Gravity exploration
DATY 21 DATY 2 − DATY 1 Saad and Bishop, 1989
4.5 Rock densities
The primary controls on the density of subsurface mate- rials are mineral composition and void space, which are largely dependent on the lithology (rock type) and the chemical and physical effects of secondary pro- cesses including rock fracturing, solutioning, and chem- ical alteration of minerals. The density of minerals varies from 1,990 kg/m3 for sylvite, the potassium salt, to about 20,000 kg/m3 for gold, but the vast majority of commonly occurring minerals range from 2,500 to 3,500 kg/m3, although ore minerals of metals are in the 4,000–6,000 kg/m3range. Examples of mineral densities are given in Tables 4.2, 4.3 and 4.4, where several gen- eralities can be noted. For example, the density of min- erals tends to decrease with an increase in SiO2 content, and rock densities decrease with an increase in H2O con- tent. Furthermore, minerals formed under high-pressure conditions in metamorphic rocks tend to have higher den- sities, and non-metalliferous resources (e.g. coal, salts, clays) have lower than average densities, while metallifer- ous ores have densities exceeding those of common rock- forming minerals. In addition, densities are affected by lithostatic pressure and temperature, both of which are primarily functions of depth within the Earth.
TABLE 4.3 Examples of densities in kg/m3of igneous and metamorphic rock minerals.
Name Density
Augite 3,300
Biotite 3,360
Ca Al pyroxene 3,360
Chlorite 2,800
Diamond 3,520
Feldspars
Albite to anorthosite 2,620–2,760
Microcline 2,560
Orthoclase 2,570
Sanadine 2,560
Hornblende 3,080
Muscovite 2,560
High-grade metamorphic minerals
Sillimanite 3,250
Kyanite to
Garnet, etc. 4,300
Olivine
Fosterite to fayalite 3,210–4,390
Quartz 2,650
Serpentine 2,600
Talc 2,780
Adapted primarily from Johnson and Olhoeft (1984).
4.5.1 Lithology
For purposes of considering density, Earth materials are conveniently classified into crystalline rocks, sedimentary rocks, and unconsolidated sediments. Crystalline rocks include igneous rocks, both plutonic and volcanic, that originate from magma that has solidified, respectively, within the Earth and at the surface. In addition, they include metamorphic rocks, derived from both igneous and sedimentary rocks that have been altered deep in the crust by increased lithostatic pressure, tectonic stress, and enhanced temperatures. Unconsolidated sediments are made up of the fragments derived from erosion of pre- existing rocks that are commonly deposited in water or less commonly in air and by chemical precipitants. Sedimen- tary rocks consist of sediments that have been lithified by lithostatic pressure and chemical precipitants. The compo- nent minerals and the origin and nature of the void space in these three classes generally occur within specified ranges, facilitating consideration of their densities.
4.5 Rock densities 71
TABLE 4.4 Examples of densities in kg/m3of ore minerals and metals, and terrestrial waters.
Name Density
Ore minerals and metals
Barite 4,480
Cinnabar 8,187
Chalcopyrite 4,200
Chalcocite 5,793
Copper 8,934
Corundum 3,987
Galena 7,600
Gold 19,282
Halite 2,163
Hematite 5,275
Iron 7,875
Kaolinite 2,594
Lead 11,343
Limonite 4,880
Magnetite 5,200
Malachite 4,031
Pyrite 5,010 (4,950–5,030)
Pyrrhotite 4,610
Sphalerite 4,089
Uraninite 10,970
Water
Fresh (at 4◦C) 1,000
Ice 890–910
Brine 1,125
Sea water 1,030
Adapted primarily from Johnson and Olhoeft (1984).
Crystalline rocks
Unaltered plutonic igneous rocks that are formed deep within the crust or upper mantle characteristically have minimal void space, generally less than 1%, and values seldom exceed 3%. A minor portion of this volume is originally due to intergrain voids, but the vast majority of voids are caused by chemical and physical weathering within the upper few hundred meters of the surface and by fracturing and faulting within the brittle crust plus cooling cracks. Generally, voids occurring within fractures such as rock joints and in faults largely, but not completely, close up under the effect of lithostatic pressure. As a result, both densities and associated seismic velocities approach constant values with increasing pressure at values of the order of 600 MPa (1 Pa = 10−8kilobars [kb]) which is reached at depths within the Earth of the order of
3.5 3.0 2.5
50
0 100%
1 2 3 4 5 6 7
12
11 10 9
(kg/m
8
Density
FIGURE 4.2 Schematic illustration showing the dependence of the density of plutonic igneous rocks upon their mineralogical composition. The lower panel shows the proportional content of the major minerals of igneous rocks, whereas the upper panel gives the range of densities of the corresponding rock types. 1 – peridotite and pyroxenite; 2 – gabbro; 3 – diorite; 4 – granodiorite;
5 – granite; 6 – syenite; 7 – nepheline syenite; 8 – nepheline (σ= 2,600 kg/m3); 9 – potassium feldspar (σ=2,500 kg/m3); 10 – q uartz (σ=2,600 kg/m3); 11 – plagioclase (σ=2,600 to 4,800 kg/m3);
12 – iron-magnesium minerals (σ=3,100 to 3,500 kg/m3). Adapted fromP i c ket al.(1973).
15–20 km. The decrease in density in fault zones is not only a result of the increase in volume during fracturing, but also the result of physiochemical alteration of the rocks to lower-density clay minerals.
Given the minor void space in plutonic rocks, the den- sity of these rocks is primarily the result of the densities and constituent proportions of the minerals present. The lowest-density minerals (see Table 4.3) commonly present in plutonic rocks are quartz and orthoclase, which have densities around 2,600 kg/m3. Rocks with relatively high proportions of these minerals are called felsic (or acidic) and have the lowest densities among plutonic igneous rocks. In contrast, rocks containing significant propor- tions of plagioclase feldspars, biotite, and hornblende, and thus rich in calcium, magnesium, and iron, have greater densities. These are informally termed mafic (or basic) rocks. The variation in the density of plutonic rocks with varying mineralogical composition is generalized in Figure 4.2.
The crystalline rocks of the continental crust in gen- eral become more mafic with depth, and thus densities increase with depth over and above the effect of increas- ing lithostatic pressure. However, the continental crust is not layered in the simple view of an upper granitic layer underlain by a basaltic layer. Rather, all evidence points to a crust, as illustrated in Figure 4.3 that varies significantly
72 Density of Earth materials
00 100 200
10 20 30 40 50 60
0 10 20 30 40 50 60 Distance (km)
Depth (km)
Strike – slip fault
6A
6A
6A 6C
6D 6D
6D 6D 6D
6D 6D
6D 6B
5B 5B
5B
5B
5C 5B
5B 4
4 4 4
3A 4 3A
3A 3A
3A
3D
2E 2E 2E
2E
2E 2E
2B 2B 2B
2C 2D 2D
2D 2B
2B
3D 3D
5B 6B
6D
6D 6A
6A 6A 6A
6A 6A
FIGURE 4.3 Hypothetical cross-section of the crystalline crust and upper mantle. (1) Sedimentary rocks; (2) metamophosed supracrustal rocks; (3) plutonic rocks; (4) gneiss and migmatite; (5) mafic rocks; (6) ultramafic rocks. The Moho is the line separating mafic gneiss (5B) from the peridotite (6D) of the upper mantle. Adapted fromF o u n t a i nandC h r i s t e n s e n(1989).
in both composition and density over short lateral and ver- tical dimensions (FountainandChristensen, 1989).
In a general sense, increasing metamorphic grade is evi- dent with increasing depth, supporting the view of increas- ing density with depth. The densities of surficial rocks vary greatly because of the presence of materials ranging from unconsolidated sediments to mafic crystalline rocks.
However, the overall average of 2,670 kg/m3reflects their general granitic (felsic) nature (Hinze, 2003).
The increase in density with depth to values around 3,100 kg/m3 in the lower continental crust results in an average crustal density of roughly 2,830 kg/m3 (Chris- tensen andMooney, 1995). The lower crust has an average chemical composition equivalent to gabbro, but garnet becomes more abundant with depth. At the base of the crust, mafic garnet granulite is most abundant. Gran- ulite is a relatively coarse-grained, high-grade metamor- phic rock that is formed at the high temperatures of the lower crust. Despite the widespread lateral density vari- ations in the crust, it is useful to consider an average density profile through the continental crust, illustrated in Figure 4.4 as a starting point for investigating hetero- geneities. However, it is important to remember that this is an average profile when developing density profiles of specific regions of the crust. For example, velocity stud- ies of the Archean Pilbara Craton of northwest Australia show two layers with a transition zone of only a few verti- cal kilometers between them that defines a density contrast of the order of 50 kg/m3at depths of about 10 to 15 km (Drummond, 1982).
3.1 2.9 2.7 2.7 2.6
2.6
3.1 2.9
Base of crust
Density (kg/m3 × 103)
Depth (km)
0
10 15 20 25 30 35 40 5
FIGURE 4.4 Average density profile through the crust. Adapted fromC h r i s t e n s e nandM o o n e y(1995).
The oceanic crust is apparently more laterally homo- geneous in both composition and density than the con- tinental crust and consists primarily of mafic extrusive and intrusive rocks beneath a layer of sediments and sedimentary rocks. However, there are lateral differences associated with varying age of the crust, alteration (or metamorphic grade), tectonic province and heat flow
4.5 Rock densities 73 (e.g.ChristensenandSalisbury, 1975;Carlson
andHerrick, 1990). Carlson and Herrick’s analysis of laboratory and downhole logging of oceanic crystalline materials and seismic investigations of the oceanic crust shows that the upper crystalline layer (“layer 2”), which consists largely of lavas extruded on the sea floor, has a density ranging from 2,620 to 2,690 kg/m3, and the low- ermost layer of the oceanic crust (“layer 3”), consisting of dikes and mafic magma chambers, has an estimated average density of 2,920 to 2,970 kg/m3. They estimate the average density of the entire crystalline oceanic crust to be 2,860±30 kg/m3, very similar to Christensen and Mooney’s estimate of the density of the continental crust.
Thus, gravity anomalies at the continent/ocean interface are primarily due to changes in the thickness of the crust and sedimentary rock accumulations.
The ultramafic upper mantle has a density of approx- imately 3,300 (3,270–3,320) kg/m3, but varies with rock composition, temperature, pressure effects, and the degree of partial melting. Peridotite is the principal lithology of the upper mantle based on geophysical studies and direct observation mantle rocks exposed on the Eart’s surface.
These rocks are rich in olivine, but another important lithology of continental upper-mantle is eclogite contain- ing garnet. These latter rocks, which are proposed to originate from tectonic processes involving remnants of subducted crust into the mantle, may have densities of 200 kg/m3greater than peridotite. Upper-mantle continen- tal rocks generally have undergone periods of multiple melting leading to depletion in Ca, Fe, Al, and Ti. These depleted rocks are rich in olivine and Mg content and have a lower density. Clearly, the upper mantle, particularly in the continents, is subject to variations leading to lateral changes in density. For example,KabanandMooney (2001) describe a±3% variation (∼100 kg/m3) in the den- sity of the upper mantle of the southwestern USA, which they ascribe largely to compositional differences rather than temperature, on the basis of seismic wave velocities.
In analyzing the gravity anomalies associated with the Kenya rift in East Africa,Ravat et al.(1999) find on the basis of seismic velocities and gravity modeling that the mantle plume beneath the rift at a depth of about 100 km has a density contrast of approximately−30 kg/m3 with the surrounding mantle.
In addition, the density of the upper mantle incorpo- rated into subducted slabs increases as the slab undergoes metamorphic reactions with increasing depth of subduc- tion (e.g. Grow and Bowin, 1975; Tessara et al., 2006). A thermal–petrologic–seismological model of sub- duction zones (Hackeret al., 2003, 2004) is available which includes provision for calculating the density of sub-
duction zone rocks at high pressures and temperature. The asthenosphere, which is recognized as the low-velocity zone at the base of the lithosphere in the upper mantle, is believed to have a small (<50 kg/m3) negative density con- trast with the surrounding mantle, and ´Swieczaket al.
(2009) suggest the density of the asthenosphere beneath Europe is in the range 3,100–3,340 kg/m3.
The density contrast across the Moho (crust/mantle interface) has been assigned a range of values from of the order of 200 to 450 kg/m3. Based on a density of the low- ermost crust of 3,100 ( Figure 4.4) and a mean density of the upper mantle of 3,300 kg/m3, the density contrast at the Moho is 200 kg/m3. But where the crust/mantle inter- face has a relief that reaches several kilometers or more, the density contrast is greater. In fact, Simpson et al.
(1986) find that a density difference across the Moho in the conterminous USA of 350 kg/m3 provides Moho depths consistent with seismic refraction measurements. How- ever,Chapin(1996) arrives at a value of 450 kg/m3 for the crust/mantle density contrast across the South Amer- ican continent, based on analysis of observed continental gravity values, andKabanandMooney(2001), on the basis of seismic velocity data, find a density contrast of 420 kg/m3 across this interface in the southwestern USA Clearly the crust/mantle interface density contrast is a non- linear function with relief as a result of the changing den- sity with depth and is likely to vary depending on the geologic terrane.
Igneous rocks also form close to or on the Earth’s sur- face producing volcanic or extrusive igneous rocks. The rapid crystallization of igneous melts in this environment leads to a fine-grained texture in contrast to the coarse texture characteristic of plutonic rocks. The fine-grained texture tends to lower the density despite the rock having a similar composition, but the effect generally is less than 10%. The occurrence of non-crystalline glasses in these rocks such as obsidian and vitrophyre, for example, leads to even lower densities. Commonly the density of volcanic rocks is lowered by the presence of voids resulting from gas cavities frozen into the rock during its rapid crystalliza- tion. This situation is particularly prevalent in the upper portion of volcanic flows where gas rising through the flow accumulates. In extreme situations, the densities of these rocks, such as pumice and scoria, will be lowered to less than that of water by the cavities filled with entrapped gas.
Volcanic ash deposits typically are highly porous, and thus are significantly lower in density than their plutonic or lava flow equivalents, but pore space is readily decreased when ash is welded during the depositional process by its internal heat or compacted by burial. For example, felsic
74 Density of Earth materials
Rhyotite lavas & tuff of Calico Hills Prow Pass Member *
Static Water Level Pah Canyon & Yucca Mountain Members, undivided +
Topopah Spring Member +
Bullfrog Member *
Depth (m)
Density (kg/m3 ×103)
Tram Member *
Dacite lava breccia
260 k g/m
3 per kilomet
er
Lithic Ridge Tuff
Other tuff deposits
Member of the Paintbrush Tuff Member of the Crater Flat Tuff Sea Level
+
*
1.5 2.0 2.5
500
1,000
1,500
Tiva Canyon Member +
FIGURE 4.5 Density versus depth plot for Tertiary tuff at Yucca Mountain, Nevada. Solid rectangles indicate drillhole gravity measurements, depth range, and error range; dashed lines are gamma–gamma density measurements (see Section 4.6.3); circles are density measurements on unsaturated core from saturated zone at Yucca Mountain. Solid line is least-squares fit to the data. Adapted fromS n y d e randC a r r(1984).
Tertiary tuffs of Southern Nevada vary in densities from about of 1,700 to 2,500 kg/m3, with respective porosities from a few percent to 50%. A density profile from a drill hole into these rocks, obtained from a variety of measure- ment methods, shows a generally linear increase in density as illustrated in Figure 4.5. In this case the increase in den- sity is linear, but in other situations the density increases exponentially with the largest increase near the surface.
Blakelyet al.(1999) report that in the volcanic basins of the Basin and Range Province, volcanic rocks increase in density by 200 kg/m3from the surface to 1.2 km. The density of volcanic flows may be lowered by fracturing associated with movement of the crystallizing melt and by the effect of chemical alteration of the flows leading to less dense minerals.
The density of metamorphic rocks that make up the majority of continental crust depends primarily on the original mineral composition of the rocks, but is also strongly influenced by the degree and type of metamor- phism which largely reflects the temperature and pres- sure to which the rocks have been subjected. In general, densities increase with chemical compositions of rocks higher in iron, magnesium, and calcium. Numerous studies show the broad range of densities of exposed metamorphic rocks of Precambrian shields (e.g.Woollard, 1962;
Gibb, 1968;Smithson, 1971;Subrahmanyamand Verma, 1981;Korhonenet al., 1993). Typically, mean densities of metamorphic rocks occur in the range 2,700 to 2,800 kg/m3, but many metamorphic rocks derived from felsic igneous rocks and metasedimentary rocks, such as granite gneiss, have densities in the range of 2,600 to 2,700 kg/m3. In contrast, intermediate to mafic metavol- canic rocks have significantly higher densities.Gupta andGrant(1985) report a mean density for these rocks of the order of 2,850 kg/m3 in the Sudbury–Cobalt area of Canada, and Gibb (1968) has found similar values for these rocks in the Precambrian shield of Manitoba, Canada. The highest-density metamorphic rocks are those that have been metamorphosed to the eclogite grade in the lower crustal environment where garnet has replaced pla- gioclase. These rocks typically have densities in the range 3,000 to 3,300 kg/m3.
Sedimentary rocks and unconsolidated sediments The densities of unconsolidated sediments and their lithi- fied equivalents, sedimentary rocks, are primarily con- trolled by their void space. Their mineral components are rather limited, and those that are present do not vary greatly in density. The most common constituent is quartz derived from the chemical and physical breakdown of igneous rocks and their metamorphic equivalents. Clay minerals originating by chemical modification of feldspars and other silicate minerals which dominate in clays and shales are another important component. Sediments and sedimentary rocks formed by chemical precipitation are largely monomineralic, and thus are of constant density except where post-depositional processes primarily cause dissolution that produces secondary porosity. Limestone (calcite) and dolomite are the principal chemical sed- iments. However, salt deposits generally consisting of halite are important locally in producing density con- trasts within sedimentary basins. All of these mineral con- stituents, quartz, clays, and calcite, have densities in the range 2,600 to 2,700 kg/m3, whereas dolomite has a den- sity 2,870 kg/m3 and salt deposits have densities in the
4.5 Rock densities 75
1.5 2.0 2.5 3.0 1.5 2.0 2.5 3.0
0
1
2
3
4
5
z (km)
Density (kg/m3 × 103)
a b c
FIGURE 4.6 Density of sedimentary rocks vs. depth (z) for the North German–Polish basin. The left panel includes sandstones and siltstones; the right panel gives the mean values for the
(a) Quaternary, Cretaceous, Jurassic, (b) Bunter sandstone (Lower Triassic), and (c) Permian stratigraphic units. Adapted fromS c h ¨o n (1996).
range 2,000 to 2,200 kg/m3. Methane hydrate, which may occur in deep oceanic sediments, has a density of the order of 900 kg/m3.
Densities of sedimentary rocks increase with depth rapidly at shallow depths and less rapidly at progressively greater depths owing to lithostatic pressure as well as lithi- fication.Athy (1930) and others have shown that this variation can be approximated by an exponential function with a negative exponent. Densities also generally increase with the age of the sedimentary rocks. Older rocks tend to occur at greater, depth or have at one time been at greater depth, and thus have undergone greater compaction and lithification which leads to increased densities. Examples of increasing density with depth and age are shown in Figures 4.6 and 4.7. Note that in the right panel of Fig- ure 4.6, the older rocks of (b) and (c) do not increase in density as rapidly at shallower depths as younger rocks, probably as a result of the lower compaction possible in the older rocks. Figure 4.7 was assembled byMooney and Kaban(2010) from drillhole logs in the upper 2 to 3 km and extended to greater depths with seismic constraints. In this figure, the density variation with depth in the Michigan and Illinois basins, which contain significant quantities of carbonate rocks, is much less than in the offshore basins which consist largely of poorly lithified relatively young sediments subject to compaction with depth.
2.7 2.6 2.5 2.4 2.3 2.2 2.1 2
2
0 4 6 8 10 12 14 16
Depth (km) Density (kg/m3 × 103 )
Density of sediments vs. depth Great Valley Offshore basins Most of intr. bas.
Will. + Powd. Riv.
Mich. + Illin.
Some Cord. bas.
FIGURE 4.7 Smoothed density of sedimentary rocks vs. depth for several sedimentary basins on the North American continent including some Cordilleran (Cord.) basins; Michigan and Illinois basins (Mich.+Illin.); Williston and Powder River basins (Will.+ Powd.); most intracratonic basins; Great Valley of California; and offshore basins. Adapted fromM o o n e yandK a b a n(2010).
4.5.2 Lithostatic pressure and void space
Lithostatic pressure derived from the weight of overlying rocks has a profound effect on the density of subsurface materials, particularly sediments and sedimentary rocks.
This is primarily due to the decrease in void space with increasing pressure. As explained above, both primary and secondary void space of crystalline rocks are minimized with increasing pressure leading to higher densities with the process being especially effective in igneous ash flows and lava flows and near the surface.
Numerous relationships have been developed to describe the changing density with depth and lithostatic pressure. One form of this is
σP2=σP1{1+[(P2−P1)/K]}, (4.9) whereσP1andσP2are the respective densities of rocks in common units at pressuresP1andP2in pascals (Pa) and Kis the bulk modulus of the rock in Pa. The assumed con- stantKfor the upper crust is roughly 52 GPa, whereas the assumed values for the lower crust and upper mantle are approximately 75 and 130 GPa, respectively (Dziewon- skiandAnderson, 1981).
Void space in sediments and sedimentary rocks has a variety of origins and tends to be highly variable account- ing for the observed wide range of densities (e.g.Hall and Hajnal, 1962;Eaton and Watkins, 1970) of