
MgO 48.9 16.7 11.5 48.2 17 11.7 47.5 17.4 12.6 12.1 47.3 16.6 12.6 46.7 16.3 10.9 46.6 16.6 10.1 45.4 16.6 39.9 15.9 6.4 6.3 6.1 
CaO 0.09 22.6  0.14 22.8  0.1 21.8   0.13 21.8  0.17 22.6  0.17 22.9  0.14 22.8  0.14 20.8    
Na2 0.54   0.44   0.62    0.5   0.42   0.46   0.42   0.62    
K2                           
NiO 0.42   0.35   0.38    0.39   0.38   0.35   0.29   0.25     
Total 99.75 99.96 97.81 100.78 100.66 97.56 99.54 99.57 97.72 97.83 99.36 99.89 99.2 100.19 100.28 98.28 99.87 100.03 98.04 100.05 101.38 98.19 100.2 100.41 96.36 94.42  
Recalc. FeO  18.2   17.2   16.7 17   16.7   19.3   19.7   21.6   25.8 24.2 36.2 
Recalc. Fe2O3  12.1   13.6   9.6 10.2   11.1   15.1   13.7   21.5   32.7 29.7 34.4 
cr-no.   63.4   67.5   60.6 64.5   62.9   63.9   69.1   61.4   59.4 62.2 27.6 
mg-no. (Fo) 89.4 90.3 52.9 87.8 89.9 54.8 87.8 89.6 57.3 56 87.6 89.4 57 86.5 88.7 50.2 86.5 89.4 47.8 84.4 86.9 42.6 76.8 80.9 30.7 31.6 23.1 
Fs  5.2   5.4   5.6   5.7     5.6   7.1   10.7     
Wo  46.8   46.5   44.7    45.8   47   47   45.9   42.9    
An 
Or 
Equilib. T (°C) 
 Dunite 
 B22 C27 A19 A11 A6 A16 B12 A18 
 
 

 

 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp sp ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx Cr-mt Cr-mt Ti-mt 
SiO2 39.9 51.9  40 53  39.6 51.6   39.5 51.6  39.8 51.9  39.6 52.8  39 52.6  38 52.2    
TiO2  0.76 2.4  0.55  0.7 2.2 1.9  0.88 1.8  0.75 2.5  0.47 1.9  0.6 1.8  0.36 3.4 2.2 14.7 
Al2O3  15.2  2.5 13.2  2.8 17.4 15.5  4.2 16.4  3.6 14.2  2.2 12.4  2.9 13.7  2.7 9.7 10 5.8 
Cr2O3  1.2 39.3  0.93 40.9  0.97 39.9 42  0.71 41.5  0.91 37.5  41.4  0.87 32.5  0.8 21.2 24.5 3.3 
FeO 10.3 3.2 29.1 12 3.4 29.5 11.8 3.6 25.4 26.1 11.9 3.5 26.7 13 3.7 32.9 13 3.5 32 15 4.5 40.9 21.5 6.7 55.2 51 67.1 
MnO 0.14 0.06 0.31 0.09 0.04 0.26 0.16 0.08 0.22 0.23 0.14 0.1 0.2 0.14 0.1 0.28 0.15 0.1 0.24 0.22 0.09 0.29 0.41 0.33 0.46 0.42 0.51 
MgO 48.9 16.7 11.5 48.2 17 11.7 47.5 17.4 12.6 12.1 47.3 16.6 12.6 46.7 16.3 10.9 46.6 16.6 10.1 45.4 16.6 39.9 15.9 6.4 6.3 6.1 
CaO 0.09 22.6  0.14 22.8  0.1 21.8   0.13 21.8  0.17 22.6  0.17 22.9  0.14 22.8  0.14 20.8    
Na2 0.54   0.44   0.62    0.5   0.42   0.46   0.42   0.62    
K2                           
NiO 0.42   0.35   0.38    0.39   0.38   0.35   0.29   0.25     
Total 99.75 99.96 97.81 100.78 100.66 97.56 99.54 99.57 97.72 97.83 99.36 99.89 99.2 100.19 100.28 98.28 99.87 100.03 98.04 100.05 101.38 98.19 100.2 100.41 96.36 94.42  
Recalc. FeO  18.2   17.2   16.7 17   16.7   19.3   19.7   21.6   25.8 24.2 36.2 
Recalc. Fe2O3  12.1   13.6   9.6 10.2   11.1   15.1   13.7   21.5   32.7 29.7 34.4 
cr-no.   63.4   67.5   60.6 64.5   62.9   63.9   69.1   61.4   59.4 62.2 27.6 
mg-no. (Fo) 89.4 90.3 52.9 87.8 89.9 54.8 87.8 89.6 57.3 56 87.6 89.4 57 86.5 88.7 50.2 86.5 89.4 47.8 84.4 86.9 42.6 76.8 80.9 30.7 31.6 23.1 
Fs  5.2   5.4   5.6   5.7     5.6   7.1   10.7     
Wo  46.8   46.5   44.7    45.8   47   47   45.9   42.9    
An 
Or 
Equilib. T (°C) 
 Wehrlites 
 C46 A45 A17 C29 C99 C30 B1 
 
 

 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx cpx Cr-mt Cr-mt pl ol cpx Cr-mt pl 
SiO2 39.6 51.8  39.4 51.9  39.1 52  39.6 52.6  39.3 51.2  39 52.2 51.1   49.7 38.7 51.7  50.1 
TiO2  0.77 2.3  0.73  0.62 2.5  0.54 2.3  0.98 3.2  0.67 6.5 8.4   0.86 7.4  
Al2O3  3.6 19.7  3.1 18.7  3.7 19.8  2.9 17  4.4 17.5  2.9 11.3 9.9 31.6  3.1 10.4 32 
Cr2O3  37.1  32.7  0.91 34.9  0.93 38.4  0.77 32.1  0.89 27.5 25.9   0.76 24.7  
FeO 13.2 4.1 28.9 14.8 4.2 33.9 14.8 4.7 28.5 14.6 4.5 28.5 15.6 5.1 36.7 17.1 5.2 44.1 45.7 0.34 17.5 5.5 47.3 0.17 
MnO 0.18 0.09 0.24 0.21 0.09 0.29 0.12 0.11 0.24 0.11 0.06 0.27 0.21 0.12 0.31 0.26 0.12 0.13 0.34 0.33  0.22 0.12 0.29  
MgO 47 17 12.2 46.1 17.2 10.6 45.8 17 12.4 46.1 17.8 11.4 44.9 16.5 10.2 43.6 16.9 16 7.5 6.6  43.9 17.1 7.6  
CaO 0.11 21  0.09 21.8  0.07 21  0.09 20.6  0.14 21.4  0.06 21.2 21.4   14.3 0.07 20.8  14.4 
Na2 0.37   0.37   0.46   0.33   0.34   0.41 0.48   3.2  0.33  
K2                    0.11    0.08 
NiO 0.34   0.33   0.32   0.29   0.3   0.3      0.34    
Total 100.43 99.73 100.44 100.93 100.39 99.19 100.21 100.5 98.34 100.79 100.26 97.87 100.45 100.81 100.01 100.32 100.4 100.2 97.24 96.83 99.25 100.73 100.27 97.69 99.75 
Recalc. FeO   18.6   21.2   17.9   18.6   22    27.3 30.1    28.1  
Recalc. Fe2O3   11.3   14.1   11.8   11   16.2    18.6 17.4    21.4  
cr-no.   55.7   54   54.2   60.2   55.2    62 63.7    61.4  
mg-no. (Fo) 86.4 88.1 53.9 84.7 88 47.1 84.7 86.6 55.2 84.6 87.6 52.2 83.7 85.2 45.2 82 85.8 84.6 32.9 28.1  81.7 84.7 32.5  
Fs  6.7   6.7   6.1   7.2   8.2   8.5     8.8   
Wo  43.9   44.5   44.2   42.2   44.3   43.6 44.9     42.6   
An                     70.8    72.3 
Or                     0.65    0.48 
Equilib. T (°C) 
 Wehrlites 
 C46 A45 A17 C29 C99 C30 B1 
 
 

 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx sp ol cpx cpx Cr-mt Cr-mt pl ol cpx Cr-mt pl 
SiO2 39.6 51.8  39.4 51.9  39.1 52  39.6 52.6  39.3 51.2  39 52.2 51.1   49.7 38.7 51.7  50.1 
TiO2  0.77 2.3  0.73  0.62 2.5  0.54 2.3  0.98 3.2  0.67 6.5 8.4   0.86 7.4  
Al2O3  3.6 19.7  3.1 18.7  3.7 19.8  2.9 17  4.4 17.5  2.9 11.3 9.9 31.6  3.1 10.4 32 
Cr2O3  37.1  32.7  0.91 34.9  0.93 38.4  0.77 32.1  0.89 27.5 25.9   0.76 24.7  
FeO 13.2 4.1 28.9 14.8 4.2 33.9 14.8 4.7 28.5 14.6 4.5 28.5 15.6 5.1 36.7 17.1 5.2 44.1 45.7 0.34 17.5 5.5 47.3 0.17 
MnO 0.18 0.09 0.24 0.21 0.09 0.29 0.12 0.11 0.24 0.11 0.06 0.27 0.21 0.12 0.31 0.26 0.12 0.13 0.34 0.33  0.22 0.12 0.29  
MgO 47 17 12.2 46.1 17.2 10.6 45.8 17 12.4 46.1 17.8 11.4 44.9 16.5 10.2 43.6 16.9 16 7.5 6.6  43.9 17.1 7.6  
CaO 0.11 21  0.09 21.8  0.07 21  0.09 20.6  0.14 21.4  0.06 21.2 21.4   14.3 0.07 20.8  14.4 
Na2 0.37   0.37   0.46   0.33   0.34   0.41 0.48   3.2  0.33  
K2                    0.11    0.08 
NiO 0.34   0.33   0.32   0.29   0.3   0.3      0.34    
Total 100.43 99.73 100.44 100.93 100.39 99.19 100.21 100.5 98.34 100.79 100.26 97.87 100.45 100.81 100.01 100.32 100.4 100.2 97.24 96.83 99.25 100.73 100.27 97.69 99.75 
Recalc. FeO   18.6   21.2   17.9   18.6   22    27.3 30.1    28.1  
Recalc. Fe2O3   11.3   14.1   11.8   11   16.2    18.6 17.4    21.4  
cr-no.   55.7   54   54.2   60.2   55.2    62 63.7    61.4  
mg-no. (Fo) 86.4 88.1 53.9 84.7 88 47.1 84.7 86.6 55.2 84.6 87.6 52.2 83.7 85.2 45.2 82 85.8 84.6 32.9 28.1  81.7 84.7 32.5  
Fs  6.7   6.7   6.1   7.2   8.2   8.5     8.8   
Wo  43.9   44.5   44.2   42.2   44.3   43.6 44.9     42.6   
An                     70.8    72.3 
Or                     0.65    0.48 
Equilib. T (°C) 
 C20 C52 C77 A28 
 
 

 

 

 
 ol cpx opx Cr-mt pl ol cpx Cr-mt pl ol cpx opx Cr-mt pl ol cpx Ti-mt pl 
SiO2 38.3 52.3 55  46.7 38.5 51.1  47.5 38.1 51.6 54.2  51 37.4 50.6  49.8 
TiO2  0.7 0.34 5.4   0.86 8.9   0.65 0.27 17.5   1.2 15  
Al2O3  1.5 13.4 33.9  12.6 33.9  2.4 1.4 8.1 31.1  4.4 5.1 31.5 
Cr2O3  0.58 0.19 25.9   0.84 21   0.5 0.15 20.4   0.66 6.6  
FeO 18.1 5.1 11.3 43.2 0.44 18.6 5.5 46.3 0.37 21.4 6.9 13.7 42.1 0.47 22.3 6.3 63.1 0.41 
MnO 0.2 0.11 0.24 0.3  0.24 0.1 0.3  0.27 0.19 0.3 0.3  0.32 0.14 0.33  
MgO 42.9 16.6 30.7 8.2  42 16.5 7.1  40.6 17.2 29.5 7.1  0.07 16 6.2  
CaO 0.01 21.5 0.9  17.6 0.1 21.1  16.7 0.07 19.2 0.84  13.6 39.6 21.1  14.3 
Na2 0.33 0.01  1.7  0.38  1.8  0.34 0.01  3.5  0.33  3.2 
K2    0.08    0.12     0.17    0.2 
NiO 0.25     0.25    0.26     0.22    
Total 99.76 100.22 100.18 96.4 100.42 99.69 100.38 96.2 100.39 100.7 98.98 100.37  99.84 99.91 100.73 96.33 99.41 
Recalc. FeO    25.3    30     36.6    35.9  
Recalc. Fe2O3    20.3    18.1     6.1    30.2  
cr-no.    56.5    52.8     62.8    46.5  
mg-no. (Fo) 80.9 85.3 82.9 36.5  80.1 84.3 29.7  77.2 81.6 79.3 25.7  76 81.9 23.5  
Fs  8.2 16.8    8.9    11.1 20.3   10.2   
Wo  44.3 1.7    43.7    39.6 1.6    43.7   
An     84.8    83.1     67.6    70.4 
Or     0.46    0.71        1.2 
Equilib. T (°C) 1000        1087         
 C20 C52 C77 A28 
 
 

 

 

 
 ol cpx opx Cr-mt pl ol cpx Cr-mt pl ol cpx opx Cr-mt pl ol cpx Ti-mt pl 
SiO2 38.3 52.3 55  46.7 38.5 51.1  47.5 38.1 51.6 54.2  51 37.4 50.6  49.8 
TiO2  0.7 0.34 5.4   0.86 8.9   0.65 0.27 17.5   1.2 15  
Al2O3  1.5 13.4 33.9  12.6 33.9  2.4 1.4 8.1 31.1  4.4 5.1 31.5 
Cr2O3  0.58 0.19 25.9   0.84 21   0.5 0.15 20.4   0.66 6.6  
FeO 18.1 5.1 11.3 43.2 0.44 18.6 5.5 46.3 0.37 21.4 6.9 13.7 42.1 0.47 22.3 6.3 63.1 0.41 
MnO 0.2 0.11 0.24 0.3  0.24 0.1 0.3  0.27 0.19 0.3 0.3  0.32 0.14 0.33  
MgO 42.9 16.6 30.7 8.2  42 16.5 7.1  40.6 17.2 29.5 7.1  0.07 16 6.2  
CaO 0.01 21.5 0.9  17.6 0.1 21.1  16.7 0.07 19.2 0.84  13.6 39.6 21.1  14.3 
Na2 0.33 0.01  1.7  0.38  1.8  0.34 0.01  3.5  0.33  3.2 
K2    0.08    0.12     0.17    0.2 
NiO 0.25     0.25    0.26     0.22    
Total 99.76 100.22 100.18 96.4 100.42 99.69 100.38 96.2 100.39 100.7 98.98 100.37  99.84 99.91 100.73 96.33 99.41 
Recalc. FeO    25.3    30     36.6    35.9  
Recalc. Fe2O3    20.3    18.1     6.1    30.2  
cr-no.    56.5    52.8     62.8    46.5  
mg-no. (Fo) 80.9 85.3 82.9 36.5  80.1 84.3 29.7  77.2 81.6 79.3 25.7  76 81.9 23.5  
Fs  8.2 16.8    8.9    11.1 20.3   10.2   
Wo  44.3 1.7    43.7    39.6 1.6    43.7   
An     84.8    83.1     67.6    70.4 
Or     0.46    0.71        1.2 
Equilib. T (°C) 1000        1087         
 Olivine clinopyroxenites 
 C78 C12 C21 
 
 

 

 
 ol cpx pl ol cpx opx Cr-mt pl ol cpx opx ilm pl 
SiO2 39.3 52.3 51 38.9 52 53.2  46.5 38.2 52 55  52 
TiO2  0.6   0.61 0.32 4.8   0.91 0.35 51  
Al2O3  2.4 31.1  14.7 34.1  2.8 1.4 0.59 30.5 
Cr2O3  0.6   0.75 0.28 27.4   0.49 0.14 1.5  
FeO 16.5 5.1 0.37 17.6 5.1 11.5 42.7 0.37 21.2 6.3 13.5 36.7 0.33 
MnO 0.22 0.12  0.15 0.16 0.22 0.28  0.25 0.25 0.26 0.24  
MgO 44.1 17.5  43 16.4 31.3 8.1  40.5 16.2 29.2 8.1  
CaO 0.07 21.2 14.1 0.04 21.8 0.62  17.3 0.04 20.9  13.2 
Na2 0.24 3.4  0.27 0.04  1.8  0.39 0.05  3.9 
K2  0.3     0.1     0.21 
NiO 0.27   0.28     0.24     
Total 100.46 100.06 100.27 99.97 100.09 98.98 97.98 100.17 100.43 100.24 100.9 98.13 100.14 
Recalc. FeO       25.7     31.2  
Recalc. Fe2O3       18.8     6.2  
cr-no.       55.6       
mg-no. (Fo) 82.7 86  81.3 85.2 82.9 36  77.3 82.1 79.4   
Fs    8.2 16.9    10.2 20.1   
Wo  42.8   44.9 1.2    43.3 1.9   
An   68.5     83.7     64.4 
Or   1.7     0.58     1.2 
Equilib. T (°C)    969     1002    
 Olivine clinopyroxenites 
 C78 C12 C21 
 
 

 

 
 ol cpx pl ol cpx opx Cr-mt pl ol cpx opx ilm pl 
SiO2 39.3 52.3 51 38.9 52 53.2  46.5 38.2 52 55  52 
TiO2  0.6   0.61 0.32 4.8   0.91 0.35 51  
Al2O3  2.4 31.1  14.7 34.1  2.8 1.4 0.59 30.5 
Cr2O3  0.6   0.75 0.28 27.4   0.49 0.14 1.5  
FeO 16.5 5.1 0.37 17.6 5.1 11.5 42.7 0.37 21.2 6.3 13.5 36.7 0.33 
MnO 0.22 0.12  0.15 0.16 0.22 0.28  0.25 0.25 0.26 0.24  
MgO 44.1 17.5  43 16.4 31.3 8.1  40.5 16.2 29.2 8.1  
CaO 0.07 21.2 14.1 0.04 21.8 0.62  17.3 0.04 20.9  13.2 
Na2 0.24 3.4  0.27 0.04  1.8  0.39 0.05  3.9 
K2  0.3     0.1     0.21 
NiO 0.27   0.28     0.24     
Total 100.46 100.06 100.27 99.97 100.09 98.98 97.98 100.17 100.43 100.24 100.9 98.13 100.14 
Recalc. FeO       25.7     31.2  
Recalc. Fe2O3       18.8     6.2  
cr-no.       55.6       
mg-no. (Fo) 82.7 86  81.3 85.2 82.9 36  77.3 82.1 79.4   
Fs    8.2 16.9    10.2 20.1   
Wo  42.8   44.9 1.2    43.3 1.9   
An   68.5     83.7     64.4 
Or   1.7     0.58     1.2 
Equilib. T (°C)    969     1002    
 Composites 
 A27 (dunite) A27 (wehrlite) C25 (dunite) C25 (ol cpx′nt) C75 (ol cpx′nt) C75 (ol gabn′rt) 
 
 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp ol cpx ol cpx Cr-mt ol cpx opx ilm pl 
SiO2 38.8 53.2  38.7 52.6  39.3 52  39.5 52.1 38.7 52  38.3 51.2 55.1  50.9 
TiO2  0.42 2.4  0.54 3.3  0.53 2.6  0.67  0.64 4.9  0.66 0.41 51.4  
Al2O3  2.2 13.7  2.3 14.1  3.3 17.9  3.6  3.1 16.4  1.2 0.2 30.9 
Cr2O3  0.57 34.6  0.78 38.2  32.2  0.66  0.77 24.7  0.71 0.08 0.39  
FeO 14.2 3.8 34.2 14.3 4.2 31.7 15.1 4.6 33.9 15.8 5.3 17.9 5.9 44 18.3 6.4 11.6 34.2 0.3 
MnO 0.2 0.12 0.28 0.2 0.12 0.28 0.12 0.09 0.25 0.14 0.16 0.24 0.13 0.19 0.23 0.17 0.24 0.42  
MgO 45.8 17.2 10.3 45.2 17.5 10 45.5 17 11 45 17.1 42.8 17.4 8.3 42.2 17.2 30 10.1  
CaO 0.1 22.3  0.09 21.1  0.06 20.7  0.07 20.6 0.04 20.6  0.04 19.8 1.2  13.3 
Na2 0.3   0.33   0.38   0.36  0.32   0.36 0.03  3.7 
K2                  0.21 
NiO 0.32   0.31   0.31   0.26  0.24   0.25     
Total 99.42 100.11 95.48 98.8 99.47 97.58 100.39 99.6 97.85 100.77 100.55 99.92 100.86 98.49 99.32 99.5 99.86 96.71 99.31 
Recalc. FeO   19.2   21.1   19.8     23.7    27.8  
Recalc. Fe2O3   16.7   11.8   15.7     17.7    7.1  
cr-no.   62.9   64.5   54.8     48.7      
mg-no. (Fo) 85.2 89 48.9 84.9 88.1 45.8 84.3 86.8 49.8 83.6 85.2 81 84 41.2 80.4 82.7 82.2   
Fs    6.7   7.5   8.5  9.3   10.2 17.4   
Wo  43.4   43.3   43.2   42.5  41.7   40.7 2.3   
An                   65.7 
Or                   1.2 
Equilib. T (°C)                1071    
 Composites 
 A27 (dunite) A27 (wehrlite) C25 (dunite) C25 (ol cpx′nt) C75 (ol cpx′nt) C75 (ol gabn′rt) 
 
 

 

 

 

 

 
 ol cpx sp ol cpx sp ol cpx sp ol cpx ol cpx Cr-mt ol cpx opx ilm pl 
SiO2 38.8 53.2  38.7 52.6  39.3 52  39.5 52.1 38.7 52  38.3 51.2 55.1  50.9 
TiO2  0.42 2.4  0.54 3.3  0.53 2.6  0.67  0.64 4.9  0.66 0.41 51.4  
Al2O3  2.2 13.7  2.3 14.1  3.3 17.9  3.6  3.1 16.4  1.2 0.2 30.9 
Cr2O3  0.57 34.6  0.78 38.2  32.2  0.66  0.77 24.7  0.71 0.08 0.39  
FeO 14.2 3.8 34.2 14.3 4.2 31.7 15.1 4.6 33.9 15.8 5.3 17.9 5.9 44 18.3 6.4 11.6 34.2 0.3 
MnO 0.2 0.12 0.28 0.2 0.12 0.28 0.12 0.09 0.25 0.14 0.16 0.24 0.13 0.19 0.23 0.17 0.24 0.42  
MgO 45.8 17.2 10.3 45.2 17.5 10 45.5 17 11 45 17.1 42.8 17.4 8.3 42.2 17.2 30 10.1  
CaO 0.1 22.3  0.09 21.1  0.06 20.7  0.07 20.6 0.04 20.6  0.04 19.8 1.2  13.3 
Na2 0.3   0.33   0.38   0.36  0.32   0.36 0.03  3.7 
K2                  0.21 
NiO 0.32   0.31   0.31   0.26  0.24   0.25     
Total 99.42 100.11 95.48 98.8 99.47 97.58 100.39 99.6 97.85 100.77 100.55 99.92 100.86 98.49 99.32 99.5 99.86 96.71 99.31 
Recalc. FeO   19.2   21.1   19.8     23.7    27.8  
Recalc. Fe2O3   16.7   11.8   15.7     17.7    7.1  
cr-no.   62.9   64.5   54.8     48.7      
mg-no. (Fo) 85.2 89 48.9 84.9 88.1 45.8 84.3 86.8 49.8 83.6 85.2 81 84 41.2 80.4 82.7 82.2   
Fs    6.7   7.5   8.5  9.3   10.2 17.4   
Wo  43.4   43.3   43.2   42.5  41.7   40.7 2.3   
An                   65.7 
Or                   1.2 
Equilib. T (°C)                1071    

Prefixes A, B, and C in sample numbers refer to locations of sample sites–cones shown in Fig. 1. ol, olivine; cpx, clinopyroxene; opx, orthopyroxene; pl, plagioclase; sp, Cr-spinel; Cr-mt, Cr-rich magnetite; Ti-mt, Ti-rich magnetite; ilm, ilmenite. ‘Recalc.’ refers to recalculated FeO and Fe2O3 based on stoichiometry of oxide phases. ‘Composites’ refers to two rock types forming a single xenolith. Equilibration T refers to clinopyroxene–orthopyroxene geothermometry (Wells, 1997).

Fig. 4.

Forsterite (Fo, mol %) variation with NiO and CaO contents (wt %) in olivine of the xenoliths from Mauna Kea volcano. Each data point represents an average value for an olivine grain acquired by 10–15 spot analyses per grain; multiple grains are plotted for some samples. For comparison, we show compositional fields for olivine in lavas and gabbros from Kilauea (Fodor & Moore, 1994), Maui (Fodor et al., 1977), Mauna Loa (Garcia et al., 1995), Mauna Kea (Baker et al., 1996), and Kahoolawe (Rudek et al., 1992; Fodor et al., 1993) volcanoes, Hawaii.

Fig. 4.

Forsterite (Fo, mol %) variation with NiO and CaO contents (wt %) in olivine of the xenoliths from Mauna Kea volcano. Each data point represents an average value for an olivine grain acquired by 10–15 spot analyses per grain; multiple grains are plotted for some samples. For comparison, we show compositional fields for olivine in lavas and gabbros from Kilauea (Fodor & Moore, 1994), Maui (Fodor et al., 1977), Mauna Loa (Garcia et al., 1995), Mauna Kea (Baker et al., 1996), and Kahoolawe (Rudek et al., 1992; Fodor et al., 1993) volcanoes, Hawaii.

Modal- and phase-layered xenoliths (Fig. 2d and f;Table 1) generally have olivine ∼0.3–1.8 mol % Fo higher in the olivine-dominant layers. As examples, modally layered gabbros A9 and A22 have the pairs Fo73.4 and Fo71.8, and Fo77 and Fo75.2, respectively, and olivine clinopyroxenite–gabbronorite composite C75 has Fo81 and Fo80.4. In gabbroic grain-size composites, however, olivine compositions are essentially the same in each layer (Fig. 2g; Table 1b).

NiO abundances in olivine correlate with Fo (Fig. 4). Ultramafic xenoliths have the highest NiO, 0.42–0.25 wt %, and gabbroic samples have lower amounts, ∼0.35–0.05 wt % NiO. Similarly, the highest CaO contents are in ultramafic xenoliths (0.06–0.20 wt %) and the lowest are in gabbroic xenoliths (0.02–0.10 wt %), where they form a flat trend when plotted against Fo (Fig. 4). The NiO values are characteristic of NiO in olivines of Hawaiian lavas and gabbros, but CaO abundances are low compared with olivine in Hawaiian lavas (Fig. 4).

Clinopyroxene

All samples have clinopyroxene, but in some gabbroic xenoliths it includes exsolved orthopyroxene and/or Fe-oxide that complicate acquisition of compositional data. For grains with lamellae, we analyzed portions that are optically clear, and our analyses, therefore, presumably represent post-exsolution compositions (Tables 2 and 3; Fig. 5).

Table 3

Olivine, clinopyroxene, orthopyroxene, and plagioclase compositions (wt %) in representative gabbroic xenoliths of Mauna Kea volcano

 C7 C17 A2 A14 A21 C8 A22 (ol) 
 
 

 

 

 

 

 

 
 ol cpx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl ol cpx pl ol cpx pl 
SiO2 39.3 51.3 47.8 38.7 51.8 54.8 51.2 38.9 52.2 55.4 47.5 38.6 51.6 54.7 48.5 38.1 51.6 53.6 49.9 38.2 51.4 49.9 37.9 51.1 48.5 
TiO2  0.81   0.96 0.4   0.61 0.78   0.89 0.41   0.86 0.38     0.98  
Al2O3  3.5 33.4  1.6 31.3  2.8 1.5 32.9  1.5 32.7  2.9 1.4 31.5  3.9 31.3  3.7 33.4 
Cr2O3  0.92   0.57 0.31   0.84 0.36   0.64 0.22   0.47 0.16   0.69   0.68  
FeO 17.3 5.6 0.29 18 5.8 11.3 0.32 18.1 5.5 11.5 0.29 19.3 5.9 11.9 0.38 20.4 6.3 14.2 0.23 21.2 6.8 0.31 21.4 6.5 0.28 
MnO 0.25 0.12  0.26 0.12 0.18  0.16 0.17 0.23  0.18 0.13 0.22  0.22 0.19 0.3  0.29 0.15  0.25 0.15  
MgO 43.6 17.1  42.8 16.4 30.3  42.5 17.3 30.8  42.4 17.2 30.6  40.5 16.8 28.4  40.8 16  40.2 16.2  
CaO 0.05 19.8 16 0.06 21 1.2 13.8 0.02 20.5 0.73 15.9 0.03 20.4 0.95 15.5 0.05 21.2 0.99  0.05 20.8 14 0.05 20.7 15.9 
Na2 0.33 2.3  0.4 0.05 3.4  0.36 0.03 2.4  0.34 0.01 2.6  0.36 0.04 14.5  0.42 3.4  0.37 2.6 
K2  0.06    0.18    0.06    0.13      0.1   0.07 
NiO 0.25   0.35    0.28    0.27    0.24   0.14 0.24   0.25   
Total 100.75 99.48 99.85 100.17 100.05 100.14 100.2 99.96 100.28 101.33 99.05 100.78 100.1 100.51 99.81 99.51 100.68 99.47 99.27 100.78 101.16 99.01 100.05 100.38 100.75 
mg-no. 81.8 84.5  80.9 83.4 82.7  80.7 84.9 82.7  79.7 83.9 82.1  78 82.6 78.1  77.4 80.8  77 81.6  
Fs  9.1   9.4 16.9   8.8 17.1   9.4 17.6   9.9 21.5   11   10.5  
Wo  41.3   43.5 2.3   42 1.4   41.7 1.8   42.9 1.9   43   42.9  
An   79.1    68.5    78.3    76.2    72.2   69.1   76.9 
Or   0.35    1.06    0.35    0.76    0.83   0.59   0.4 
Equilib. T (°C)      1011   1055    1051    999         
 C7 C17 A2 A14 A21 C8 A22 (ol) 
 
 

 

 

 

 

 

 
 ol cpx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl ol cpx pl ol cpx pl 
SiO2 39.3 51.3 47.8 38.7 51.8 54.8 51.2 38.9 52.2 55.4 47.5 38.6 51.6 54.7 48.5 38.1 51.6 53.6 49.9 38.2 51.4 49.9 37.9 51.1 48.5 
TiO2  0.81   0.96 0.4   0.61 0.78   0.89 0.41   0.86 0.38     0.98  
Al2O3  3.5 33.4  1.6 31.3  2.8 1.5 32.9  1.5 32.7  2.9 1.4 31.5  3.9 31.3  3.7 33.4 
Cr2O3  0.92   0.57 0.31   0.84 0.36   0.64 0.22   0.47 0.16   0.69   0.68  
FeO 17.3 5.6 0.29 18 5.8 11.3 0.32 18.1 5.5 11.5 0.29 19.3 5.9 11.9 0.38 20.4 6.3 14.2 0.23 21.2 6.8 0.31 21.4 6.5 0.28 
MnO 0.25 0.12  0.26 0.12 0.18  0.16 0.17 0.23  0.18 0.13 0.22  0.22 0.19 0.3  0.29 0.15  0.25 0.15  
MgO 43.6 17.1  42.8 16.4 30.3  42.5 17.3 30.8  42.4 17.2 30.6  40.5 16.8 28.4  40.8 16  40.2 16.2  
CaO 0.05 19.8 16 0.06 21 1.2 13.8 0.02 20.5 0.73 15.9 0.03 20.4 0.95 15.5 0.05 21.2 0.99  0.05 20.8 14 0.05 20.7 15.9 
Na2 0.33 2.3  0.4 0.05 3.4  0.36 0.03 2.4  0.34 0.01 2.6  0.36 0.04 14.5  0.42 3.4  0.37 2.6 
K2  0.06    0.18    0.06    0.13      0.1   0.07 
NiO 0.25   0.35    0.28    0.27    0.24   0.14 0.24   0.25   
Total 100.75 99.48 99.85 100.17 100.05 100.14 100.2 99.96 100.28 101.33 99.05 100.78 100.1 100.51 99.81 99.51 100.68 99.47 99.27 100.78 101.16 99.01 100.05 100.38 100.75 
mg-no. 81.8 84.5  80.9 83.4 82.7  80.7 84.9 82.7  79.7 83.9 82.1  78 82.6 78.1  77.4 80.8  77 81.6  
Fs  9.1   9.4 16.9   8.8 17.1   9.4 17.6   9.9 21.5   11   10.5  
Wo  41.3   43.5 2.3   42 1.4   41.7 1.8   42.9 1.9   43   42.9  
An   79.1    68.5    78.3    76.2    72.2   69.1   76.9 
Or   0.35    1.06    0.35    0.76    0.83   0.59   0.4 
Equilib. T (°C)      1011   1055    1051    999         
 A22 (pl) A7 A13 A5 C22 C15 
 
 

 

 

 

 

 
 ol cpx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl pl ol cpx pl pl ol cpx pl 
SiO2 37.8 51.4 49.2 37.9 51.9 54.8 49.1 37.6 51.9 54.6 49.8 37.9 52 54.5 49.3 52.7 38.1 48.9 47.1 49 38 51.7 51 
TiO2  1.1   0.46   0.85 0.45   0.65 0.35    2.1    1.1  
Al2O3  3.3 32.7  1.5 32.4  2.6 1.4 31.8  2.6 1.3 31.3 28.5  5.2 34.1 32.2  3.3 31.5 
Cr2O3  0.49   0.36 0.15   0.29 0.15   0.28 0.11    0.02    0.32  
FeO 22.8 7.1 0.31 21.5 6.7 14.2 0.41 21.8 6.8 13.5 0.31 21.8 6.3 14.2 0.33 0.35 22 6.7 0.34 0.29 22.4 5.6 0.33 
MnO 0.26 0.17  0.2 0.15 0.18  0.26 0.11 0.18  0.26 0.13 0.23   0.42 0.12   0.28 0.13  
MgO 38.8 16.1  40.1 16.6 28.6  40.4 16.4 28.2  40.4 15.9 28.5   39.4 14.1   39.4 15.8  
CaO 0.06 21.2 15.6 0.02 20.4 0.74 15.3 0.06 20.9 1.7 14.7 0.06 21.1 1.3 14.7 11.9 0.09 21.6 17.1 15.3 0.04 22.3 14 
Na2 0.41 2.8  0.35 0.01 2.7  0.4 0.05  0.4 0.03 2.9 4.3  0.46 1.7 2.7  0.39 3.4 
K2  0.1    0.13    0.17    0.12 0.2   0.1 0.12   0.11 
NiO 0.24   0.27    0.28    0.28     0.09    0.25   
Total 99.96 101.27 100.71 99.99 100.46 100.64 100.04 100.4 100.25 100.23 99.78 100.7 99.36 100.52 98.65 97.95 100.1 99.2 100.44 99.61 100.37 100.64 100.34 
mg-no. 75.2 80.2  76.9 81.5 78.2  76.8 81.1 78.8  76.8 81.2 78.2   75.9 79   75.8 83.4  
Fs  11.3   10.7 21.5   10.8 20.5   10.2 21.3    11.2     
Wo  43.2   41.9 1.4   42.7 3.3   43.9 2.5    46.5    45.9  
An   75.1    75.3    72.3    73.6 59.8   84.3 75.3   69.1 
Or   0.57    0.76    0.99    0.71 1.2   0.59 0.7   0.65 
Equilib. T (°C)    1040    1013    985           
 A22 (pl) A7 A13 A5 C22 C15 
 
 

 

 

 

 

 
 ol cpx pl ol cpx opx pl ol cpx opx pl ol cpx opx pl pl ol cpx pl pl ol cpx pl 
SiO2 37.8 51.4 49.2 37.9 51.9 54.8 49.1 37.6 51.9 54.6 49.8 37.9 52 54.5 49.3 52.7 38.1 48.9 47.1 49 38 51.7 51 
TiO2  1.1   0.46   0.85 0.45   0.65 0.35    2.1    1.1  
Al2O3  3.3 32.7  1.5 32.4  2.6 1.4 31.8  2.6 1.3 31.3 28.5  5.2 34.1 32.2  3.3 31.5 
Cr2O3  0.49   0.36 0.15   0.29 0.15   0.28 0.11    0.02    0.32  
FeO 22.8 7.1 0.31 21.5 6.7 14.2 0.41 21.8 6.8 13.5 0.31 21.8 6.3 14.2 0.33 0.35 22 6.7 0.34 0.29 22.4 5.6 0.33 
MnO 0.26 0.17  0.2 0.15 0.18  0.26 0.11 0.18  0.26 0.13 0.23   0.42 0.12   0.28 0.13  
MgO 38.8 16.1  40.1 16.6 28.6  40.4 16.4 28.2  40.4 15.9 28.5   39.4 14.1   39.4 15.8  
CaO 0.06 21.2 15.6 0.02 20.4 0.74 15.3 0.06 20.9 1.7 14.7 0.06 21.1 1.3 14.7 11.9 0.09 21.6 17.1 15.3 0.04 22.3 14 
Na2 0.41 2.8  0.35 0.01 2.7  0.4 0.05  0.4 0.03 2.9 4.3  0.46 1.7 2.7  0.39 3.4 
K2  0.1    0.13    0.17    0.12 0.2   0.1 0.12   0.11 
NiO 0.24   0.27    0.28    0.28     0.09    0.25   
Total 99.96 101.27 100.71 99.99 100.46 100.64 100.04 100.4 100.25 100.23 99.78 100.7 99.36 100.52 98.65 97.95 100.1 99.2 100.44 99.61 100.37 100.64 100.34 
mg-no. 75.2 80.2  76.9 81.5 78.2  76.8 81.1 78.8  76.8 81.2 78.2   75.9 79   75.8 83.4  
Fs  11.3   10.7 21.5   10.8 20.5   10.2 21.3    11.2     
Wo  43.2   41.9 1.4   42.7 3.3   43.9 2.5    46.5    45.9  
An   75.1    75.3    72.3    73.6 59.8   84.3 75.3   69.1 
Or   0.57    0.76    0.99    0.71 1.2   0.59 0.7   0.65 
Equilib. T (°C)    1040    1013    985           
 C50 C18 C3 C59 C16 C49 A3 
 
 

 

 

 

 

 

 
 ol cpx opx pl ol cpx pl ol cpx opx pl ol cpx cpx pl ol cpx pl ol cpx pl ol cpx opx pl 
SiO2 38.1 51.5 54.3 50.7 37.5 49.8 51.1 37.5 50.9 53.2 48.6 37.6 50.8 50.6 49.5 37.5 51.5 51.3 37.8 53.3 52.6 37.2 51.7 54.8 52.2 
TiO2  0.92 0.48   1.3   0.33   1.1 1.1   1.1   0.43   0.73 0.28  
Al2O3  2.8 1.5 32.1  4.1 31.8  3.5 1.3 32.8  4.1 3.6 32  3.1 30.7  1.8 30.3  2.6 1.1 30.5 
Cr2O3  0.32 0.12   0.34   0.27 0.04   0.53 0.25   0.11   0.03   0.14 0.04  
FeO 22.6 6.4 13.9 0.35 23.6 7.4 0.42 23.2 7.4 16.7 0.42 23.6 6.3 0.33 24 7.1 0.39 23.8 6.6 0.33 24.6 7.3 14.9 0.32 
MnO 0.25 0.16 0.24  0.33 0.14  0.21 0.14 0.31  0.35 0.14 0.17  0.32 0.18  0.33 0.13  0.11 0.18 0.22  
MgO 39.3 15.9 28.3  39.2 15.4  38.4 15.9 26.4  38.5 15.9 15.5  38.2 15.9  37.5 16  38.5 16 28.7  
CaO 0.1 21.2 1.5 14.4 0.09 20.2 13.6 0.03 20.8 0.93 15.9 0.08 20.6 20.8 14.8 0.05 21.4 12.9 0.07 22.3 13 0.02 20.6 0.86 13.2 
Na2 0.33 0.02 3.1  0.39 3.6  0.36 0.01 2.5  0.4 0.42 2.9  0.45 4.3  0.32 3.9  0.37 0.04 3.8 
K2   0.14   0.23    0.14    0.09   0.12   0.2    0.18 
NiO 0.3    0.18   0.24    0.17    0.22   0.12   0.29    
Total 100.65 99.53 100.36 100.79 100.9 99.07 100.75 99.58 100.27 99.22 100.36 100.3 99.87 99.44 99.62 100.29 100.84 99.71 99.62 100.91 100.33 100.72 99.62 100.94 100.2 
mg-no. 75.6 81.6 78.4  74.8 78.9  74.7 79.3 73.8  74.4 81.2 79.8  73.9 80  73.8 81.2  73.6 79.6 77.5  
Fs  10.3 21   12.1   11.9 25.7   10.3 11.4   11.3   10.4   11.7 22.2  
Wo  43.9 2.9   42.7   42.7 1.8   43.3 43.5   43.6   44.9   42.5 1.6  
An    71.4   66.7    77.3    73.5   62   64.1    65.1 
Or    0.83   1.3    0.81    0.53   0.69   1.2    1.06 
Equilib. T (°C) 987       997              1000   
 C50 C18 C3 C59 C16 C49 A3 
 
 

 

 

 

 

 

 
 ol cpx opx pl ol cpx pl ol cpx opx pl ol cpx cpx pl ol cpx pl ol cpx pl ol cpx opx pl 
SiO2 38.1 51.5 54.3 50.7 37.5 49.8 51.1 37.5 50.9 53.2 48.6 37.6 50.8 50.6 49.5 37.5 51.5 51.3 37.8 53.3 52.6 37.2 51.7 54.8 52.2 
TiO2  0.92 0.48   1.3   0.33   1.1 1.1   1.1   0.43   0.73 0.28  
Al2O3  2.8 1.5 32.1  4.1 31.8  3.5 1.3 32.8  4.1 3.6 32  3.1 30.7  1.8 30.3  2.6 1.1 30.5 
Cr2O3  0.32 0.12   0.34   0.27 0.04   0.53 0.25   0.11   0.03   0.14 0.04  
FeO 22.6 6.4 13.9 0.35 23.6 7.4 0.42 23.2 7.4 16.7 0.42 23.6 6.3 0.33 24 7.1 0.39 23.8 6.6 0.33 24.6 7.3 14.9 0.32 
MnO 0.25 0.16 0.24  0.33 0.14  0.21 0.14 0.31  0.35 0.14 0.17  0.32 0.18  0.33 0.13  0.11 0.18 0.22  
MgO 39.3 15.9 28.3  39.2 15.4  38.4 15.9 26.4  38.5 15.9 15.5  38.2 15.9  37.5 16  38.5 16 28.7  
CaO 0.1 21.2 1.5 14.4 0.09 20.2 13.6 0.03 20.8 0.93 15.9 0.08 20.6 20.8 14.8 0.05 21.4 12.9 0.07 22.3 13 0.02 20.6 0.86 13.2 
Na2 0.33 0.02 3.1  0.39 3.6  0.36 0.01 2.5  0.4 0.42 2.9  0.45 4.3  0.32 3.9  0.37 0.04 3.8 
K2   0.14   0.23    0.14    0.09   0.12   0.2    0.18 
NiO 0.3    0.18   0.24    0.17    0.22   0.12   0.29    
Total 100.65 99.53 100.36 100.79 100.9 99.07 100.75 99.58 100.27 99.22 100.36 100.3 99.87 99.44 99.62 100.29 100.84 99.71 99.62 100.91 100.33 100.72 99.62 100.94 100.2 
mg-no. 75.6 81.6 78.4  74.8 78.9  74.7 79.3 73.8  74.4 81.2 79.8  73.9 80  73.8 81.2  73.6 79.6 77.5  
Fs  10.3 21   12.1   11.9 25.7   10.3 11.4   11.3   10.4   11.7 22.2  
Wo  43.9 2.9   42.7   42.7 1.8   43.3 43.5   43.6   44.9   42.5 1.6  
An    71.4   66.7    77.3    73.5   62   64.1    65.1 
Or    0.83   1.3    0.81    0.53   0.69   1.2    1.06 
Equilib. T (°C) 987       997              1000   
 C10c C10f A9 (ol) A9 (pl) A10 C6 C14 A15 
 
 

 

 

 

 

 

 

 
 ol cpx pl ol cpx pl ol cpx pl ol cpx pl ol cpx opx pl ol cpx pl ol cpx pl ol cpx pl 
SiO2 37.1 51.6 50.8 37.6 51.1 50.3 38.2 51 51.4 37.4 50.6 51 37.2 51.6 54 51.7 37.2 52.1 51.3 37.4 52.1 52.5 37.2 49.6 50.7 
TiO2  1.1   1.1   0.96   1.1   0.82 0.48   0.7   0.63   1.3  
Al2O3  3.4 31.4  3.3 31.8  3.2 30.8  3.6 31.2  2.6 1.4 30.4  2.6 30.3  2.2 30  31.8 
Cr2O3  0.22   0.16   0.15   0.15   0.37 0.13   0.05   0.05  25.7 0.09  
FeO 24.2 7.6 0.38 23.9 7.7 0.37 24.1 7.2 0.43 25.4 7.5 0.36 25.4 7.1 14.6 0.42 25.6 7.2 0.47 25.5 6.8 0.48 0.3 7.7 0.39 
MnO 0.29 0.15  0.29 0.17  0.32 0.19  0.35 0.16  0.39 0.2 0.22  0.26 0.16  0.32 0.18  37 0.14  
MgO 37.5 15.9  37.5 16.2  37.3 15.9  36.2 15.4  37.3 16.2 27.9  37.4 15.8  37.2 16.1  0.09 15.1  
CaO 0.05 20.9 14.1 0.06 20.5 14.6 0.04 20.8 13.3 0.04 20.6 13.7 0.03 20.7 1.6 13.5 0.05 21.8 13.7 0.05 21.4 12.5  21.2 14.4 
Na2 0.39 3.3  0.35 3.2  0.38 3.6  0.4 3.3  0.4 0.01 3.6  0.33 3.6  0.29 4.1  0.42 3.2 
K2  0.13   0.06   0.16   0.14    0.22   0.19   0.2   0.19 
NiO 0.17   0.23   0.19   0.19   0.15    0.12   0.13   0.13   
Total 99.31 101.26 100.11 99.58 100.58 100.33 100.15 99.78 99.69 99.58 99.51 99.7 100.47 99.99 100.34 99.84 100.63 100.74 99.56 100.6 99.75 99.78 100.42 99.55 100.68 
mg-no. 73.4 79  73.6 79.1  73.4 79.3  71.8 78.5  72.4 80.3 77.3  72.3 79.6  72.3 80.9  72 77.8  
Fs  12.2   12.2   11.9   12.2   11.4 22   11.4   10.8   12.5  
Wo  41.8   42.6   42.7   43.1   42.5 3.1   44.2   43.6   44  
An   69.8   71.4   66.5   69.1    66.6   67.1   62.1   70.6 
Or   0.76   0.35   0.95   0.84    1.3   1.1   1.2   1.1 
Equilib. T (°C)             1005            
 C10c C10f A9 (ol) A9 (pl) A10 C6 C14 A15 
 
 

 

 

 

 

 

 

 
 ol cpx pl ol cpx pl ol cpx pl ol cpx pl ol cpx opx pl ol cpx pl ol cpx pl ol cpx pl 
SiO2 37.1 51.6 50.8 37.6 51.1 50.3 38.2 51 51.4 37.4 50.6 51 37.2 51.6 54 51.7 37.2 52.1 51.3 37.4 52.1 52.5 37.2 49.6 50.7 
TiO2  1.1   1.1   0.96   1.1   0.82 0.48   0.7   0.63   1.3  
Al2O3  3.4 31.4  3.3 31.8  3.2 30.8  3.6 31.2  2.6 1.4 30.4  2.6 30.3  2.2 30  31.8 
Cr2O3  0.22   0.16   0.15   0.15   0.37 0.13   0.05   0.05  25.7 0.09  
FeO 24.2 7.6 0.38 23.9 7.7 0.37 24.1 7.2 0.43 25.4 7.5 0.36 25.4 7.1 14.6 0.42 25.6 7.2 0.47 25.5 6.8 0.48 0.3 7.7 0.39 
MnO 0.29 0.15  0.29 0.17  0.32 0.19  0.35 0.16  0.39 0.2 0.22  0.26 0.16  0.32 0.18  37 0.14  
MgO 37.5 15.9  37.5 16.2  37.3 15.9  36.2 15.4  37.3 16.2 27.9  37.4 15.8  37.2 16.1  0.09 15.1  
CaO 0.05 20.9 14.1 0.06 20.5 14.6 0.04 20.8 13.3 0.04 20.6 13.7 0.03 20.7 1.6 13.5 0.05 21.8 13.7 0.05 21.4 12.5  21.2 14.4 
Na2 0.39 3.3  0.35 3.2  0.38 3.6  0.4 3.3  0.4 0.01 3.6  0.33 3.6  0.29 4.1  0.42 3.2 
K2  0.13   0.06   0.16   0.14    0.22   0.19   0.2   0.19 
NiO 0.17   0.23   0.19   0.19   0.15    0.12   0.13   0.13   
Total 99.31 101.26 100.11 99.58 100.58 100.33 100.15 99.78 99.69 99.58 99.51 99.7 100.47 99.99 100.34 99.84 100.63 100.74 99.56 100.6 99.75 99.78 100.42 99.55 100.68 
mg-no. 73.4 79  73.6 79.1  73.4 79.3  71.8 78.5  72.4 80.3 77.3  72.3 79.6  72.3 80.9  72 77.8  
Fs  12.2   12.2   11.9   12.2   11.4 22   11.4   10.8   12.5  
Wo  41.8   42.6   42.7   43.1   42.5 3.1   44.2   43.6   44  
An   69.8   71.4   66.5   69.1    66.6   67.1   62.1   70.6 
Or   0.76   0.35   0.95   0.84    1.3   1.1   1.2   1.1 
Equilib. T (°C)             1005            
 C4 A2-2 B4 C1c C1f C19c C19f 
 
 

 

 

 

 

 

 
 ol cpx opx pl ol cpx opx pl ol cpx opx pl cpx pl cpx pl ol cpx opx pl ol cpx opx pl 
SiO2 37 51.6 52.2 52.1 36.2 51.5 53.4 52.8 37 51.5 54.1 52.3 50.8 51.5 51.1 49.8 36.2 51.8 52.3 52.2 35.8 51.7 53.5 51.4 
TiO2  0.76 0.35   0.74 0.39   0.87 0.29  0.97  0.94   0.72 0.24   0.64 0.32  
Al2O3  0.27 1.2 30.6  2.4 1.4 30.3  2.8 1.3 30.8 3.2 30.9 3.1 32  2.7 1.1 30.3  2.3 1.3 30.6 
Cr2O3  0.03 0.03   0.06 0.03   0.08 0.01       0.01 0.01   0.01 0.01  
FeO 26.6 7.5 18.2 0.38 28.2 7.9 16.8 0.34 28.8 7.8 16.8 0.31 7.5 0.4 7.6 0.45 33 7.6 20 0.28 32.8 7.9 19.1 0.25 
MnO 0.34 0.18 0.36  0.46 0.23 0.42  0.41 0.14 0.26  0.21  0.21  0.45 0.37 0.43  0.49 0.35 0.39  
MgO 35.9 15.5 25.8  35.8 15.5 26.5  34.8 15.5 26.5  15.9  15.9  29.9 15.5 24.6  29.2 15.6 24.7  
CaO 0.03 21.3 0.99 12.7 0.04 19.9 1.2 12 0.11 21.1 1.2 12.8 20.7 13.2 20.9 14.6 0.06 21.1 0.86 12.6 0.11 21 1.1 12.9 
Na2 0.38 0.05  0.35 0.03 4.3  0.36 0.01 0.38 3.7 0.37 3.1  0.34 0.01  0.35 0.02 3.8 
K2   0.14    0.17    0.22  0.17  0.1    0.17    0.16 
NiO 0.11    0.13    0.08        0.08    0.09    
Total 99.98 99.95 99.18 99.92 100.83 98.58 100.17 99.91 101.2 100.15 100.47 100.43 99.66 99.87 100.12 100.05 99.69 100.14 99.55 99.55 98.49 99.85 100.44 99.11 
mg-no. 70.6 78.7 71.7  69.4 77.8 73.8  68.3 78 73.7  79.1  78.8  61.8 78.4 68.7  61.4 77.9 69.6  
Fs  12 27.8   12.9 25.6   12.5 25.6  12  12.1   12.2 30.8   12.6 29.6  
Wo  43.8 1.9   41.8 2.4   43.3 2.4  42.6  42.7   43.5 1.7   43 2.2  
An    63.2    60.1    63.1  65.7  71.9    62.9    64.7 
Or    0.83    1.01    1.3  1.01  0.59    1.01    0.95 
Equilib. T (°C) 953    1024    974        966    972   
 C4 A2-2 B4 C1c C1f C19c C19f 
 
 

 

 

 

 

 

 
 ol cpx opx pl ol cpx opx pl ol cpx opx pl cpx pl cpx pl ol cpx opx pl ol cpx opx pl 
SiO2 37 51.6 52.2 52.1 36.2 51.5 53.4 52.8 37 51.5 54.1 52.3 50.8 51.5 51.1 49.8 36.2 51.8 52.3 52.2 35.8 51.7 53.5 51.4 
TiO2  0.76 0.35   0.74 0.39   0.87 0.29  0.97  0.94   0.72 0.24   0.64 0.32  
Al2O3  0.27 1.2 30.6  2.4 1.4 30.3  2.8 1.3 30.8 3.2 30.9 3.1 32  2.7 1.1 30.3  2.3 1.3 30.6 
Cr2O3  0.03 0.03   0.06 0.03   0.08 0.01       0.01 0.01   0.01 0.01  
FeO 26.6 7.5 18.2 0.38 28.2 7.9 16.8 0.34 28.8 7.8 16.8 0.31 7.5 0.4 7.6 0.45 33 7.6 20 0.28 32.8 7.9 19.1 0.25 
MnO 0.34 0.18 0.36  0.46 0.23 0.42  0.41 0.14 0.26  0.21  0.21  0.45 0.37 0.43  0.49 0.35 0.39  
MgO 35.9 15.5 25.8  35.8 15.5 26.5  34.8 15.5 26.5  15.9  15.9  29.9 15.5 24.6  29.2 15.6 24.7  
CaO 0.03 21.3 0.99 12.7 0.04 19.9 1.2 12 0.11 21.1 1.2 12.8 20.7 13.2 20.9 14.6 0.06 21.1 0.86 12.6 0.11 21 1.1 12.9 
Na2 0.38 0.05  0.35 0.03 4.3  0.36 0.01 0.38 3.7 0.37 3.1  0.34 0.01  0.35 0.02 3.8 
K2   0.14    0.17    0.22  0.17  0.1    0.17    0.16 
NiO 0.11    0.13    0.08        0.08    0.09    
Total 99.98 99.95 99.18 99.92 100.83 98.58 100.17 99.91 101.2 100.15 100.47 100.43 99.66 99.87 100.12 100.05 99.69 100.14 99.55 99.55 98.49 99.85 100.44 99.11 
mg-no. 70.6 78.7 71.7  69.4 77.8 73.8  68.3 78 73.7  79.1  78.8  61.8 78.4 68.7  61.4 77.9 69.6  
Fs  12 27.8   12.9 25.6   12.5 25.6  12  12.1   12.2 30.8   12.6 29.6  
Wo  43.8 1.9   41.8 2.4   43.3 2.4  42.6  42.7   43.5 1.7   43 2.2  
An    63.2    60.1    63.1  65.7  71.9    62.9    64.7 
Or    0.83    1.01    1.3  1.01  0.59    1.01    0.95 
Equilib. T (°C) 953    1024    974        966    972   

Prefixes A, B, and C in sample numbers refer to locations of sample sites–cones shown in Fig. 1. ol, olivine; cpx, clinopyroxene; opx, orthopyroxene; pl, plagioclase. For modally layered and grain-size layered xenoliths, phases in each layer are presented with mineral reference to the dominant phase (ol or pl) or to grain-size (f, fine; c, coarse). Equilibrium T refers to clinopyroxene–orthopyroxene geothermometry (Wells, 1997).

Fig. 5.

Clinopyroxene mg-number [100×atomic Mg/(Mg+Fe)] plotted against various clinopyroxene components (wt %) and coexisting olivine forsterite (Fo) values for xenoliths from cones on the southern flank of Mauna Kea volcano (Fig. 1). (a) Each data point represents an average value for a clinopyroxene grain acquired by 10–15 spot analyses per grain; multiple grains are plotted for most samples. For comparison, we include compositional fields for clinopyroxene from Hawaiian tholeiitic magmas represented by Kilauea volcano lavas and gabbros (data from Moore et al., 1980; Ho & Garcia, 1988; Nichols & Stout, 1988; Garcia et al., 1989, 1992; Helz & Wright, 1992; Fodor & Moore, 1994), and for clinopyroxene in Hawaiian alkalic basalts (Fodor et al., 1975; Beeson, 1976; Chen et al., 1990; Frey et al., 1990). Data points for selected xenoliths are encircled. (b) Intra-sample compositional zoning for clinopyroxene in selected xenoliths to illustrate the ways by which Al2O3 varies with respect to mg-number; C1 is a grain-size composite; small circles represent the fine-grained portion.

Fig. 5.

Clinopyroxene mg-number [100×atomic Mg/(Mg+Fe)] plotted against various clinopyroxene components (wt %) and coexisting olivine forsterite (Fo) values for xenoliths from cones on the southern flank of Mauna Kea volcano (Fig. 1). (a) Each data point represents an average value for a clinopyroxene grain acquired by 10–15 spot analyses per grain; multiple grains are plotted for most samples. For comparison, we include compositional fields for clinopyroxene from Hawaiian tholeiitic magmas represented by Kilauea volcano lavas and gabbros (data from Moore et al., 1980; Ho & Garcia, 1988; Nichols & Stout, 1988; Garcia et al., 1989, 1992; Helz & Wright, 1992; Fodor & Moore, 1994), and for clinopyroxene in Hawaiian alkalic basalts (Fodor et al., 1975; Beeson, 1976; Chen et al., 1990; Frey et al., 1990). Data points for selected xenoliths are encircled. (b) Intra-sample compositional zoning for clinopyroxene in selected xenoliths to illustrate the ways by which Al2O3 varies with respect to mg-number; C1 is a grain-size composite; small circles represent the fine-grained portion.

Average mg-numbers for both interstitial and granular clinopyroxene in plagioclase-free (largely porphyroclastic) dunite, wehrlite, and olivine clinopyroxenite xenoliths are in the range 90.3–83.6, with dunite occupying the high end of that range, except for ‘evolved’ dunite A18 which has clinopyroxene beyond the range at mg-number 80.9. Plagioclase-bearing wehrlite and olivine clinopyroxenite have a lower range of clinopyroxene mg-numbers, 86–81.6, and gabbroic xenoliths have mg-numbers 84.9–77.3, where the fine-grained variety provides the majority of mg-numbers <80. Intra- and inter-grain variations are generally ∼1–2 mg-number units but can be up to ∼7 mg-number units (Fig. 5b).

Among modal- and phase-layered xenoliths, average mg-numbers in juxtaposed layers can be essentially identical or differ by up to ∼1 mg-number unit, where olivine-dominant layers have the higher mg-numbers (Table 1a and b). For example, dunite–olivine clinopyroxenite C25 has mg-numbers 86.6 and 86.3, olivine clinopyroxenite–gabbronite C75 has mg-numbers 84 and 82.7, and olivine-rich–plagioclase-rich portions in gabbro A22 have clinopyroxene mg-numbers 76.9 and 75.9.

Clinopyroxene mg-numbers plotted against coexisting olivine Fo values form a continuum between ultramafic and gabbroic xenoliths (Fig. 5a). However, clinopyroxene mg-numbers are higher than coexisting Fo values, and differences between the coexisting values increase from ∼2 to 8 mg-number units from ultramafic to gabbroic xenoliths.

Average Al2O3 contents in clinopyroxene can be viewed as two subgroups (Fig. 5a). (1) One has Al2O3 less than ∼3 wt %, and its clinopyroxene often coexists with orthopyroxene; compositions plot from mg-number ∼90 into and across the reference field for clinopyroxene in Kilauea tholeiitic lava. Troctolite, which does not have orthopyroxene, belongs to this subgroup by its low clinopyroxene Al2O3, ∼2 wt %, which is among the lowest known for Hawaiian clinopyroxene. (2) The second subgroup has clinopyroxene Al2O3 >3 wt %, no associated orthopyroxene, and plots above the tholeiitic reference field toward or in the alkalic basalt reference field. The ∼6 wt % Al2O3 of gabbro C22 is an extreme example of this ‘alkalic’ subgroup.

Only some clinopyroxenes show significant intra- and inter-grain Al2O3 variations. These are usually increases in Al2O3 with decreasing mg-number, but there are also examples of varying Al2O3 and nearly constant mg-number (e.g. C25c), and the reverse relationship (e.g. C18) (Fig. 5b).

Clinopyroxenes of all xenolith types collectively form an essentially flat field for TiO2 with decreasing mg-number, where gabbronorites and troctolites overall have lower TiO2 than gabbros. Compositions compare with those for Kilauea (tholeiitic) clinopyroxene, but >1.8 wt % TiO2 in clinopyroxene of C22 (Fig. 5a) is ‘alkalic’. Clinopyroxene Cr2O3 abundances among all xenoliths correlate positively with clinopyroxene mg-numbers, where Cr detection limits are reached at mg-number ∼77 (Fig. 5a).

Orthopyroxene-bearing xenoliths have clinopyroxene Na2O ∼0.3–0.4 wt %, and most gabbro xenoliths have Na2O >0.4 wt %; the highest is in C22, 0.45–0.60 wt %. Ultramafic xenoliths without orthopyroxene have a clinopyroxene Na2O range equal to that of all gabbroic xenoliths. Collectively, Na2O abundances are higher than the ∼0.20–0.35 wt % typically found in clinopyroxene of Hawaiian tholeiitic lavas and instead resemble values characteristic of clinopyroxene in alkalic basalt (Fig. 5a).

Figure 6A shows the average wollastonite (Wo) and ferrosilite (Fs) components in clinopyroxenes. Ultramafic and gabbroic xenoliths form a field of Wo decreasing with increasing Fs. Only gabbro C22 is distinct by its relatively high Wo, ∼46 mol %.

Fig. 6.

(a) Variation between clinopyroxene ferrosilite (Fs) and wollastonite (Wo) molecular end-members in xenoliths from Mauna Kea volcano. Data points are average values, one representative per xenolith (Tables 2 and 3), except for multiple grains for C22. (b) Average mg-number [100×atomic Mg/(Mg+Fe)] values in coexisting clinopyroxene and orthopyroxene in plagioclase-bearing olivine clinopyroxenite and wehrlite, and in gabbronorite xenoliths.

Fig. 6.

(a) Variation between clinopyroxene ferrosilite (Fs) and wollastonite (Wo) molecular end-members in xenoliths from Mauna Kea volcano. Data points are average values, one representative per xenolith (Tables 2 and 3), except for multiple grains for C22. (b) Average mg-number [100×atomic Mg/(Mg+Fe)] values in coexisting clinopyroxene and orthopyroxene in plagioclase-bearing olivine clinopyroxenite and wehrlite, and in gabbronorite xenoliths.

Rare-earth element (REE) abundances for clinopyroxenes (Table 4) separated from wehrlite, olivine clinopyroxenite, and gabbroic xenoliths are plotted in Fig. 7a according to the clinopyroxene–Al2O3 subgroups (e.g. Al2O3 <3 wt % associates with orthopyroxene; Fig. 5a). The negative Ce anomaly for C21 clinopyroxene may reflect leaching owing to weathering (e.g. Fodor et al., 1992, 1994). All patterns have shapes characteristic of clinopyroxene (e.g. Irving & Frey, 1984), and, in general, clinopyroxenes in the low-Al2O3 subgroup have lower REE abundances.

Table 4

Trace-element abundances (parts per million) in clinopyroxene and plagioclase grains separated from ultramafic and gabbroic xenoliths of Mauna Kea and from a lava of Mauna Loa volcano

 Clinopyroxene Plagioclase 
 
 

 
 Flank-cone xenoliths Summit-cone xenoliths Flank-cone xenoliths Mauna Loa lava 
 
 

 

 

 
 C46 A27w A45 C25c C78 C20 A14 A15 C18 C21 C22 MKS2B MKS9 MKS14 A5 A7 C49  
La 3.9 1.7 2.7 1.7 1.25 1.7 2.36 2.4 2.1 3.72 2.5 1.8 8.5 1.87 1.41 
Ce  7.3   6.4 7.3 15.5 10  11.3 25.9 6.7 3.6 1.9 1.5 
Nd   9.3  16.1 7.5  7.2  19.6 2.9 1.7  0.6 
Sm 2.39 1.75 2.13 2.6 1.88 1.85 2.54 4.37 3.7 2.79 5.27 2.31 2.59 7.51 0.27 0.25 0.16 0.21 
Eu 0.83 0.65 0.76 0.85 0.65 0.65 0.94 1.34 1.24 0.96 1.67 0.9 0.98 2.27 0.99 0.55 0.42 0.48 
Tb 0.25 0.33 0.36 0.32 0.35 0.33 0.52 0.79 0.55 0.5 0.91 0.41 0.5 1.33 0.05 0.06 0.06 0.03 
Yb 0.9 1.1 1.1 1.1 0.9 1.3 1.32 1.29 1.04 1.5 0.87 1.22 2.28 0.1   0.09 
Lu       0.24 0.23 0.29 0.3 0.25 0.13 0.16 0.25 0.01 0.03   
Sc 77.7 65 66.6 80.9 68.4 70.4 69.3 85.1 85.2 64.9 71.1 83.7 86.5 63.9     
Hf 1.5 0.7 1.2 1.4 0.94 1.2 0.99 2.69 1.9 1.4 4.01 1.4 1.6 5.6  0.07 0.06 0.09 
Sr               707 562 721 755 
Ba               99 76 88 73 
 Clinopyroxene Plagioclase 
 
 

 
 Flank-cone xenoliths Summit-cone xenoliths Flank-cone xenoliths Mauna Loa lava 
 
 

 

 

 
 C46 A27w A45 C25c C78 C20 A14 A15 C18 C21 C22 MKS2B MKS9 MKS14 A5 A7 C49  
La 3.9 1.7 2.7 1.7 1.25 1.7 2.36 2.4 2.1 3.72 2.5 1.8 8.5 1.87 1.41 
Ce  7.3   6.4 7.3 15.5 10  11.3 25.9 6.7 3.6 1.9 1.5 
Nd   9.3  16.1 7.5  7.2  19.6 2.9 1.7  0.6 
Sm 2.39 1.75 2.13 2.6 1.88 1.85 2.54 4.37 3.7 2.79 5.27 2.31 2.59 7.51 0.27 0.25 0.16 0.21 
Eu 0.83 0.65 0.76 0.85 0.65 0.65 0.94 1.34 1.24 0.96 1.67 0.9 0.98 2.27 0.99 0.55 0.42 0.48 
Tb 0.25 0.33 0.36 0.32 0.35 0.33 0.52 0.79 0.55 0.5 0.91 0.41 0.5 1.33 0.05 0.06 0.06 0.03 
Yb 0.9 1.1 1.1 1.1 0.9 1.3 1.32 1.29 1.04 1.5 0.87 1.22 2.28 0.1   0.09 
Lu       0.24 0.23 0.29 0.3 0.25 0.13 0.16 0.25 0.01 0.03   
Sc 77.7 65 66.6 80.9 68.4 70.4 69.3 85.1 85.2 64.9 71.1 83.7 86.5 63.9     
Hf 1.5 0.7 1.2 1.4 0.94 1.2 0.99 2.69 1.9 1.4 4.01 1.4 1.6 5.6  0.07 0.06 0.09 
Sr               707 562 721 755 
Ba               99 76 88 73 

Clinopyroxene from summit-cone xenoliths (Fodor & Vandermeyden, 1988): new data, presented for comparison. Plagioclase from Mauna Loa lava: new data, presented for comparison.

Fig. 7.

(a) Rare-earth element (REE) abundance patterns for clinopyroxene and plagioclase grains from xenoliths and for whole-xenoliths (most of which are fine-grained)-all from the southern flank of Mauna Kea volcano. The plagioclase patterns include new data for plagioclase from a Mauna Loa volcano lava (tholeiitic), and the whole-rock patterns include, for reference, REE abundances of a Mauna Kea tholeiitic lava (Frey et al., 1991; sample MU-8). (b) We make comparisons with new REE data for clinopyroxenes from Mauna Kea summit-cone gabbro xenoliths (olivine gabbros 2B and 9, and opaque-oxide gabbro 14; Fodor & Vandermeyden, 1988); we also calculated four REE patterns for liquids in equilibrium with clinopyroxenes, one of which is for the relatively REE-rich clinopyroxene in opaque-oxide gabbro 14 (summit cone), and compare all ‘liquid’ patterns with those for various Mauna Kea postshield lava types-tholeiite (MU8), alkalic (LP9), and Fe–Ti (MU6), all from Frey et al., (1991).

Fig. 7.

(a) Rare-earth element (REE) abundance patterns for clinopyroxene and plagioclase grains from xenoliths and for whole-xenoliths (most of which are fine-grained)-all from the southern flank of Mauna Kea volcano. The plagioclase patterns include new data for plagioclase from a Mauna Loa volcano lava (tholeiitic), and the whole-rock patterns include, for reference, REE abundances of a Mauna Kea tholeiitic lava (Frey et al., 1991; sample MU-8). (b) We make comparisons with new REE data for clinopyroxenes from Mauna Kea summit-cone gabbro xenoliths (olivine gabbros 2B and 9, and opaque-oxide gabbro 14; Fodor & Vandermeyden, 1988); we also calculated four REE patterns for liquids in equilibrium with clinopyroxenes, one of which is for the relatively REE-rich clinopyroxene in opaque-oxide gabbro 14 (summit cone), and compare all ‘liquid’ patterns with those for various Mauna Kea postshield lava types-tholeiite (MU8), alkalic (LP9), and Fe–Ti (MU6), all from Frey et al., (1991).

Orthopyroxene

Orthopyroxene mg-numbers range from ∼83 to 69 (Table 3) and correlate positively with clinopyroxene mg-numbers (Fig. 6b). Orthopyroxene begins forming when clinopyroxene mg-numbers are ∼85 and olivine ∼Fo81. Throughout its crystallization, orthopyroxene mg-numbers remain lower by ∼3 mg-number units than the mg-numbers of coexisting clinopyroxenes.

Plagioclase

Average plagioclase compositions (Tables 2 and 3) are An87Or0.4 to An60Or1.5; zoning extends the range to ∼An50Or1.5 (Fig. 8). Wehrlite and olivine clinopyroxenite contain the most calcic plagioclase, >An81, and gabbroic xenoliths have comparatively evolved plagioclase, ∼An80–60, with the fine-grained variety <An70 (Fig. 9a). Compositional zoning within grains and variations within samples can be anywhere from ∼5 to 20 mol % An. Figure 8b shows representative zoning patterns.

Fig. 8.

Plagioclase endmember An vs Or (mol %) plots. (a) Composite diagrams for individual analysis points on multiple plagiocase grains in xenoliths collected from cones A, B, and C, southern flank of Mauna Kea volcano, compared with a compositional field of plagioclase from Kilauea volcano tholeiitic lavas and gabbros (data from Moore et al., 1980; Ho & Garcia, 1988; Garcia et al., 1989, 1992; Helz & Wright, 1992, Fodor & Moore, 1994).(b) Examples of compositional zoning per sample, sometimes extreme, as in gabbronorite A5, and sometimes limited in An but relatively extensive in Or, as in fine-grained gabbronorite C4.(c) An–Or diagrams for fine- and coarse-grained portions of grain-size composite gabbro xenoliths. In general, fine grained portions have plagioclase with higher An contents than coarse-grained portions, but there is substantial overlap.

Fig. 8.

Plagioclase endmember An vs Or (mol %) plots. (a) Composite diagrams for individual analysis points on multiple plagiocase grains in xenoliths collected from cones A, B, and C, southern flank of Mauna Kea volcano, compared with a compositional field of plagioclase from Kilauea volcano tholeiitic lavas and gabbros (data from Moore et al., 1980; Ho & Garcia, 1988; Garcia et al., 1989, 1992; Helz & Wright, 1992, Fodor & Moore, 1994).(b) Examples of compositional zoning per sample, sometimes extreme, as in gabbronorite A5, and sometimes limited in An but relatively extensive in Or, as in fine-grained gabbronorite C4.(c) An–Or diagrams for fine- and coarse-grained portions of grain-size composite gabbro xenoliths. In general, fine grained portions have plagioclase with higher An contents than coarse-grained portions, but there is substantial overlap.

Spot analyses expressed collectively show that An–Or relationships in cone-A xenoliths and in most cone-C xenoliths conform to the compositional field in Kilauea tholeiitic lavas and gabbros (Fig. 8a). Some compositional points for cone-C xenoliths fall outside the Kilauea field, largely owed to plagioclase in fine-grained gabbroic xenoliths. Plagioclase there began crystallization with Or values below the Kilauea trend and resembling the low Or of mid-ocean ridge basalt (MORB) plagioclase; progressive crystallization extended the compositions toward or into the Kilauea field (e.g. C4 in Fig. 8b). There are also cone-C xenoliths with Or values greater than Kilauea Or; olivine clinopyroxenite C78 has extreme Or, ∼1.7 mol% at An65 (Fig. 8b).

Comparisons of plagioclase in fine- and coarse-grained composites show that smaller grains, by and large, have lower Or and higher An than coexisting larger grains (Fig. 8c). Plagioclase compositions in the olivine-rich and plagioclase-rich portions of modal- and phase-layered gabbroic xenoliths overlap (e.g. A9 and A22, Fig. 2f).

REE compositions for plagioclase grains separated from gabbro, gabbronorite, and troctolite are shown in Table 4 and Fig. 7a. The grains are respectively An79–68, An77–55, and An68–62. They are enriched in light REE (LREE) when compared with the REE abundances of plagioclase (An68–62) from a Mauna Loa tholeiitic lava (Fig. 7a).

Average An values in uniform (non-layered) olivine+clinopyroxene framework and plagioclase-rich xenoliths have a rough inverse correlation with modal percentages of plagioclase (Fig. 9a). Plagioclase that deviates from the apparent trend is interstitial; namely, occurrences of <5 vol. % in cone-C xenoliths are either high An (85–83 mol %) or low An (72–65 mol %) (Fig. 9a). The An–Fo plot (Fig. 9b) shows that cone-A xenoliths conform well to the An–Fo trend for Kilauea tholeiitic gabbro xenoliths, but cone-C xenoliths scatter from this Kilauea trend, owing mainly to their wide An range for interstitial plagioclase. Figure 9c shows a relationship between increasing modal plagioclase and decreasing Fo values (i.e. lowest volume of plagioclase associates with highest Fo olivine, and the reverse).

Fig. 9.

(a) Plagioclase An content (mol %) plotted against plagioclase abundance in vol. % in Mauna Kea uniform (non-layered) xenoliths according to collection sites (cones A, B, C). Approximately 35 vol. % marks the change from olivine+clinopyroxene frameworks to plagioclase frameworks. There is a rough correlation between modal percentages and An contents; C3 is removed largely because of its modal heterogeneity (see Fig. 2b). (b) Representative average plagioclase An and coexisting olivine Fo in xenoliths compared with a reference field for gabbroic xenoliths from the Kilauea 1960 lava (Fodor & Moore, 1994). Cone A xenoliths conform well to intergranular plagioclase–olivine of Kilauea gabbro, but cone C data are 'scattered’ owing largely to varying compositions of interstitial plagioclase. (c) Representative olivine Fo content plotted against plagioclase abundance per xenolith demonstrates a rough correlation; not all samples are represented, because of oxidized, unanalyzable olivine in some.

Fig. 9.

(a) Plagioclase An content (mol %) plotted against plagioclase abundance in vol. % in Mauna Kea uniform (non-layered) xenoliths according to collection sites (cones A, B, C). Approximately 35 vol. % marks the change from olivine+clinopyroxene frameworks to plagioclase frameworks. There is a rough correlation between modal percentages and An contents; C3 is removed largely because of its modal heterogeneity (see Fig. 2b). (b) Representative average plagioclase An and coexisting olivine Fo in xenoliths compared with a reference field for gabbroic xenoliths from the Kilauea 1960 lava (Fodor & Moore, 1994). Cone A xenoliths conform well to intergranular plagioclase–olivine of Kilauea gabbro, but cone C data are 'scattered’ owing largely to varying compositions of interstitial plagioclase. (c) Representative olivine Fo content plotted against plagioclase abundance per xenolith demonstrates a rough correlation; not all samples are represented, because of oxidized, unanalyzable olivine in some.

Oxides

Each xenolith has at least one oxide phase-Cr-spinel (Cr2O3 >30 wt %), Cr-magnetite (Cr2O3 ∼10–30 wt %), Ti-magnetite (Cr2O3 <10 wt %), or ilmenite (Tables 2 and 5). Oxide-phase mg-numbers correlate with Fo in coexisting olivine (Fig. 10a). Plotting oxide mg-number against oxide cr-number, however, shows a ‘dogleg’ trend where cr-numbers remain essentially flat until mg-numbers decrease to <25 in the Cr- and Ti-magnetites of gabbroic xenoliths (with or without ilmenite present) (Fig. 10a).

Fig. 10.

(a) Average mg-numbers [100×Mg/(Mg+Fe)] for oxide phases in gabbros from Mauna Kea volcano, plotted against coexisting olivine Fo contents and oxide cr-numbers [100×Cr/(Cr+Al)]. Data are compared with Cr-rich oxide phases in Kilauea and Mauna Loa tholeiitic lavas and gabbros (after Evans & Wright, 1972; Nicholls & Stout, 1988; Wilkinson & Hensel, 1988; Scowen et al., 1991; Fodor & Moore, 1994; Garcia et al., 1995). (b) Negative correlation between Al2O3 in clinopyroxene and cr-number of coexisting spinel in porphyroclastic ultramafic xenoliths.

Fig. 10.

(a) Average mg-numbers [100×Mg/(Mg+Fe)] for oxide phases in gabbros from Mauna Kea volcano, plotted against coexisting olivine Fo contents and oxide cr-numbers [100×Cr/(Cr+Al)]. Data are compared with Cr-rich oxide phases in Kilauea and Mauna Loa tholeiitic lavas and gabbros (after Evans & Wright, 1972; Nicholls & Stout, 1988; Wilkinson & Hensel, 1988; Scowen et al., 1991; Fodor & Moore, 1994; Garcia et al., 1995). (b) Negative correlation between Al2O3 in clinopyroxene and cr-number of coexisting spinel in porphyroclastic ultramafic xenoliths.

The cluster of oxide-phase data with mg-number >40 (i.e. Cr-spinel) represents the plagioclase-free (largely porphyroclastic) ultramafic xenoliths, many of which overlap the compositional reference field for Cr-rich oxide phases in Hawaiian tholeiitic lavas and gabbroic xenoliths (Fig. 10a). Others fall beneath that tholeiitic field. Accordingly, Fig. 10b shows that, in general, lower spinel cr-numbers among plagioclase-free ultramafic xenoliths equate with higher Al2O3 in coexisting clinopyroxene. This relationship diminishes at oxide mg-number ∼40, or when the oxide-phase compositions approach closer to that of titaniferous magnetite; that is, there are also relatively low cr-numbers in the Cr-magnetite of some orthopyroxene-bearing ultramafic xenoliths (Fig. 10a and b).

Among composites, dunite and wehrlite layers of A27 have spinel mg-numbers of 64.5 and 62.9, respectively. In modally layered gabbro A22 (Fig. 2f), olivine-rich and plagioclase-rich portions have variable, but distinctive, Cr-magnetite mg-numbers, 31–24 and 24–21, respectively.

Whole-Rock Compositions

Compositions of selected xenoliths are shown in Table 6 and Fig. 11. Loss-on-ignition determinations yield weight gains for most xenoliths owing to Fe oxidation during ignition that offset any volatile loss; this suggests that the xenoliths, overall, are ‘fresh’. For xenoliths that lost weight during ignition, such as those with iddingsite (e.g. C1), the amounts were <0.75 wt %.

Fig. 11.

Whole-rock MgO variation diagrams for selected xenoliths from the southern flank of Mauna Kea volcano compared with whole-rock compositions of selected Mauna Kea lavas (data from Frey et al., 1990, 1991).

Fig. 11.

Whole-rock MgO variation diagrams for selected xenoliths from the southern flank of Mauna Kea volcano compared with whole-rock compositions of selected Mauna Kea lavas (data from Frey et al., 1990, 1991).

Major-element abundances have little overlap with major-element abundances in Mauna Kea lavas, and have a wider range. Al2O3, CaO, and TiO2 contents increase with decreasing MgO. The increase in Al2O3 is smooth, but the change in TiO2 is gradual until ∼10–15 wt % MgO, where it increases without regard for MgO. Incompatible elements K, Zr, and Y have notably lower abundances than Mauna Kea lavas, and increase only slightly with decreasing MgO. In contrast, Sr contents are high, and, like Al2O3, reflect modal plagioclase.

Among REE patterns (Fig. 7; Table 4), dunite A18 is LREE enriched, unusual for dunite except that it has apatite-bearing magnetite veins among its Fo77 olivine. Gabbronorite A2 with Fo80.7 has LREE depletion and a positive Eu anomaly. Fine-grained gabbroic xenoliths (C1f,c; C19, B3, A2–2), all compositionally evolved (Fo72), have essentially flat patterns with positive Eu anomalies. The pattern for the coarse portion of grain-size composite C1 (C1c) is subparallel to that for the fine portion (C1f), but lower to perhaps reflect a slightly different mode (Table 1b).

Isotopic Compositions (mineral and Whole Rock)

The 87Sr/86Sr isotopic ratios for one cone-A and one cone-C gabbro are 0.703614±22 and 0.703556±9, respectively-values characteristic of Mauna Kea postshield tholeiitic lavas (Kennedy et al., 1991) and slightly less than 87Sr/86Sr for Mauna Kea shield lavas (Yang et al., 1994) (Table 7). Pb isotope ratios for one fine-grained cone-B gabbronorite and for clinopyroxene separates from two cone-C and one cone-A gabbroic xenoliths are like those for Mauna Kea postshield lavas (Kennedy et al., 1991) (Table 7). This is true even to the extent of a notably high 207Pb/204Pb ratio of 15.630 because Pb isotope data for Mauna Kea lavas include, too, an unusually high 207Pb/204Pb of 15.550 (Kennedy et al., 1991).

δ18O values for clinopyroxene and plagioclase from A and C xenoliths are 4.05–5.62 (Table 7), compatible with the range observed for Mauna Kea lavas (Eiler et al., 1996), and, in general, with basaltic samples unaffected by secondary processes. However, they are lower than the 6–6.1 δ18O reported for Mauna Kea xenoliths by Kyser et al., (1982). There is oxygen-isotope disequilibrium between most coexisting clinopyroxene and plagioclase (Table 7), which introduces some doubt about the meaning of whole-xenolith oxygen isotope analyses (e.g. Kyser et al., 1982).

Discussion

Xenolith sources

The Sr and Pb isotopic compositions of xenoliths and mineral separates resemble those of Mauna Kea lavas (Kennedy et al., 1991). Additionally, compositions for olivine, clinopyroxene, orthopyroxene, plagioclase, and oxides are consistent with the compositions of minerals in Hawaiian tholeiitic and alkalic lavas (e.g. Keil et al., 1972; Fodor et al., 1975, 1977, 1993, 1994; Moore et al., 1980; Ho & Garcia, 1988; Garcia et al., 1989, 1992, 1995; Fodor & Moore, 1994; Clague et al., 1995). Comparatively, the mg-numbers of the xenolith minerals are lower than those observed in upper-mantle phases. Oceanic crust beneath Hawaii is also excluded by the relatively high amounts of K2O (molecular Or) in xenolith plagioclase relative to oceanic basalt plagioclase-comparisons previously used by Bohrson & Clague, (1988), Rudek et al., (1992) and Fodor et al., (1993, 1994) to discriminate between Hawaiian and ocean-ridge sources. Although some occurrences of xenoliths contain plagioclase with low Or, the Or values are not as low as in ocean-crust plagioclase, and they are, in addition, associated with phases characteristic of Hawaiian tholeiite. We conclude that the xenoliths from the southern flank of Mauna Kea (Fig. 1) have origins in Mauna Kea magmas.

Magma parentage: tholeiitic or alkalic?

Frey et al., (1990, 1991) pointed out that Mauna Kea stratigraphy and magma types are (1) shield tholeiites, (2) postshield Hamakua Volcanics, which are tholeiitic and alkalic basalts and ankaramites, and alkalic-affinity Fe–Ti basalts, and (3) the youngest eruptives, the Laupahoehoe Volcanics, composed of hawaiites and mugearites that serve as hosts for the xenoliths. Both tholeiitic and alkalic magmas were therefore available as parental magmas for the xenoliths.

Orthopyroxene indicates tholeiitic parentage for xenoliths, and it follows that clinopyroxene with ∼2–3 wt % Al2O3 (and ∼0.50–1.25 wt % TiO2) in clinopyroxene coexisting with orthopyroxene is also a tholeiitic characteristic (Fig. 5a). Most plagioclase-bearing xenoliths are therefore tholeiitic. Parentage is less clear for the gabbroic xenoliths with ∼3.5–4.5 wt % Al2O3 (∼1.25–1.5 wt % TiO2) in clinopyroxene because these values match clinopyroxene in both tholeiitic and alkalic lavas of Hawaii (Fig. 5a). To the extent that these xenoliths lack orthopyroxene, we consider them to be alkalic or at least transitional to alkalic. Clinopyroxene in C22, with Al2O3 4.5–6.5 wt % (TiO2 >1.7 wt %), seems decisively alkalic (Fig. 5).

To evaluate parentages of orthopyroxene-free, plagioclase-free ultramafic xenoliths in Hawaiian lavas, Sen & Presnall, (1986), Clague, (1988), Clague & Bohrson, (1991), and Chen et al., (1992) arrived at relatively low spinel cr-numbers and high clinopyroxene Al2O3 and clinopyroxene TiO2 for discriminating alkalic from tholeiitic. Figure 10b shows this general correlation for Mauna Kea plagioclase-free, orthopyroxene-free (and largely porphyroclastic) ultramafic xenoliths. That is, this relationship, taken with clinopyroxene Al2O3mg-number correlations (Fig. 5a), identifies these particular ultramafic xenoliths as collectively having both tholeiitic and alkalic parentages.

Other characteristics to evaluate for establishing parentages are clinopyroxene Na2O, Wo, and REE contents, and plagioclase Or and REE contents. Na2O (∼0.3–0.4 wt %) in clinopyroxene of the xenoliths that we judge from Al2O3 and TiO2 to be tholeiitic is too high to correspond to Na2O in tholeiitic clinopyroxene (Fig. 5a). Although this relatively high Na2O may instead represent re-equilibration with plagioclase in plutonic environments, the notably high Na2O in C22 clinopyroxene (>0.45 wt %) can nonetheless, in a relative sense, be viewed as characteristic of alkalic parentage.

The Wo of clinopyroxene can distinguish tholeiitic from alkalic parentage for lavas (e.g. Fodor et al., 1975), but Wo values in these plutonic rocks are dubious. This is because of possible subsolidus adjustments-CaO, in this case, diffusing among clinopyroxene, orthopyroxene, and olivine (e.g. Kohler & Brey, 1990). That caveat notwithstanding, it appears (Fig. 6a) that high Wo (e.g. >45 mol %) in gabbro C22 is in accordance with its other indications for alkalic parentage. Beyond that, Wo does not discern tholeiitic from alkalic affinities among gabbroic xenoliths.

Evaluation of REE abundances in clinopyroxenes (Table 4) is complicated by the wide range in partition coefficients empirically observed for clinopyroxene-liquid equilibrium. This is generally due to P, T, and compositional variables that influence partitioning, and to apparent sensitivity of REE abundances to melt compositions (Green & Pearson, 1985; Gallahan & Nielson, 1992; Gaetani & Grove, 1995). Plutonic crystallization, in particular, creates inconsistencies for REE partitioning (Cawthorn, 1996). To nonetheless make relative comparisons of parental-liquid characteristics, we calculated some clinopyroxene–equilibrium-liquid compositions using the mid-range partition coefficients of Gallahan & Nielson, (1992). In a general way, the calculated liquids (Fig. 7b) support our assessments for tholeiitic and alkalic parentages based on Al2O3 in clinopyroxenes. They show that low-REE, low-Al2O3 clinopyroxene corresponds to Mauna Kea tholeiitic lava, and that clinopyroxenes (>4 wt % Al2O3) in gabbros plot between representations for alkalic basalts and incompatible element enriched Fe–Ti basalt of Mauna Kea.

Plagioclase An–Or relationships, where Or contents by and large resemble those of Kilauea tholeiitic plagioclase (Fig. 8), also imply tholeiitic parentage for many xenoliths. Curiously, this is also true for ‘alkalic’ C22 (Fig. 8b). In contrast, some occurrences of interstitial plagioclase, such as in olivine clinopyroxenite C78, which is tholeiitic in terms of clinopyroxene composition, have unusually high Or contents (Fig. 8b). These examples suggest that Or is not necessarily a reliable indicator of magma parentage in plutonic environments where interstitial liquids can either escape crystallization sites with portions of incompatible element concentrations, or become trapped and concentrate incompatible elements.

LREE in plagioclase may also be an unreliable indicator of parentage because plagioclase from gabbronorite A5 (tholeiitic) is the most LREE enriched of those analyzed, even compared with the Mauna Loa (also tholeiitic) reference plagioclase (Fig. 7a). Its comparatively high La (Table 4) is probably due to higher Ab content rather than to any characterization for parentage because La partitioning is sensitive to plagioclase composition (e.g. plagioclase–liquid La KD for intermediate magmas is twice that for basaltic magmas; Henderson, 1982).

To summarize, most xenoliths can be decisively identified as tholeiitic by the presence of orthopyroxene, low Al2O3 (<3 wt %) and TiO2 (<1 wt %) in clinopyroxene, and REE abundances in clinopyroxene that can be assigned to Mauna Kea tholeiitic magmas. A minority of xenoliths have clinopyroxene that suggests parentage transitional to alkalic, and alkalic, based on clinopyroxene Al2O3 >3 wt %, lack of coexisting orthopyroxene, clinopyroxene REE abundances that can be assigned to basaltic liquids more enriched in REE than Mauna Kea alkalic lavas, and, among the porphyroclastic, plagioclase-free ultramafic xenoliths, relatively low spinel cr-numbers. Clinopyroxene in gabbro C22 has the 'strongest’ alkalic characteristics.

Xenoliths associated with stages of volcano development

The highly forsteritic olivines, ∼Fo89–84, in plagioclase-free ultramafic xenoliths compare with Fo-rich olivine phenocrysts (and xenocrysts) in Mauna Kea shield and postshield tholeiites (Yang et al., 1994; Garcia et al., 1995; Baker et al., 1996).The magmas that crystallized Fo89–84 olivine in the xenoliths-of both tholeiitic and alkalic compositions-would have yielded liquids during that crystallization that were appropriately parental for the plagioclase-bearing ultramafic xenoliths, and, ultimately, for the gabbroic xenoliths. Because xenoliths have both tholeiitic and alkalic affinities, and because Hamakua postshield lavas consist of tholeiitic, transitional, and alkalic varieties (e.g. Frey et al., 1991; Yang et al., 1996), and, finally, because Sr-isotopic ratios for two xenoliths agree with those of Hamakua lavas rather than shield lavas, we assign cone-A, -B, and -C xenoliths to the postshield, or Hamakua, stage of volcanism.

Comparison with other Mauna Kea xenoliths

The only other detailed assessment for Mauna Kea xenoliths addresses olivine gabbros and opaque-oxide gabbros (i.e. Fe–Ti-oxide rich) of the summit cone (Fodor & Vandermeyden, 1988). Based largely on clinopyroxene Al2O3 (∼3.0–4.5 wt %) and TiO2 abundances, those xenoliths have transitional and alkalic parentages. [Fodor & Vandermeyden, (1988) did not observe any orthopyroxene-bearing, or tholeiitic, xenoliths at the summit cone.] Accordingly, REE patterns for clinopyroxene from two summit-cone olivine gabbros conform to those of the gabbroic xenoliths from the southern flank that we report here as having transitional to alkalic parentage (Fig. 7b). We note, then, that summit-cone olivine gabbros represent the kind of magma parentage manifested by the minority of flank-cone gabbroic xenoliths.

On the other hand, REE abundances in clinopyroxene from a summit-cone opaque-oxide gabbro (Fig. 7b) are notably enriched and identify a sub-variety of magma parentage not found among flank xenoliths. Its calculated parental liquid seems most appropriate for the Fe–Ti basaltic lavas on Mauna Kea (Fig. 7b), which Frey et al., (1991) interpreted as fractionates of Hamakua alkalic basalts produced by shallow-level removal of mainly plagioclase.

Crystallization environments

Textures and whole-rock compositions of flank-location xenoliths indicate that they do not represent direct isochemical solidification of mafic liquids. Specifically, xenolith compositions do not overlap those for Mauna Kea lavas, and, in comparison, they have low incompatible element abundances. Also, dunite and wehrlite, phase and modal layering, and poikilitic textures do not conform to melt compositions, and positive Eu anomalies reflect plagioclase concentrations. Our interpretation, then, is that the xenoliths as a group represent crystallization of Hamakua magmas within Mauna Kea where crystals, to some extent, segregated from magmatic liquids and accumulated. To ascribe the xenoliths to crystallization environments, we refer to those recognized for active Hawaiian shields: central conduits that feed magmas to reservoirs beneath summit calderas, and deep and shallow dike-, sill-, and room-shaped ‘reservoirs’ extending from conduits to beneath flank rift zones (e.g. Ryan, 1988; Hoffmann et al., 1990). Whereas rift-zone magmatism may have been at comparatively lower rates during postshield stages, there is no reason to not consider that shallow magma-storage reservoirs existed beneath Mauna Kea's flanks.

Porphyroclastic (plagioclase-free) ultramafic xenoliths

The Fo89–84 kink-banded olivine and porphyroclastic textures of the plagioclase-free ultramafic xenoliths are consistent with crystallization from high-MgO magmas under relatively high pressures, stresses, and temperatures (e.g. Kirby & Green, 1980; Helz, 1987), and with subsequent recrystallization. Helz, (1987) offered conduits as the environment for kink-banded dunitic assemblages. This is because conduits probably provide sustained non-hydrostatic pressures necessary to promote kink-banding and high, thermally buffering temperatures to restrict modal assemblages to essentially Fo-rich olivine and Cr-spinel. Additionally, conduits allow concentric plating of olivine crystallizing from magmas ascending from mantle sources, and provide stress fields to strain olivine and create annealed porphyroclastic textures.

For an alternate environment, Clague & Denlinger, (1994) promoted porphyroclastic dunite formation by olivine accumulation at the bottoms of magma reservoirs. They emphasized the relative ease by which olivine settles by gravity from high-MgO magmas, and suggested that the strained dunite in the Hawaiian magma system is a consequence of cumulate flow within the volcanoes (dunite has the rheology of ice) and attending the plastic deformation. For example, 1-mm olivine in a melt of 9–16 wt % MgO would settle at 1.2–20 m/h.

Dunite origins entirely by gravity settling, however, may be oversimplification, as Tait & Jaupart, (1996) modeled and promoted in situ crystallization as the dominant method of solidification on reservoir floors. For this study, bottom-reservoir solidification zones reconcile with the sharp interfaces in layered dunite–olivine clinopyroxenite composites (Fig. 2d) because concentrations of clinopyroxene could represent liquids pressed from pores among cumulus olivine grains. Wehrlite, on the other hand, with varying proportions of mixed olivine and clinopyroxene, is, like dunite, a better candidate for gravity-settled origin.

Discerning between conduit and reservoir crystallization may not be possible from dunites alone, but dunite–olivine clinopyroxenite and dunite–wehrlite layered composites may help interpretations (Fig. 2d). That is, if conduits are thermally buffering (Helz, 1987) and sustain high temperatures, then a contrasting and suitable environment for olivine yielding to cumulus clinopyroxene is a reservoir that can accommodate a gradual drop from olivine liquidus temperature of ∼1260°C to clinopyroxene liquidus temperatures of ∼1170°C (Helz & Thornber, 1987).

Liquids crystallizing earliest olivine had FeO/MgO ∼0.72, or ∼15 wt % MgO, based on Fo89.4 in dunite B22 and assuming FeO/MgO crystal–liquid partitioning of 0.3. The full range, ∼Fo89.4–83.6, in this group of plagioclase-free (porphyroclastic) ultramafic xenoliths points to magmas achieving FeO/MgO ∼1.17, or ∼9–10 wt % MgO. Figure 13 (below) illustrates the crystallization of olivine and clinopyroxene from a parental liquid (in ternary i) to form a compacted cumulate pile of dunite, olivine clinopyroxenite, and wehrlite.

The Fo77 olivine in dunite A18 (Table 1a) suggests that some dunite formed from evolved liquids, but the apatite-bearing magnetite in this dunite suggests Fe–P metasomatism to account for its relatively Fe-rich olivine. Metasomatism may have also modified original clinopyroxene and spinel mg-numbers and enriched Na2O in the interstitial clinopyroxene (Fig. 5a). A clue to this dunite having originally had characteristic dunitic olivine (>Fo85) comes from high clinopyroxene–Cr2O3 (∼0.8 wt %) more appropriate for that associated with ∼Fo88–86 (Fig. 5a). Alternatively, Fe-rich dunite A18 represents a combination of processes: accumulated olivines from evolved basaltic liquids that were permeated by Fe–P-enriched residual fluids.

To summarize, strained, high-Fo olivines and porphyroclastic textures are most suitable with ultramafic xenoliths having formed in bottoms of reservoirs from tholeiitic and alkalic magmas having 10–15 wt % MgO. The physical processes of formation were probably a combination of gravity-settling and in situ crystallization.

Gabbroic and plagioclase-bearing ultramafic xenoliths

The correlations between clinopyroxene mg-number and Fo (Fig. 5), between oxide mg-number and Fo (Fig. 10), and between olivine NiO and Fo (Fig. 4) are all consistent with the gabbroic xenoliths representing crystal fractionation products of basaltic liquids. Plagioclase compositions are not as definitive about fractional crystallization relationships owing largely to their wide (∼20 mol %) An range for plagioclase amongst olivine grains having highest Fo. Nonetheless, Fig. 9a shows a general fractional crystallization correlation between plagioclase and olivine. Namely, plagioclase first appears in small volumes with ∼Fo82.7 (e.g. Table 1a), crystallizing, therefore, from liquids having FeO/MgO ∼1.25, or ∼9 wt % MgO; compositionally evolved plagioclase of lowest An (e.g. <An70) is in high abundance, such as in troctolite and fine-grained gabbros, and associated with low-Fo olivine, namely <Fo75 (Table 1b).

The compositional continua between minerals in the ultramafic and gabbroic xenoliths link gabbroic xenoliths to the magmas that produced ultramafic assemblages, on, as we propose, reservoir bottoms. However, the absence of strained characteristics in the gabbroic xenoliths (only a few have kinked olivine) suggests that their crystallization environments differed from those for porphyroclastic ultramafic xenoliths. On the other hand, crystallization pressures for gabbroic xenoliths did not differ enough from those for porphyroclastic ultramafic xenoliths to become apparent in the Aliv/Alvi ratios in clinopyroxene, which can be relative indicators of crystallization pressure (e.g. Wass, 1979; Fodor et al., 1995). Figure 12 shows that these ratios are essentially the same for gabbroic and porphyroclastic ultramafic xenoliths. In comparison, websterites of Hualalai, which have Mg-rich cumulus orthopyroxene formed under high pressure (e.g. lower crust; Bohrson & Clague, 1988), have lower Aliv/Alvi ratios (Fig. 12). Absence of Mg-rich cumulus orthopyroxene in the Mauna Kea xenoliths restricts their crystallization pressures to less than ∼5 kbar (e.g. Bohrson & Clague, 1988). This is within or above ocean crust because crust beneath Mauna Kea is between ∼13 and 18 km thick (Moore, 1987).

Fig. 12.

Clinopyroxene Aliv/Alvi ratios can reflect pressures of crystallization. Ratios for Mauna Kea xenoliths do not distinguish the various ultramafic and gabbroic xenoliths, but those for clinopyroxene in websterite xenoliths of Hualalai (Bohrson & Clague, 1988) have lower ratios (i.e. formed under higher pressure) (a 2:1 reference line is shown). We also show Aliv/Alvi ratios for other ultramafic xenoliths of Hawaii for comparison: Loihi dunite (Clague, 1988), Puu Waa Waa dunite of Hualalai (Clague & Bohrson, 1991), and Koolau dunite (Sen & Presnall, 1986) mainly overlap with Mauna Kea xenoliths, whereas wehrlite and dunite of Hualalai (Chen et al., 1992) have ratios both lower and overlapping with those of Mauna Kea xenoliths.

Fig. 12.

Clinopyroxene Aliv/Alvi ratios can reflect pressures of crystallization. Ratios for Mauna Kea xenoliths do not distinguish the various ultramafic and gabbroic xenoliths, but those for clinopyroxene in websterite xenoliths of Hualalai (Bohrson & Clague, 1988) have lower ratios (i.e. formed under higher pressure) (a 2:1 reference line is shown). We also show Aliv/Alvi ratios for other ultramafic xenoliths of Hawaii for comparison: Loihi dunite (Clague, 1988), Puu Waa Waa dunite of Hualalai (Clague & Bohrson, 1991), and Koolau dunite (Sen & Presnall, 1986) mainly overlap with Mauna Kea xenoliths, whereas wehrlite and dunite of Hualalai (Chen et al., 1992) have ratios both lower and overlapping with those of Mauna Kea xenoliths.

Fig. 13.

Schematic solidification zones and cumulate piles for reservoirs (regardless of whether tholeiitic or alkalic) within Mauna Kea volcano to illustrate and account for the origins of various xenoliths erupted from cones along the southern flank. Fo–An–Di ternary diagrams show liquid compositions [arrowhead and eutectic (E) positions in triangles] that yield mineral plus liquid assemblages at various stages of cumulate development. Fo–An–Di ternary i (bottom) and adjacent olivine+clinopyroxene mush and cumulate pile represent gravity-settled and in situ origins for porphyroclastic ultramafic xenoliths. Ternaries iii and iv, and adjacent mush and cumulate pile represent reservoir-margin solidification zones. Liquids ‘filter-pressed’ from spaces between solidification-zone grains migrate (wavy arrows) into reservoir main bodies of liquid to differentiate magma reservoirs; in turn, reservoir magmas progressively precipitate olivine, clinopyroxene, and plagioclase that are progressively more compositionally evolved. This produces large-scale modally and cryptically layered solidification zones. Fo-Di-SiO2 ternary v illustrates the eventual addition of orthopyroxene to some gabbroic assemblages. The compacted pile illustrates rock types formed, including fine-grained gabbros originating as late-stage liquids mobile within cumulate piles.

Fig. 13.

Schematic solidification zones and cumulate piles for reservoirs (regardless of whether tholeiitic or alkalic) within Mauna Kea volcano to illustrate and account for the origins of various xenoliths erupted from cones along the southern flank. Fo–An–Di ternary diagrams show liquid compositions [arrowhead and eutectic (E) positions in triangles] that yield mineral plus liquid assemblages at various stages of cumulate development. Fo–An–Di ternary i (bottom) and adjacent olivine+clinopyroxene mush and cumulate pile represent gravity-settled and in situ origins for porphyroclastic ultramafic xenoliths. Ternaries iii and iv, and adjacent mush and cumulate pile represent reservoir-margin solidification zones. Liquids ‘filter-pressed’ from spaces between solidification-zone grains migrate (wavy arrows) into reservoir main bodies of liquid to differentiate magma reservoirs; in turn, reservoir magmas progressively precipitate olivine, clinopyroxene, and plagioclase that are progressively more compositionally evolved. This produces large-scale modally and cryptically layered solidification zones. Fo-Di-SiO2 ternary v illustrates the eventual addition of orthopyroxene to some gabbroic assemblages. The compacted pile illustrates rock types formed, including fine-grained gabbros originating as late-stage liquids mobile within cumulate piles.

Petrographically uniform. Mineral compositional trends, modes, and textures of plagioclase-bearing xenoliths allow the use of a series of Fo–An–Di ternary diagrams (ii–iv in Fig. 13) with schematic diagrams to depict crystallization that led to cumulus mush and rock. In Fig. 13, ii–iv, ‘parental’ liquid crystallizes olivine (+Cr-oxide), then clinopyroxene, and then becomes saturated in plagioclase. The straightforward translation of this liquid line of descent in conjunction with the modes, textures, and phase compositions is to reservoir-margin solidification zones composed of olivine+clinopyroxene frameworks enclosing ‘eutectic’ composition (intercumulus) liquids. Geometries of the reservoirs are uncertain and could have ranged from dike-shaped (e.g. Kilauea; Hoffmann et al., 1990) to ‘room’-shaped extensions of central conduits. The main criterion is a boundary layer of crystallization between a wall and an accumulation of magma.

The relatively evolved 'starting’ liquid composition-estimated from olivine in plagioclase-bearing ultramafics to be ≤9 wt % MgO (C78 for tholeiitic; C30 for alkalic)-lends support to this reservoir-margin in situ crystallization model for plagioclase-bearing xenoliths. That is, magma viscosities increase with decreasing MgO (e.g. Shaw, 1972; Clague & Denlinger, 1994) and lessen the likelihood for olivine to settle to reservoir floors before capture by crystallization fronts growing from walls and ceilings (e.g. Marsh, 1996). Depending on depth (pressure) differences between cumulate piles on magma-reservoir floors and in solidification zones along walls and ceilings, this model reconciles the apparently different stress, pressure, and recrystallization histories for porphyroclastic ultramafic and gabbroic xenoliths while still allowing them common parental magmas.

Some percentages of intercumulus liquids in this in situ environment would escape solidification-zone frameworks to enter and chemically differentiate reservoir magmas (e.g. Langmuir, 1989). Some liquids, however, would ‘pool’ within incipient cumulate piles to crystallize as plagioclase-rich lenses and layers within wehrlites and olivine clinopyroxenites (Fig. 2a,b). Additionally, some intercumulus liquids would remain to create plagioclase-as observed among olivine+clinopyroxene framework gabbroic xenoliths-up to ∼35 vol. % (Fig. 3c-e). Poikilitic textures of two wehrlites (Fig. 3h) suggest that some early-formed olivine became enclosed by clinopyroxene-saturated, essentially cotectic, interstitial liquid.

The ∼20 mol % An compositional range for small-volume interstitial plagioclase in framework-textured xenoliths (Fig. 9a) may reflect the various histories for intercumulus liquids. In one case, the pore liquids had ‘continuity’ with reservoir magmas. That is, interstitial plagioclase with ‘high’ An relative to the Fo (Fig. 9a,b) of its framework olivine represents intercumulus liquid that sustained high An by diffusive ‘communication’ with the reservoir. In another case, isolated pore liquids were able to fractionate to Ab enrichment within pore spaces, and the result was interstitial plagioclase with ‘low’ An relative to Fo (Fig. 9b). High Or contents of some interstitial plagioclase, such as 1.7 mol % Or in olivine clinopyroxenite C78 (Fig. 8b), may be a product of isolated, interstitial crystallization during which pore liquid could not escape with incompatible element K.

Continuous reservoir-magma crystallization (e.g. Fig. 13, ii–iv) in advancing fronts of solidification zones (e.g. Marsh, 1996) progressively yielded more evolved olivine and clinopyroxene and greater proportions of plagioclase, also more evolved than earlier generations. Reservoir liquids eventually formed gabbroic assemblages that represented ‘eutectic’ composition liquids (Fig. 13, iv); in quaternary portrayal (e.g. Philpotts, 1990, fig. 10–34), that ‘eutectic’ is a Di–Fo–An–Ab cotectic that represents increasingly more albitic plagioclase crystallization (e.g. to form fine-grained mosaic-textured gabbros, their complementary coarse-grained gabbros in composite xenoliths, Fig. 2g, and some plagioclase-framework xenoliths, such as C16, Fig. 3f). Substantial amounts of orthopyroxene entered some assemblages, too, as tholeiitic reservoir magmas evolved toward SiO2 saturation (Fig. 13, v); for example, gabbronorite A5 (Fo76) has ∼25 vol. %. Gabbronorite C19 marks the most evolved tholeiitic magma, and gabbro A15 represents the most evolved transitional to alkalic magma (Table 1b).

Figure 9c generalizes the cryptic and modal layerings formed along reservoir margins in terms of decreasing Fo and increasing modal plagioclase to yield a continuum from plagioclase-bearing wehrlite and olivine clinopyroxenite to gabbroic zones. Figure 13 schematically shows this large-scale (e.g. in meters) in situ crystallization environment progressing with respect to mineral compositions, textures, and growth inward from a reservoir margin. Specifically, a gradually thickening solidification zone becomes modally and cryptically layered on a large scale and creates a general correlation between Fo and modal plagioclase (Fig. 9c) to reflect phase proportions according to Fo–An–Di crystallization (i.e. modes agree with expected original solidification-zone proportions).

Processes other than eutectic crystallization were apparently active because the percentages of modal plagioclase in some xenoliths are at disparity with ‘eutectic’ proportions of plagioclase+clinopyroxene+olivine, such as in troctolites (all have some clinopyroxene; Table 1b) and plagioclase-rich gabbro C3 (∼65 vol. %; Figs 2b and 9). Because Hawaiian tholeiitic magmas remain saturated in olivine and clinopyroxene while crystallizing plagioclase [e.g. multiphase crystallization (Clague et al., 1995); although orthopyroxene can ultimately supplant olivine (e.g. Fodor & Moore, 1994)], unusually plagioclase-rich xenoliths probably represent ‘mechanical’ concentrations of plagioclase (discussed in a following section).

Localized layering. We believe that plagioclase-bearing xenoliths, collectively, represent large-scale (meters) modal-cryptic layerings (summarized in Table 1a and b) of tholeiitic and alkalic solidification-zone crystallizations (e.g. ∼Fo83–70; Fig. 13). We further suggest that localized processes superimposed small-scale modal, phase, grain-size, and cryptic layerings on these assemblages. For example, the olivine clinopyroxenite of composite xenolith C75, which is layered olivine clinopyroxenite and gabbronorite (Fig. 2b), has no plagioclase, and its clinopyroxene mg-number ∼84 is slightly more than that of juxtaposed gabbronorite (∼83). These layered physical and chemical relationships may therefore represent the pore-liquid process addressed earlier, whereby cumulus olivine+clinopyroxene frameworks expel intercumulus liquids saturated in plagioclase to ‘pool’ nearby (e.g. small-scale phase-cryptic layering). Besides compaction to filter-press liquids, other methods for intercumulus-liquid ejection lie with its buoyancy, with development of polygonal grain boundaries and triple-point junctures (e.g. Sparks et al., 1985; Shirey, 1987), and with compositional convection (Tait & Kerr, 1987).

Filter-pressed liquids may also account for the fine-grained, generally evolved (<Fo72) mosaic-textured gabbros-some of which are composites with coarse-grained gabbros that accordingly define grain-size layering (Fig. 2g,h). These composites have only little compositional distinction between portions; namely, there is slightly higher An and lower Or in fine-grained plagioclase (Fig. 8c). Perhaps the fine-grained xenoliths represent late-stage liquids that were mobilized into cooler zones of a reservoir, where supercooling and high nucleation rates created fine textures at the margins of such ‘dikes’ (Brandeis & Jaupert, 1987). Whereas these fine-grained gabbros can represent undercooling (e.g. small grains; many nucleation sites), they are nonetheless cumulates (e.g. positive Eu anomalies, A2–2; Fig. 7a), or, at least, concentrations of crystals separated from liquids. In particular, liquid adjacent to the numerous plagioclase crystals at a ‘dike’ margin, and accordingly depleted in An compared with original liquid composition, may have moved to the ‘dike’ interior. The admixture of original gabbroic liquid (dike interior) and liquid filter-pressed from margins would have lower than original An content and, upon crystallization, account for the coarse-grained, lower An portions of composite xenoliths.

The grain-size layering that attends some modal and cryptic layering-namely, olivine-rich (coarse)–plagioclase-rich (fine) gabbros A9 and A22 (Fig. 2f; Table 1b)-is difficult to explain. It may be related to localized concentration gradients in the liquids. That is, any grain-size differences established across only small distances (e.g. 1 mm to a few centimeters) owing to, perhaps, differential cooling, may lead to modal and cryptic layering (and enhanced grain-size differences) because large crystals, because of lower surface energies, grow larger and extract components at the expense of small ones that dissolve and relinquish components (McBirney, 1993, p. 184).

Some layering may have been superimposed on cumulus mush by physical movements among the accumulations of crystals. Localized dynamic disturbances could be related to compaction, sloughing, convective processes, or hydraulic pressures (e.g. Irvine, 1980; Garcia et al., 1992). The small-amplitude folds and the small-scale faults in some xenoliths (Fig. 2e) may be the products of stresses induced by such physical processes. Movements within solidification zones could have mechanically segregated plagioclase to concentrate as ‘mono’-mineralogic mush-perhaps as a result of the lower density of plagioclase relative to associated olivine and clinopyroxene-and therefore account for modal layering (Fig. 2c) and for the ‘mechanical’ formation of troctolite and plagioclase-rich gabbro. The pyroxene and plagioclase compositions are similar enough across interfaces in modal- and phase-layered gabbroic xenoliths to suggest that if these layerings have mechanical origins, the mineral segregations and concentrations were confined to local zones (e.g. <100 m3) within cumulate piles.

Finally, mineral compositions in layered xenoliths do not provide any evidence for within-xenolith layering caused by introduction and crystallization of replenishment magmas. That is, small-scale modal, phase, and grain-size layerings cannot be attributed to sudden compositional changes in reservoir magmas because there are no wide compositional gaps among minerals in coexisting layers (e.g. Table 1b).

Adcumulus textures

Near-solidus processes in the crystallization environments formed polygonal, mutual-interference grain boundaries (Fig. 3c-f) during cooling and annealing (granoblastic) and/or during crystallization of interstitial liquids. In the context of cumulates, polygonal grain boundaries depict adcumulus growth (e.g. Hunter, 1987). Interstitial liquids producing adcumulus textures would have differed according to their cumulus environments. For example, liquids trapped among gravity-settled olivine would have had essentially ‘parental’, high-MgO compositions (original magma at reservoir bottom) consistent with crystallizing high-mg-number phases, but not plagioclase. In contrast, liquids interstitial during reservoir-margin solidification would have been compositionally influenced by local fractional crystallization processes occurring in situ, within solidification zones.

Cumulate-pile temperatures and phase equilibration

Phase equilibration at subsolidus temperatures is not surprising for cumulates. Equilibration temperatures for 21 orthopyroxene-clinopyroxene pairs are ∼953–1087°C (Tables 2 and 3) when calculated according to Wells, (1977); those from cone C span the entire range. Pairs of coexisting Ti-magnetite and ilmenite in gabbroic xenoliths yield lower equilibration temperatures. Using Fe–Ti oxide compositions mainly with <5 wt % Cr2O3, but two with 12 and 15 wt %, we calculate 704–890°C (Table 5). This range, however, subdivides to roughly correlate with sampling sites. Namely, cone-A and cone-B xenoliths have 704–823°C, and cone-C xenoliths have a higher range, 801–890°C.

Table 5

Cr-magnetite, Ti-magnetite, and ilmenite compositions (wt %) in representative gabbroic xenoliths of Mauna Kea

 C7 A2 A14 A21 A22 (ol) A22 (pl) A7 A7 A13 A13 C22 C22 C3 C3 A3 A3 
TiO2 7.4 6.2 9.6 10 7.5 14.9 11.7 52.3 10.6 50.1 14.8 48.8 13.3 49.7 7.7 49.5 
Al2O3 9.5 12.3 8.6 6.9 8.6 5.9 5.5 0.47 6.3 0.57 4.7 0.5 7.2 0.63 4.1 0.38 
Cr2O3 23.3 26.3 20.7 14.3 21.1 9.7 12 0.92 15.5 0.7 0.68 0.04 4.9 0.31 3.8 0.13 
FeO 49.2 43.8 51 60 54.5 59.4 60.6 37.5 58.1 38.6 71.1 42.3 64.8 38.4 77.8 43.8 
MnO 0.33 0.35 0.31 0.34 0.25 0.3 0.6 0.36 0.5 0.28 0.37 0.41 0.33 0.34 0.38 0.37 
MgO 7.3 6.5 7.4 6.2 5.6 5.5 7.4 7.9 5.8 4.7 7.5 4.7 8.3 2.9 6.4 
Sum 97.03 95.45 97.61 97.74 97.55 95.7 97.8 99.45 96.8 99.25 96.35 99.55 95.23 97.68 96.68 100.58 
Recalc. 28.2 28.1 30.1 32.1 31 36.8 31.4 32.5 32.7 28.7 38 30.1 36.7 29.6 34.6 32.7 
FeO                 
Recalc. 23.4 17.4 23.3 31 26.1 25.2 32.5 5.5 28.2 11 36.8 13.6 31.3 9.8 48 12.3 
Fe2O3 
Total 99.43 97.15 100.01 100.84 100.15 98.3 101.1 99.95 99.6 100.35 100.05 100.95 98.43 98.68 101.48 101.78 
mg-no. 31.6 58.9 30.5 25.6 24.3 21.1 29.6  24  18.1  18.6 18.6 13  
cr-no. 62.2 29.2 61.8 58.2 62.2 52.4 59.4  62.3  8.8  31.3  38  
usp       18.3  21.9  34.3  33.2  15.6  
hem         14.7  16.9  13  14.5 
Equil. T (°C)      704  814  890  848  784  
 C7 A2 A14 A21 A22 (ol) A22 (pl) A7 A7 A13 A13 C22 C22 C3 C3 A3 A3 
TiO2 7.4 6.2 9.6 10 7.5 14.9 11.7 52.3 10.6 50.1 14.8 48.8 13.3 49.7 7.7 49.5 
Al2O3 9.5 12.3 8.6 6.9 8.6 5.9 5.5 0.47 6.3 0.57 4.7 0.5 7.2 0.63 4.1 0.38 
Cr2O3 23.3 26.3 20.7 14.3 21.1 9.7 12 0.92 15.5 0.7 0.68 0.04 4.9 0.31 3.8 0.13 
FeO 49.2 43.8 51 60 54.5 59.4 60.6 37.5 58.1 38.6 71.1 42.3 64.8 38.4 77.8 43.8 
MnO 0.33 0.35 0.31 0.34 0.25 0.3 0.6 0.36 0.5 0.28 0.37 0.41 0.33 0.34 0.38 0.37 
MgO 7.3 6.5 7.4 6.2 5.6 5.5 7.4 7.9 5.8 4.7 7.5 4.7 8.3 2.9 6.4 
Sum 97.03 95.45 97.61 97.74 97.55 95.7 97.8 99.45 96.8 99.25 96.35 99.55 95.23 97.68 96.68 100.58 
Recalc. 28.2 28.1 30.1 32.1 31 36.8 31.4 32.5 32.7 28.7 38 30.1 36.7 29.6 34.6 32.7 
FeO                 
Recalc. 23.4 17.4 23.3 31 26.1 25.2 32.5 5.5 28.2 11 36.8 13.6 31.3 9.8 48 12.3 
Fe2O3 
Total 99.43 97.15 100.01 100.84 100.15 98.3 101.1 99.95 99.6 100.35 100.05 100.95 98.43 98.68 101.48 101.78 
mg-no. 31.6 58.9 30.5 25.6 24.3 21.1 29.6  24  18.1  18.6 18.6 13  
cr-no. 62.2 29.2 61.8 58.2 62.2 52.4 59.4  62.3  8.8  31.3  38  
usp       18.3  21.9  34.3  33.2  15.6  
hem         14.7  16.9  13  14.5 
Equil. T (°C)      704  814  890  848  784  
 A9 (ol) A9 (ol) A9 (pl) C6 C6 A15 A15 C5 C5 C4 C4 A2-2 A2-2 B4 B4 C19 (c) C19 (c) C19 (f) C19 (f) 
TiO2 8.7 48.3 46.4 14 49.3 13.4 50.8 14.5 51.1 11.4 47.5 8.7 49.6 12.2 51.1 13.8 48.4 14.1 49.3 
Al2O3 4.3 0.21 0.36 4.8 0.46 4.9 0.47 5.4 0.28 3.5 0.36 3.2 0.3 4.7 0.46 3.5 0.21 3.5 0.36 
Cr2O3 4.4 0.33 0.25 2.9 0.15 1.8 0.05 1.7 0.22 0.31 0.01 1.3 0.12 1.3 0.08 0.46 0.01 0.67 0.01 
FeO 73.3 42.1 42.9 71 43 70.7 38.1 68.9 39.7 78.2 43.9 78.8 42.7 72.7 38.7 75.4 44.8 74.5 41.9 
MnO 0.25 0.31 0.34 0.36 0.36 0.36 0.38 0.32 0.41 0.3 0.28 0.4 0.43 0.28 0.28 0.31 0.38 0.34 0.3 
MgO 3.3 7.7 6.6 4.9 6.9 4.1 7.4 4.3 6.9 3.4 6.4 2.6 5.9 3.7 6.9 2.4 4.4 3.2 6.5 
Sum 94.25 98.95 96.85 97.96 100.17 95.26 97.2 95.12 98.61 97.11 98.45 95 99.05 94.88 97.52 95.87 98.2 96.31 98.37 
Recalc. 34.2 29.4 29.6 37.6 31.7 37.4 32.1 38.1 33.2 37.2 31 35.3 33.7 36.9 33.4 40.4 35.3 39.6 32.4 
FeO                    
Recalc. 43.5 14.1 14.8 37.2 12.6 37 6.7 34.2 7.2 45.5 14.3 48.4 10.1 39.8 5.9 38.9 10.6 38.8 10.5 
Fe2O3                    
Total 98.65 100.35 98.35 101.76 101.47 98.96 97.9 98.52 99.31 101.61 99.85 99.9 100.15 98.88 98.12 99.77 99.3 100.21 99.37 
mg-no. 14.7   18.9  16.4  16.8  14  11.6  15.2  9.6  12.6  
cr-no. 40.7   28.8  19.8  17.4  5.6  21.4  15.7  8.1  11.4  
usp 18.9   25.1  32.1  35.9  25.3  19.2  28.9  36  35  
hem  17.8 18.3  15.2  8.5  8.9  17.2  11.9  7.4  11.9  12.7 
Equil. T (°C) 823   833  779  801  849  778  746  848  853  
 A9 (ol) A9 (ol) A9 (pl) C6 C6 A15 A15 C5 C5 C4 C4 A2-2 A2-2 B4 B4 C19 (c) C19 (c) C19 (f) C19 (f) 
TiO2 8.7 48.3 46.4 14 49.3 13.4 50.8 14.5 51.1 11.4 47.5 8.7 49.6 12.2 51.1 13.8 48.4 14.1 49.3 
Al2O3 4.3 0.21 0.36 4.8 0.46 4.9 0.47 5.4 0.28 3.5 0.36 3.2 0.3 4.7 0.46 3.5 0.21 3.5 0.36 
Cr2O3 4.4 0.33 0.25 2.9 0.15 1.8 0.05 1.7 0.22 0.31 0.01 1.3 0.12 1.3 0.08 0.46 0.01 0.67 0.01 
FeO 73.3 42.1 42.9 71 43 70.7 38.1 68.9 39.7 78.2 43.9 78.8 42.7 72.7 38.7 75.4 44.8 74.5 41.9 
MnO 0.25 0.31 0.34 0.36 0.36 0.36 0.38 0.32 0.41 0.3 0.28 0.4 0.43 0.28 0.28 0.31 0.38 0.34 0.3 
MgO 3.3 7.7 6.6 4.9 6.9 4.1 7.4 4.3 6.9 3.4 6.4 2.6 5.9 3.7 6.9 2.4 4.4 3.2 6.5 
Sum 94.25 98.95 96.85 97.96 100.17 95.26 97.2 95.12 98.61 97.11 98.45 95 99.05 94.88 97.52 95.87 98.2 96.31 98.37 
Recalc. 34.2 29.4 29.6 37.6 31.7 37.4 32.1 38.1 33.2 37.2 31 35.3 33.7 36.9 33.4 40.4 35.3 39.6 32.4 
FeO                    
Recalc. 43.5 14.1 14.8 37.2 12.6 37 6.7 34.2 7.2 45.5 14.3 48.4 10.1 39.8 5.9 38.9 10.6 38.8 10.5 
Fe2O3                    
Total 98.65 100.35 98.35 101.76 101.47 98.96 97.9 98.52 99.31 101.61 99.85 99.9 100.15 98.88 98.12 99.77 99.3 100.21 99.37 
mg-no. 14.7   18.9  16.4  16.8  14  11.6  15.2  9.6  12.6  
cr-no. 40.7   28.8  19.8  17.4  5.6  21.4  15.7  8.1  11.4  
usp 18.9   25.1  32.1  35.9  25.3  19.2  28.9  36  35  
hem  17.8 18.3  15.2  8.5  8.9  17.2  11.9  7.4  11.9  12.7 
Equil. T (°C) 823   833  779  801  849  778  746  848  853  

Equilibration temperatures based on titaniferous magnetite and ilmenite coexisting in the same samples; usp, ulvöspinel molecule; hem, hematite molecule.

Table 6

Whole-rock compositions (wt %) of ultramafic and gabbroic xenoliths from the southern flank of Mauna Kea volcano, Hawaii (in order of decreasing mg-number)

 dun dun dun gab nor ol cpx′nt gab gab nor ol cpx′nt gab nor gab nor gab nor gab nor gab nor gab gab gab 
 A6 A16 A19 A2 C9 C7 C17 C21 A14 A12 A13 A21 A7 C8 C1(c) C16 
SiO2 40.04 39.57 39.94 45.82 46.07 40.92 42.37 48.51 45.62 49.36 49.9 47.53 47.9 44.26 48.9 49.4 
TiO2 0.05 0.06 0.07 0.29 0.42 0.27 0.33 0.83 0.51 0.62 0.78 0.71 0.62 0.45 0.5 0.76 
Al2O3 0.64 0.53 0.98 4.8 3.93 3.84 4.11 2.62 4.14 7.7 9.44 7.58 9.37 8.89 16.74 16.8 
FeO 13.48 13.59 11.99 11.46 10.47 15.23 14.76 9.91 11.95 8.27 7.17 8.96 8.58 12.38 6.41 16.41 
MnO 0.26 0.26 0.23 0.21 0.2 0.26 0.25 0.18 0.22 0.17 0.14 0.17 0.16 0.2 0.1 0.1 
MgO 45.46 44.26 46.03 27.61 26.28 33.92 32.68 21.94 26.39 18.12 15.52 18.69 17.6 22.86 9.87 9.83 
CaO 0.3 0.36 0.31 8.98 11.34 4.28 4.37 15.27 10.3 14.29 15.12 13.88 13.58 8.36 14.37 14.05 
Na20.01 0.01 0.04 0.35 0.12 0.19 0.47 0.2 0.29 0.75 1.13 0.84 0.95 1.18 1.64 2.48 
K20.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.03 0.03 0.02 0.08 0.1 0.05 
P2O5 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.03 0.01 0.03 0.03 0.06 0.03 0.01 0.01 0.03 
Total 100.26 98.66 99.62 99.54 98.85 98.93 99.36 99.5 99.44 99.33 99.26 98.45 98.81 98.67 98.64 99.64 
LOI          0.15   0.2  0.69 0.06 
mg-no. 87 86.6 88.4 88.4 83.3 81.5 81.4 81.4 81.4 81.4 81.3 80.5 80.3 78.5 76.2 76 
Rb 
Sr 77 56 85 79 27 58 155 231 180 212 280 519 528 
Zr 24 12 21 18 16 11 19 12 14 
12 10 11 11  10 
Ni 1990 1870 2300 732 725 1200 1378 459 782 415 318 448 483 675 192 208 
Zn 102 125 101 104 72 124 118 80 88 52 38 67 58 103  42 
Cu 18 19  27 46 91 60 113 230 10 16 27 135  10 
Sc    27.1           32.8  
Hf    0.1           0.6  
La    0.5           1.4  
Ce    1.5           3.5  
Nd                 
Sm    0.57           1.31  
Eu    0.3           0.62  
Tb    0.16           0.2  
Yb    0.7           0.58  
Lu    0.11           0.08  
 dun dun dun gab nor ol cpx′nt gab gab nor ol cpx′nt gab nor gab nor gab nor gab nor gab nor gab gab gab 
 A6 A16 A19 A2 C9 C7 C17 C21 A14 A12 A13 A21 A7 C8 C1(c) C16 
SiO2 40.04 39.57 39.94 45.82 46.07 40.92 42.37 48.51 45.62 49.36 49.9 47.53 47.9 44.26 48.9 49.4 
TiO2 0.05 0.06 0.07 0.29 0.42 0.27 0.33 0.83 0.51 0.62 0.78 0.71 0.62 0.45 0.5 0.76 
Al2O3 0.64 0.53 0.98 4.8 3.93 3.84 4.11 2.62 4.14 7.7 9.44 7.58 9.37 8.89 16.74 16.8 
FeO 13.48 13.59 11.99 11.46 10.47 15.23 14.76 9.91 11.95 8.27 7.17 8.96 8.58 12.38 6.41 16.41 
MnO 0.26 0.26 0.23 0.21 0.2 0.26 0.25 0.18 0.22 0.17 0.14 0.17 0.16 0.2 0.1 0.1 
MgO 45.46 44.26 46.03 27.61 26.28 33.92 32.68 21.94 26.39 18.12 15.52 18.69 17.6 22.86 9.87 9.83 
CaO 0.3 0.36 0.31 8.98 11.34 4.28 4.37 15.27 10.3 14.29 15.12 13.88 13.58 8.36 14.37 14.05 
Na20.01 0.01 0.04 0.35 0.12 0.19 0.47 0.2 0.29 0.75 1.13 0.84 0.95 1.18 1.64 2.48 
K20.01 0.01 0.02 0.01 0.01 0.01 0.01 0.01 0.01 0.02 0.03 0.03 0.02 0.08 0.1 0.05 
P2O5 0.01 0.01 0.01 0.01 0.01 0.01 0.01 0.03 0.01 0.03 0.03 0.06 0.03 0.01 0.01 0.03 
Total 100.26 98.66 99.62 99.54 98.85 98.93 99.36 99.5 99.44 99.33 99.26 98.45 98.81 98.67 98.64 99.64 
LOI          0.15   0.2  0.69 0.06 
mg-no. 87 86.6 88.4 88.4 83.3 81.5 81.4 81.4 81.4 81.4 81.3 80.5 80.3 78.5 76.2 76 
Rb 
Sr 77 56 85 79 27 58 155 231 180 212 280 519 528 
Zr 24 12 21 18 16 11 19 12 14 
12 10 11 11  10 
Ni 1990 1870 2300 732 725 1200 1378 459 782 415 318 448 483 675 192 208 
Zn 102 125 101 104 72 124 118 80 88 52 38 67 58 103  42 
Cu 18 19  27 46 91 60 113 230 10 16 27 135  10 
Sc    27.1           32.8  
Hf    0.1           0.6  
La    0.5           1.4  
Ce    1.5           3.5  
Nd                 
Sm    0.57           1.31  
Eu    0.3           0.62  
Tb    0.16           0.2  
Yb    0.7           0.58  
Lu    0.11           0.08  
 gab gab nor gab nor gab nor dun troc fine gabn fine gabn gab troc fine gabn fine gabn gab fine gab 
 C15 A3-1 A5 A3 A18 C6 C2 C5 A15 C49 C4 A2-2 C22 C2 (f) 
SiO2 40.52 48.78 47.94 46.6 35.22 46.01 47.7 46.76 46.55 40.45 45.74 46.95 40.22 40.04 
TiO2 0.56 0.96 3.47 1.9 1.14 0.93 1.56 3.75 1.42 5.95 2.29 2.09 3.79 4.91 
Al2O3 4.14 9.9 9.8 8.54 0.94 19.81 13.38 14.13 17.52 16.44 13.34 14.11 7.87 12.65 
FeO 18.93 10.64 10.38 11.85 25.44 9.56 9.73 9.02 13.61 12.81 11.3 18.14 16.15 
MnO 0.3 0.19 0.17 0.19 0.58 0.11 0.13 0.12 0.11 0.12 0.14 0.17 0.18 0.15 
MgO 30.21 16.57 14.79 15.66 32.7 10.97 10.29 9.42 9.15 11.93 10.89 9.38 14.46 10.95 
CaO 4.31 10.9 11.36 12.87 0.62 13.65 14.08 13.75 7.66 12.51 12.03 14.15 13.61 
Na20.47 1.24 1.54 1.27 0.01 2.73 2.21 1.82 1.91 2.35 1.95 2.73 0.64 0.63 
K20.03 0.04 0.05 0.04 0.01 0.09 0.04 0.05 0.09 0.11 0.05 0.11 0.02 0.07 
P2O5 0.01 0.03 0.02 0.03 0.31 0.02 0.02 0.03 0.03 0.02 0.02 0.04 0.03 0.01 
Total 99.48 99.25 99.52 98.95 96.97 99.23 98.71 99.16 99.55 98.64 99.74 98.91 99.5 99.17 
LOI        0.2 0.33      
mg-no. 76 75.5 73.8 72.4 71.8 69.5 67.7 67.5 66.8 63.5 62.8 62.2 61.2 57.3 
Rb 
Sr 98 252 272 220 10 782 385 405 604 565 369 443 222 415 
Zr 19 66 37 14 12 14 15 59 48 
11 13 10 12 11 16 
Ni 868 448 352 371 930 226 153 179 171 323 191 158 211 181 
Zn 137 73 62 72 212 66 50 62 61  63 83 117  
Cu 85 228 102 44 22 63 17 332 63   40  
Sc     3.1      32.7 31.8  41.5 
Hf           0.4 0.75  2.5 
La     0.6      0.9 2.8  3.5 
Ce            4.9  8.1 
Nd     0.19       4.9  10.9 
Sm     0.19      0.67 1.48  3.45 
Eu     0.06      0.48 1.04  1.29 
Tb     0.01       0.33  0.51 
Yb            1.06  0.95 
Lu     0.01      0.12   0.13 
 gab gab nor gab nor gab nor dun troc fine gabn fine gabn gab troc fine gabn fine gabn gab fine gab 
 C15 A3-1 A5 A3 A18 C6 C2 C5 A15 C49 C4 A2-2 C22 C2 (f) 
SiO2 40.52 48.78 47.94 46.6 35.22 46.01 47.7 46.76 46.55 40.45 45.74 46.95 40.22 40.04 
TiO2 0.56 0.96 3.47 1.9 1.14 0.93 1.56 3.75 1.42 5.95 2.29 2.09 3.79 4.91 
Al2O3 4.14 9.9 9.8 8.54 0.94 19.81 13.38 14.13 17.52 16.44 13.34 14.11 7.87 12.65 
FeO 18.93 10.64 10.38 11.85 25.44 9.56 9.73 9.02 13.61 12.81 11.3 18.14 16.15 
MnO 0.3 0.19 0.17 0.19 0.58 0.11 0.13 0.12 0.11 0.12 0.14 0.17 0.18 0.15 
MgO 30.21 16.57 14.79 15.66 32.7 10.97 10.29 9.42 9.15 11.93 10.89 9.38 14.46 10.95 
CaO 4.31 10.9 11.36 12.87 0.62 13.65 14.08 13.75 7.66 12.51 12.03 14.15 13.61 
Na20.47 1.24 1.54 1.27 0.01 2.73 2.21 1.82 1.91 2.35 1.95 2.73 0.64 0.63 
K20.03 0.04 0.05 0.04 0.01 0.09 0.04 0.05 0.09 0.11 0.05 0.11 0.02 0.07 
P2O5 0.01 0.03 0.02 0.03 0.31 0.02 0.02 0.03 0.03 0.02 0.02 0.04 0.03 0.01 
Total 99.48 99.25 99.52 98.95 96.97 99.23 98.71 99.16 99.55 98.64 99.74 98.91 99.5 99.17 
LOI        0.2 0.33      
mg-no. 76 75.5 73.8 72.4 71.8 69.5 67.7 67.5 66.8 63.5 62.8 62.2 61.2 57.3 
Rb 
Sr 98 252 272 220 10 782 385 405 604 565 369 443 222 415 
Zr 19 66 37 14 12 14 15 59 48 
11 13 10 12 11 16 
Ni 868 448 352 371 930 226 153 179 171 323 191 158 211 181 
Zn 137 73 62 72 212 66 50 62 61  63 83 117  
Cu 85 228 102 44 22 63 17 332 63   40  
Sc     3.1      32.7 31.8  41.5 
Hf           0.4 0.75  2.5 
La     0.6      0.9 2.8  3.5 
Ce            4.9  8.1 
Nd     0.19       4.9  10.9 
Sm     0.19      0.67 1.48  3.45 
Eu     0.06      0.48 1.04  1.29 
Tb     0.01       0.33  0.51 
Yb            1.06  0.95 
Lu     0.01      0.12   0.13 

Major-element analyses are reported volatile-free, after ignition at 1000°C. LOI, loss on ignition, or devolatilization, at 1000°C; if no LOI data are listed, sample gained weight upon ignition (i.e. Fe oxidation offsets H2O loss).mg-number=[100 × Mg/(Mg + Fe2+), where Fe2+ assumes that 90% of all iron is ferrous. dun, dunite; ol cpx'nt, olivine clinopyroxenite; gab, gabbro; gab nor; gabbronorite; fine gabn, fine-grained gabbronorite; fine gab, fine-grained gabbro; troc, troctolite; (c) and (f) refer to coarse- and fine-grained portions of a grain-size composite; sample A3-1 is a split of A3.

Table 7

Isotope compositions for selected xenoliths and clinopyroxene and plagioclase grains from xenoliths, Mauna Kea volcanos

  87Sr/86Sr 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb δ18
A2 whole-rock 0.703556 ± 09     
C1c whole-rock 0.703614 ± 22     
B3 whole-rock  18.391 ± 4 15.464 ± 4 37.933 ± 10  
A13 clinopyroxene  18.444 ± 9 15.630 ± 11 38.276 ± 27 4.05 
A13 plagioclase     5.09 
C18 clinopyroxene  18.417 ± 6 15.479 ± 5 37.976 ± 14 5.36 
C18 plagioclase     5.28 
C22 clinopyroxene  18.395 ± 5 15.493 ± 5 37.995 ± 13 4.98 
C22 plagioclase     5.62 
B2 clinopyroxene     4.38 
B2 plagioclase     5.23 
A17 clinopyroxene     5.07 
  87Sr/86Sr 206Pb/204Pb 207Pb/204Pb 208Pb/204Pb δ18
A2 whole-rock 0.703556 ± 09     
C1c whole-rock 0.703614 ± 22     
B3 whole-rock  18.391 ± 4 15.464 ± 4 37.933 ± 10  
A13 clinopyroxene  18.444 ± 9 15.630 ± 11 38.276 ± 27 4.05 
A13 plagioclase     5.09 
C18 clinopyroxene  18.417 ± 6 15.479 ± 5 37.976 ± 14 5.36 
C18 plagioclase     5.28 
C22 clinopyroxene  18.395 ± 5 15.493 ± 5 37.995 ± 13 4.98 
C22 plagioclase     5.62 
B2 clinopyroxene     4.38 
B2 plagioclase     5.23 
A17 clinopyroxene     5.07 

c refers to coarse-grained portion of a grain-size composite xenolith.

The CaO in xenolith olivines offers another means of estimating temperatures. CaO is typically <0.2 wt %, lower than the >0.2 wt % present in olivine phenocrysts of Hawaiian lavas (Fig. 4). The interpretation (Bohrson & Clague, 1988) is that Ca solubility in olivine coexisting with clinopyroxene and orthopyroxene decreases with temperature, and that subsolidus equilibration among these three phases at temperatures of ∼800–900°C-the same as that indicated by Fe–Ti oxide compositions-depleted olivine of its original CaO concentrations. Whereas only some Mauna Kea xenoliths contain two pyroxenes, CaO olivine in xenoliths without orthopyroxene is also low because the temperature effect on Ca solubility in olivine still holds when only olivine and clinopyroxene equilibrate (Takahashi, 1980).

We explore geothermometry related to Ca solubility in olivine according to Kohler & Brey, (1990) for upper-mantle xenoliths, which the Mauna Kea porphyroclastic ultramafic xenoliths resemble modally and texturally. The first requirement is an estimate of pressure. In the fashion of the study by Bohrson & Clague, (1988) of Hualalai xenoliths, we constrain pressure for the porphyroclastic xenoliths to less than ∼5 kbar (because orthopyroxene precedes or attends olivine between ∼5 and 11 kbar). By assigning cumulate equilibration to ∼4.5 kbar, equilibration temperatures for the porphyroclastic xenoliths range from ∼1000°C for those with lowest CaO in olivine (∼0.06 wt %) to ∼1150°C for xenoliths with highest CaO in olivine (∼0.21 wt %). (Temperatures decrease only slightly with a decrease in estimated pressure.) This a reasonable temperature range in light of its overlap with the high end of the temperature range for pyroxene equilibration in the gabbroic xenoliths, and in view of our model, which places porphyroclastic ultramafic xenoliths below the gabbroic xenoliths and therefore hotter.

Conclusions about Hawaiian magmatism as drawn from xenoliths

Ultramafic and gabbroic xenoliths at cones A, B, and C on the southern flank of Mauna Kea mainly represent tholeiitic, and to a lesser extent alkalic, magmatism of one period of the volcano's multi-stage construction-Hamakua postshield. Interpreting that this xenolith location represents Hamakua tholeiitic, transitional, and alkalic magmas relies heavily on the amount of Al2O3 and REE in clinopyroxene, the presence or absence of orthopyroxene, the spinel cr-numbers in ultramafic xenoliths, and on Sr isotopic ratios.

Dunite formation at Mauna Kea-of both tholeiitic and alkalic parentages-originated at <15 km depth from magmas having ∼15 wt % MgO. Olivine in olivine clinopyroxenites and wehrlites associated with dunite indicates that the ‘original’ magmas evolved to ∼9–10 wt % MgO. These cumulates were probably produced by both gravity-settling and in situ crystallization at reservoir bottoms. Based on their textures, thermal and stress regimes allowed interstitial liquids to produce adcumulus textures and ‘metamorphic’ conditions to strain the cumulus olivines and recrystallize assemblages to porphyroclastic textures.

The liquids residual after formation of plagioclase-free ultramafic cumulates had ∼8–9 wt % MgO. Differentiation thereof produced liquids that evolved to FeO/MgO 2.5–3.0 (e.g. <Fo72 in fine-grained gabbroic xenoliths). These <9 wt % MgO magmas had viscosities that restricted crystal segregation and they therefore crystallized in solidification zones along reservoir margins-a history supported by framework textures and by the progressively evolved modal and mineral compositions manifested by the plagioclase-bearing xenoliths (i.e. modal and cryptic layering). We view this sequential formation of plagioclase-bearing wehrlite, olivine clinopyroxenite, gabbro, gabbronorite, and troctolite as growing inward from reservoir margins, and caused largely by differentiation (e.g. Fo83–61) accomplished by interstitial liquids that re-entered reservoir(s) from solidification zones. In detail, high-Fo olivine was in olivine-rich, plagioclase-poor zones closest to reservoir margins, and low-Fo olivine was in olivine-poor, plagioclase-rich zones farthest (most inward) from wallrock. The An–Or compositions of interstitial plagioclase of some reservoir products indicate that some pore liquids were ‘trapped’ amidst cumulus crystals to fractionate to relatively high normative-Or liquids.

The modes of xenoliths with olivine+clinopyroxene framework and plagioclase framework textures approximate original crystallization proportions within their representative solidification zones; some fine-grained and plagioclase-rich gabbros probably represent eutectic compositions and phase proportions as expressed in the Fo-Di-An-Ab system. Small-scale modal, phase, cryptic, and grain-size layerings were superimposed on the large-scale modal-cryptic layerings (e.g. Fo83–61) of solidification zones. They are due to processes such as liquid segregation, where liquids ‘pooled’ or sometimes acquired concentration gradients. Such small-scale modal and cryptic layerings manifest small compositional differences where, for example, the olivine-dominant portions are ∼1 mg-number unit richer. Fine-grained gabbroic xenoliths point to residual melts having been mobilized into fractures that developed in cooling, solidifying crystal mushes. Additionally, mechanical segregations of crystals as a result of convective, hydraulic, and gravitational disturbances on mushes probably created some modal layerings.

Another perspective that the xenoliths provide is the extent to which Hawaiian magmas compositionally evolve, which otherwise can only be assessed from lava compositions. Evaluated from xenolith phase compositions of ∼Fo89–61, ∼An85–50, and clinopyroxene mg-number ∼90–76, and from the compositions of Hamakua lavas (Frey et al., 1990, 1991; Yang et al., 1994), magmas undergoing crystallization in the subsurface are not necessarily manifested as Mauna Kea lavas. For example, the most evolved Hamakua lavas, such as Fe–Ti basalt KI-12 (Frey et al., 1991), are in equilibrium with ∼Fo72 and therefore too ‘primitive’ to represent the liquids that produced the highly evolved, generally fine-grained, gabbroic xenoliths (e.g. <Fo72–61; Tables 1b and 3). This holds even though an Fo60 microphenocryst was observed in KI-12 (Frey et al., 1991), because that olivine represents liquid with FeO/MgO >3 and is therefore in disequilibrium with its host-rock composition (FeO/MgO ∼2.3). Apparently, some liquids that crystallized fine-grained gabbroic xenoliths (<Fo72) are small-percent residual melts that did not erupt but instead only pervaded plutonic environments.

The extensive compositional range of plutonic crystallization—representing a continuum from MgO-rich to Fo–Di–An ‘eutectic’ magmas—was duplicated at Hualalai volcano, which also yields dunite–gabbro xenolith suites. There are no comparably evolved mineral compositions yet analyzed for Hualalai xenoliths, but it is established that there are gabbros with MgO 8–9 wt %, and with CaO and Al2O3 abundances and mineral modes (Jackson et al., 1981) like those we observe among Mauna Kea flank cone xenoliths (Table 6).

The preponderance of gabbroic xenoliths at Mauna Kea distinguishes the xenolith suite from that of Hualalai (D. A. Clague, personal communication, 1996), which is dominated by ultramafics. The inference, then, is that magma-production rates influenced fractionation in the reservoirs represented. That is, there were insufficient replenishments during Hamakua magmatism to buffer reservoir compositions and prevent copious amounts of magma from achieving FeO/MgO >2.

The admixture of tholeiitic and alkalic xenoliths on Mauna Kea's south flank suggests that eruptions (cones) sampled cumulate regimes that are different from, as well as perhaps identical to, those sampled by summit-cone eruptions. Namely, summit-cone olivine gabbros and opaque-oxide gabbros have, respectively, alkalic and Fe–Ti basalt affinities. As far as we can determine, the subsurface source for opaque-oxide gabbro xenoliths—namely, Fe–Ti-rich magma (Fig. 7b)—was not disrupted by magmatism along Mauna Kea's flank. Conversely, ‘alkalic’ gabbro C22, with clinopyroxene enriched in Wo, Al2O3, and TiO2, is a type not yet observed elsewhere on Mauna Kea. In reality, however, there are probably many more cumulate regimes than ever sampled by eruptions. One indication of variety, albeit somewhat subtle, comes from neighboring cones A and C both having sampled predominantly tholeiitic cumulates, but with Fe–Ti-oxide equilibration temperatures to suggest different milieux.

All told, Mauna Kea xenoliths studied demonstrate that the subsurfaces of Hawaiian shield volcanoes share igneous processes intrinsic to continental mafic-layered intrusions. These are cryptic, phase, modal, and grain-size layerings, gravity-settled and in situ crystallization, poikilitic and adcumulus textures, a variety of differentiated tholeiitic liquids, plus some transitional to alkalic, alkalic, and Fe–Ti rich. Collectively, then, they show that the Mauna Kea postshield magma system had compositions and solidification zones that developed extensively and effectively enough to facilitate this variety of processes.

Acknowledgements

This project was funded by NSF Grant EAR9104927, but we additionally want to thank the Geological Society of America for a student grant to P.G. for isotope analyses, Kreuger Enterprises for a grant to P.G. for oxygen isotope analyses, and the Oregon State University Radiation Center for providing REE analyses under their Department of Energy Reactor Sharing grant. Steve Goldberg enabled us to obtain Sr and Pb isotopic ratios at the University of North Carolina mass spectrometry laboratory. Finally, R.F. thanks former HVO staff members, particularly E. Wolfe and R. B. Moore, for providing the opportunity to begin this Mauna Kea study in 1984, G. R. Bauer for suggesting it, and D. A. Clague and J. M. Rhodes for insightful, critical manuscript reviews.

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