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Chapter 11 Abnormally low formation pressures
295
Chapter 11 ABNORMALLY LOW FORMATION PRESSURES V.A. SEREBRYAKOV, G.V. CHILINGAR and J.O. ROBERTSON JR.
INTRODUCTION
The existence of underpressured fluid chambers has enormous significance to oil and gas exploration and production in the world. Such fluid compartments are determinative elements for undetected hydrocarbon traps (i.e., so-called subtle traps). Traditionally, most hydrocarbon production in the world has been from conventional structural and stratigraphic traps. Traps of a newly identified type, the underpressured hydrocarbon traps, may evolve from conventional traps as a result of changes in temperature and pressure. This kind of underpressured traps can be created by considerable overburden removal and local temperature change due to uplift and erosion giving rise to decreased pore pressure. Chapter 11 is an outline of a theoretical basis for an investigation of the validity of this concept. Examples of these kinds of traps can be found in the Denver and Oklahoma basins (Russell, 1972), the Alberta Basin (Hitchon, 1969), and the Volga-Ural and Middle Kura basins (Dobrynin and Serebryakov, 1989). It is important to construct a model for potential hydrocarbon sources in geological sections with abnormally low pressure. This model can be constructed using the change of rock temperature in local zones with significant uplift and erosion. For estimating the thicknesses of eroded deposits the authors used the method of compression curves, that indicate the presence of unconformities and the thickness of eroded deposits not only near the surface, but also deep in the geologic section. The modeling of subsurface underpressured zones may indicate a technique for making underpressured compartment traps viable exploration targets, because exploration strategies can be made significantly more effective if the mechanism of their formation is well understood. Subnormal pressures were discussed in three very important books: Gurevich et al. (1987), Dobrynin and Serebryakov (1989) and Dobrynin and Kuznetsov (1993). Gurevich et al. (1987) briefly analyzed all possible mechanisms of pressure subnormality, including permafrost degradation, fast leakage of gas from gas pools, decrease in temperature, formation of gas hydrates, expansion of rocks owing to the reduction in the overburden caused by erosion and disappearance of ice cover, etc. These authors believed that temperature decrease is possibly the most common cause of subnormally low pressures and occurs in formations overlain by permafrost. It does not include cases of pressure subnormality caused by distribution of the piezometric head in deep layers under the highest portions of the Earth's surface and by fluid withdrawal. Subnormal pressure distributions in several regions were analyzed. Dobrynin and Serebryakov (1989) believed that the major origin of subnormally low pressures is associated with permafrost and temperature changes. Subnormal pressures
296
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
in the Siberian Platform and their origin due to temperature decrease in geologic section were analyzed quantitatively. Using the regional situation as an example of the origin of pressure subnormality, Dobrynin and Serebryakov introduced the concept of thermodynamic gradient. They took the bottom surface of the permafrost (instead of the watertable surface) as the reference surface to determine the hydrostatic pressure in formations below the permafrost. Dobrynin and Kuznetsov (1993) concentrated on the thermodynamic gradient and its role in originating the pressure subnormality. They analyzed, from the viewpoint of this concept, different aspects of pressure subnormality origin, migration and accumulation of hydrocarbons, and sealing properties of formations with higher pressures than pressures of the lower formations. These authors also took into account the possible viscoplastic property of water in narrow pore channels and its influence on the water flow through formations.
ORIGIN OF A B N O R M A L PRESSURES
The origin and characteristics of abnormally low pressure may be related to regional phenomena, or it can be local. Different origin was ascribed by different scientists. For example, Berry (1959), Hill et al. (1961) and Breeze (1970) attributed the abnormally low pressure of the Alberta (Canada), San Juan (New Mexico and Colorado) basins and the Morrow sands of northern Oklahoma to osmotic-pressure differences. Russell (1972) in the Appalachian region, and A.G. Durmishyan (pers. commun., 1976) in the northern Caucasus described the low-pressured reservoirs in well-consolidated rocks, which have been uplifted and eroded in the geologic past. Barker (1972) ascribed the subnormal pressure to the removal of overburden, which would cause a drop in pressure of the pore fluids. Erosional unloading has been suggested to explain certain abnormally low pressures by Louden (1972) and by Dickey and Cox (1977), although a quantitative analysis of the process has not yet been presented. Later, Neuzil and Pollock (1983), using the mass-balance equation (Domenico and Palciauskas, 1979) for water and grains in a small control volume of saturated porous medium, described the unloading of saturated elastic rocks caused by decreasing thickness. Thermal effects, however, were not included. In the opinion of the writers, the thermal effects play the main role in pore pressure changes. Basically, the influence of temperature on pore pressure is strongest in regions with more compacted rocks. In undercompacted, plastic rocks with a high coefficient of compressibility, thermal expansion of fluids can be compensated by deformation of rock pore spaces. Neuzil and Pollock (1983) noted the same effect for the overburden removal. They stated that only in rocks of low permeability is pore pressure likely to be affected by erosion, and if the permeable unit is effectively isolated by surrounding 'tight' rocks, erosional unloading could cause pressure lowering. Changes in temperature of rocks could occur in two ways. The first (regional one), caused by the changes in the temperature at the surface of the Earth in geologic time, affects usual and unusual fluid filtrations. The second (local one), can be caused by changes in the temperature of rocks during the significant uplift and erosion, or subsidence and aggradation. The first phenomenon is a fundamental process in the
297
ABNORMALLY LOW FORMATION PRESSURES
Earth's crust, whereas the second one is a local process in some areas. But both of these processes are associated with well-compacted rocks. In this chapter, the authors discuss only the second phenomenon. In order to estimate and compare compaction of rocks in different basins and regions, the writers used the coefficient of irreversible compaction of rocks /3(t, T), where t is time and T is temperature. The change in porosity of a sedimentary rock with depth can be presented as follows (Dobrynin, 1970): 0(~p = /3 (t, T) x d(~ - p) (1 -- ~p)~p
(11- l)
where a is the overburden pressure, p is the pore pressure, (or - p) is the effective stress, and ~bp is the value of porosity in the highest part of the interval of interest. Effective stress can be estimated as follows: (or -
p) = g(Pr -
(11-2)
pw)h
where g is the gravitational acceleration, fir is the average rock density, Pw is the average density of water, and h is the depth. Using fir 2.5 g/cm 3 and Pw -- 1.1 g/cm 3 as the average density of rocks and water, respectively, one can estimate the average effective stress for a geological section as (c~ - p) = 0.014 h, and determine the coefficient of irreversible compaction/3 (t, T) as: -
~(t, T) ~
1
Aq~p
0 . 0 1 4 ( 1 -- ~bp)~bp A h
-
(11-3)
where A~bp/Ah is the porosity gradient in the depth interval of interest. One can estimate and compare the coefficient of irreversible compaction of shales /3(t, T) in different basins. On comparing shale porosity-depth relationships of the Gulf Coast (Dickinson, 1953), the Oklahoma Basin (Athy, 1930), the West Kuban Depression (Popov, pers. commun., 1970), the North Caspian Basin (Dobrynin and Serebryakov, 1978) and the Powder River Basin, one can observe that shale porositydepth relationships in various basins are quite different (also see Rieke and Chilingarian, 1974). Using these data and Eq. 11-3, one can estimate the coefficient of irreversible compaction for these four basins (Table 11-1). In three of these basins (Gulf Coast, Oklahoma and West Kuban Depression), the coefficient of irreversible compaction varies greatly with depth. There is a fourfold change with a depth of 2 km in the Gulf Coast, almost twice in Oklahoma, and one and a half times in the West Kuban Depression. This variation shows that rocks in these basins are not fully compacted and that the compaction of these rocks may be continuing even now. Only two basins in Table 11-1, the North Caspian and Powder River, have highly compacted rocks, as shown by the fact that the coefficient of irreversible compaction is not changing much with depth. Compaction of these rocks has stopped. The probability of creating abnormal pressure in the more compacted, isolated rocks of these basins is significantly greater, because the influence of temperature on pore pressure in these rocks is significantly greater. Unusual underpressured hydrocarbon traps, and the seals that isolate them, result from global temperature change at the Earth's surface or from local temperate change
298
V.A.SEREBRYAKOV,G.V.CHILINGARANDJ.O. ROBERTSONJR.
TABLE 11-1 Coefficient of irreversible compaction for various basins Basin (source)
Depth (m)
Coefficient of irreversible compaction/~ (t, T) x 103 (mPa-~ )
Gulf Coast, USA (Dickinson, 1953)
0 1000 2000 3000 4000
0 101.5 24.6 14.8 14.5
Oklahoma (Athy, 1930)
0 1000 2000
98.0 51.0
West Kuban Depression, Russia (Popov, 1970, personal communication)
0 1000 2000 3000
66.0 42.0 29.0
Caspian, Russia (Dobrynin and Serebryakov, 1978)
0 1000 2000 3000 4000
28.2 21.1 26.6 26.3
Powder River
0 1000 2000 3000 4000
30.4 30.9 25.0 24.4
N.
due to significant subsidence and aggradation, or uplift and erosion. The existence of abnormally low pressured zones due to global temperature change at the Earth's surface was first observed in eastern Siberia at a depth of 2.0-2.5 km near the crystalline basement by Dobrynin and Serebryakov (1989). This phenomenon was also discovered in other parts of the former USSR (C.I.S.): the Volga-Ural Province, northern part of West Siberia, Ukraine and Georgia. Mostly, this phenomenon is attributed to the changes in temperature at the Earth's surface during geologic time, with consequent changes in hydrogenetic processes in geologic sections with compacted rocks (eastern and western Siberia). In the Ukraine and Georgia, abnormally high and abnormally low pressures are caused by subsidence and uplift, respectively. The theoretical basis of this phenomenon can be analyzed as follows. At a certain time in the basin's history, there was a hydrologic equilibrium defined by" p l _ gPw (h - hst)
(11-4)
where pl is the pore pressure, g is the gravitational acceleration, Pw is the average density of water, h is the depth, and hst is the depth of the static water level. At a later time, the depth changed. The overburden pressure ~r changed because of subsidence,
299
ABNORMALLY LOW FORMATION PRESSURES
permafrost, or tectonic uplift and erosion. The temperature changed in response to the change in the global Earth's surface temperature or to the change in depth. Such changes may create volumetric changes in (1) rocks, (2) pore space, and (3) the interstitial water. Relative pore volume change d V p / V p and pore water volume change dVw/Vw with respect to a change in temperature can be expressed as follows (Dobrynin, 1970):
Vp
- - j~r d(a - p) + fis dp - as dT
(11-5)
dVw = flw dp - O~wdT Vw
(11-6)
where fir, fis, and fiw, are the coefficients of compressibility of pore volume, solids and water, respectively; and O~s, and O~w are the coefficients of thermal expansion of solids (minerals) and water, respectively. In the case of static pore water (full hydrodynamic isolation of pores), a relative change in the volume of pores must be equal to the relative change of pore water volume: dVr
Vr
dVw
=
(11-7)
Vw
Equating the right-hand sides of Eqs. 11-5 and 11-6, one obtains" flw dp - OewdT -- fir d(a - p) + fis dp - Ors dT
(11-8)
Thus, the change in pore pressure (dp) that occurs during a change in the thermodynamic conditions of a deposit is equal to: fir
dp -
da +
/~r -'[- flw -- fls
lYw -- tYs
dT
11-9
fir -[- flw -- fls
2 1 The average normal stress (a) for horizontal layers is a -- 5ax + 5az, where ax is the horizontal component of stress (and ax -- ay), and az is the vertical component of stress. Using the mean lateral compression: l)
Crx -- Oy = ~1 _ a v z
(11-10)
The total overburden stress is equal to" -
a
-
3
1-v
az
(11-11)
where v is Poisson's ratio. Differentiating Eq. 11-11 one obtains: da--~
1 [l+V]d, 1-v
(11-12)
where da z is the change in the vertical component of stress during sedimentation or erosion. Expressing da z in terms of change in sedimentary overburden, one obtains: da-~
1 [l+VlgprAh
1-v
(11-13)
300
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
where Ah is the thickness of rocks accumulated during subsidence or thickness of eroded deposits during uplift, and Dr is the average density of the newly deposited or eroded rocks. Combining Eq. 11-13 with Eq. 11-9, one obtains: dP--•
ill+v] t..jl-v
fir
gprAh+
fir nt- flw - fls
Otw - Ors fir + fiw --/3s d r
(11-14)
where plus (+) designates the change in pressure due to subsidence, whereas minus is used to designate the change in pressure due to uplift and erosion. Usually, for sedimentary rocks/~r -+" /~w >) /~s and O~w >> ot~. Thus, Eq. 11-14 can be simplified to: l[l+v] fir gprAh+ Otw dp -- 4-~ 1 - V fir + fl--------~ fir +/3--------~dT
(11-15)
where + and - are treated as in Eq. 11-14. Combining Eqs. 11-4 and 11-15, the new value of pore pressure (p) after thermodynamic change can be estimated as follows: 1 [l+v] p = p l _4_ d p -- g p w ( h
- hst) 4- 5
1 - v
/3r /~r nt- /~w
gprAh+
Otw fir +/3---------~dT(ll_16 )
where pl is the original pore pressure. Using Eq. 11-16, one can estimate pressure after a change in pore pressure in isolated pores, when rock temperature and stress have changed because of sedimentation due to subsidence or erosion owing to uplift, or due to the creation of permafrost. It is necessary to pay special attention to the coefficient of thermal expansion of water, Otw, which increases with increasing temperature and, therefore, with depth. On the basis of the experimental data (Vukolovich et al., 1969), the equation for thermal expansion of fresh water was determined in the temperature interval of 5~ to 200~ to be (Dobrynin and Serebryakov, 1989): Otw = (0.694T ~
1.446) x 10-4
(11-17)
This equation shows dependence of the thermal expansion coefficient on temperature, which is most important in its influence on pore pressure.
ESTIMATION OF THE EFFECTS OF TEMPERATURE CHANGE AND EROSION ON PORE PRESSURE
Serebryakov and Chilingar (1994) discussed the effects of temperature change and erosion on pore pressure in the northem part (Recluse area) of the Powder River Basin (T55-58N, R73-76W) (Fig. 11-1). DST data from 59 wells showed the existence of abnormally low pressure in the Cretaceous deposits at a depth of 7200-7900 ft (2-2.4 km), where the abnormal pressure coefficient K, = P" ranges from 0.4 to 0.8 (Tables 11-2 and 11-3); Pa is the abnormal pressure and Pn Pls the normal hydrostatic pressure. Using the process-oriented conceptual model, presented above, for the generation of abnormal pressure, one can explain the existence of underpressured reservoirs in this
301
ABNORMALLY LOW FORMATION PRESSURES
Hammond
__..__.__._. _ _ _ _ ~Creek F e r " c~. e
--•T
Bell Creek
MONTAN_..AA
9 S
T 57 N R64W
Gas Draw
Springen Ranch
T R64w
L~yB
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R61W
i
T.
45 N 1 |
CJareton Ar~
Sfeinle Ranch
T 58 N R61W
WYOMIN
R8Zw
R71W
y
,
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area by considerable overburden removal and local temperature change owing to uplift and erosion. One can use Eq. 11-16, which consists of three parts. The first part is: g p w ( h - hst)
(11-18)
which is normal hydrostatic pressure at a certain depth before uplift and erosion. The second part is:
, I,+vl 1 - v
-3
-[-/3--------~ j~r g pr A h
(11-19)
The latter shows the influence of overburden removal on pore pressure. The third part is: O~W
dT
(11-20)
fir -'~--j~ w
which describes the changes of pore pressure due to the local temperature changes during uplift and erosion. Dobrynin and Serebryakov (1989) suggested that subnormal pressure could be predicted as the algebraic sum of hydrostatic pressure and pressure change in closed pores due to temperature change. According to these authors, for formations below the permafrost, the hydrostatic pressure is equal to the weight of the water column from the permafrost bottom down to the point where pressure is to be determined. They applied this approach to the examples from the Siberian Platform and Kura Region (Caucasus). The same approach, but at a greater detail, was presented by Dobrynin and Kuznetsov (1993). The first step in studying the underpressured zones in the Powder River Basin was to estimate the thickness of eroded rocks in the target area. The authors used the method of compression curves. Compression curves show the relationship between different parameters (porosity, density, resistivity, transit time, etc.) and effective stress (i.e., overburden pressure minus pore pressure). The parameters n (slope of straight line), and m (intercept) of the compression curve are very important because of their geologic significance. Slope n~ characterizes the compaction of that rock as a function of geologic age, mineralogy, etc. This parameter has to be estimated for each area under investigation or perhaps for each well, using normal compaction trend. Parameter mx (y-intercept), at the beginning of the compaction curve, enables one to estimate the physical rock property of interest near the surface, where effective stress (overburden pressure minus pore pressure) is zero. This parameter depends on the attributes of the entire geologic section, including eroded deposits and major unconformities. It has to be estimated in each well, using the normal compaction trend. The equation for estimating the thickness of an eroded deposit is as follows: Aher =
m x -- mxl g(Pr- pw)nx
(11-21)
where the taxi parameter is for the compression curve in a section without erosion; nx and mx parameters are for the compression curve of interest; g is the gravitational acceleration; Pr is the average density of the rocks, and Pw is the average density of water. Serebryakov and Chilingar (1994) estimated parameters of compression curves
ABNORMALLY LOW FORMATION PRESSURES
305
(Table 11-2) and thicknesses of eroded deposits (Table 11-3) in ten wells of the Powder River Basin using resistivity data, and in five of these wells using sonic data. For the estimation of thicknesses of eroded deposits, it is very important to know the value of parameter mx~ of compression curve in a geologic cross-section devoid of erosion. For the estimation of this parameter, one has to know values of geophysical data (transit time or resistivity) near the surface. In geologic sections without erosion, the value of transit time is close to 200 ~s/ft (660 ~s/m) (Magara, 1978). Using this value, one can estimate the value of mx~ in the cross-section without erosion: mx~ = 3.3. For resistivity value, the authors used estimates that had been made in Russia (Alexandrov, 1987; Dobrynin and Serebryakov, 1989): the value of shale resistivity near the surface in geologic section without erosion is 0.8 ohm m. This value was used for the Powder River Basin. The mx~ value used in calculating the amount of erosion was 0.033. An earlier study of sonic logs in the Powder River Basin (Fig. 11-2a: area without abnormal pressure; Fig. l l-2b: area with abnormally high pressure) showed some difficulty in estimating the thicknesses of eroded deposits by extrapolating the trend of normally compacted shale in terms of sonic travel time in geologic sections without erosion (Magara, 1978). In the Powder River Basin there is no single exponential relationship between transit time and depth. Instead, there are two relationships: one at a depth below 3000-3500 ft, and the other near the surface (Fig. 11-2). The change of sonic trend at shallow depth is possibly related to a significant change in pore tortuosity. Serebryakov and Chilingar (1994) had less difficulty in estimating the thickness of eroded deposits using the resistivity normal compaction trend (Fig. 11-3). It is necessary to correct for the influence of water salinity change near the surface when plotting the resistivity normal compaction trend (Dobrynin and Serebryakov, 1989). Values of estimated thicknesses of eroded deposits vary from 3707 ft (1130 m) to 5141 ft (1567 m). Values of eroded thicknesses estimated by using sonic data are greater in the majority of wells, because there were not enough sonic data at shallow depth and the authors had to use the normal trend of deep deposits. In one well (#1 Southland Govt.), these authors obtained a lower value of eroded thickness using sonic data, but these sonic data were not of good quality. Another way to estimate the approximate thickness of eroded deposits is to determine shale density near the surface. Unfortunately, core data are not available for the near-surface sediments in the target area. Serebryakov and Chilingar (1994), however, estimated the grey shale density of outcrop samples (Oedekoven area) near the surface: 2.35 g/cm 3. The shale density at a depth of 300-400 ft (90-120 m) was estimated to be 2.2-2.3 g/cm 3, using density logs. In geologic sections without erosion, such values of shale density usually denote a depth of 3300-5000 ft (1000-1500 m) (Alexandrov, 1987). These results are thus indirect confirmation of the erosion of 1000-1500 m of overburden in this area. Serebryakov and Chilingar (1994) estimated the influence of overburden removal and temperature change, owing to erosion, from the thicknesses of eroded deposits (resistivity log data). Using the second part of Eq. 11-16, the decrease in pressure due to the removal of overburden was determined. These authors used a value of 0.25 for the Poisson ratio for compacted rocks (Means, 1985), a value of 1 x 10 -3 for the coefficient of pore volume compressibility, a value of 0.5 x 10 -3 for compressibility of water
306
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
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(Dobrynin and Serebryakov, 1989), and estimated value of average density of eroded deposits, ,Or, and an estimated thickness of eroded deposits a h (Table 11-2). Values for the pressure decrease (Table 11-3) are 1180-1689 psi (8.3-11.9 MPa). These authors used the third part of Eq. 11-16 for estimating the decrease in pore pressure owing to temperature change. The average geothermal gradient in this area is 0.03~ m. Due to the erosion (1000-1500 m), the temperature could have changed 30~176 The main parameter in the third part of Eq. 11-16 is the coefficient of thermal expansion of water. To determine this coefficient, the authors used Eq. 11-17 for different values of temperature before uplift and erosion. Average values of this coefficient, which was changing during geologic time, were used in Eq. 11-16. Values of pore-pressure decrease due to temperature change vary from 1565 to 2233 psi (11 to 15.7 MPa; Table 11-3).
307
ABNORMALLY LOW FORMATION PRESSURES
3000
A
4000 m~
c-
5000
I I I I I I I I I !
O. 0
6000
7000
v a
8000
I
lo0
10
1000
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4000
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"k
r
5000
O. (!)
l $ l $
6000
l I
,f
7000
o~Im
8000
i
l0
100
1000
Shale transit time, las/ft Fig. 11-3. Relationship between geophysical parameters and depth in the Arco Kendrick 3-47 well: (A) shale transit time, and (B) shale resistivity. (Modified after Serebryakov and Chilingar, 1994, fig. 3, p. 258.) On using the values of decrease of pore pressure due to overburden removal and temperature change, Serebryakov and Chilingar (1994) estimated the values of abnormally low pressure in each well owing to uplift and erosion. Then the estimated coefficients of abnormal pressure were compared with those derived from the DST data (Table 11-3). Results are close, but the estimated coefficients of abnormal pressure are usually lower than those derived from DST data. One possible explanation for this discrepancy is that during geological time the pore pressure was changing due to leaking
308
V.A. SEREBRYAKOV,G.V. CHILINGARAND J.O. ROBERTSONJR.
seals. These calculations show the possibility of existence of underpressured zones not only in the Recluse area but also elsewhere in the Powder River Basin. Overburden removal and temperature change could both cause decrease in pore pressure. One more significant condition, however, must exist for the existence of underpressured zones at present: good seals having low permeability, which could hold pore pressure during geologic time. One can evaluate the sealing capacity of Cretaceous shale in the target area using the sonic travel time data that have been used by Magara (1978) for the Alberta and Saskatchewan areas. Using the linear relationship between the shale porosity (4~) and transit time (At) in the Cretaceous shale (4~ = 0.00466 At --0.317), Serebryakov and Chilingar (1994) estimated the porosity of shale in the Recluse area at a depth of 7000-8000 ft to be 10-13%. Using the porosity-permeability relationship for the Cretaceous shale (Magara, 1978), the permeability of the shale was found to be less than 5 x 10 -3 mD. These shales could act as good seals and could hold pore pressure over a long period of geologic time in the absence of fracturing. As a rule, in the zones where a seal has been fractured, there is normal hydrostatic pressure. Underpressured zones are not present where liquid hydrocarbons have been converted to natural gas, with creation of abnormally high pore pressures. Such examples can be found in the southern part of the Powder River Basin.
SUMMARY
Underpressured reservoirs could form as a result of removal of overburden (erosional unloading). Thermal effects (decrease in temperature) could play a major role in causing underpressure in well-compacted rocks. In estimating the thickness of eroded deposits, the writers recommend the use of "compression curves method" (Dobrynin et al., 1982). Underpressured hydrocarbon reservoirs in the Powder River Basin of Wyoming and Montana have been studied. A significant amount of research work, however, still remains to be done in this field in order to reach definite conclusions.
BIBLIOGRAPHY Abasov, M.T., Azimov, E.Kh., Aliyarov, P.Yu. et al., 1991. The Theory and Practice of GeologicGeophysical Exploration and Development of the Offshore Oil and Gas Fields. Elm Publ., Baku, Azerbaijan, 428 pp. Alexandrov, B.L., 1987. Abnormally High Formation Pressures in Oil-Gas-Bearing Basins. Nedra Publ., Moscow, 216 pp. Athy, L.E, 1930. Density, porosity and compaction of sedimentary rocks. Am. Assoc. Pet. Geol. Bull., 14: 1-24. Barker, C., 1972. Aquathermal pressuring role of temperature in development of abnormal pressure zones. Am. Assoc. Pet. Geol. Bull., 56: 2068-2871. Berry, E, 1959. Hydrodynamics and Geochemistry of the Jurassic and Cretaceous System in the San Juan Basin, Northwest New Mexico and Southwestern Colorado. Ph.D. Thesis, Stanford University, Stanford, CA, 213 pp.
ABNORMALLYLOW FORMATIONPRESSURES
309
Breeze, A., 1970. Abnormal-Subnormal Relationships in the Morrow Sands of Northwestern Oklahoma. M.Sc. Thesis, University Oklahoma, Tulsa, OK, 122 pp. Dickey, L. and Cox, W., 1977. Oil and gas reservoirs with subnormal pressure. Am. Assoc. Pet. Geol. Bull., 61: 2134-2142. Dickinson, G., 1953. Geological aspects of abnormal reservoir pressures in Gulf Coast Louisiana. Am. Assoc. Pet. Geol. Bull., 37: 410-432. Dobrynin, V., 1970. Deformation and Physical Properties Change in the Oil and Gas Reservoir Rocks. Nedra, Moscow, 288 pp. Dobrynin, V.M. and Kuznetsov, O.L., 1993. Thermoelastic Processes in the Rocks of Sedimentary Basins. VNII Geosystem, Moscow, 169 pp. Dobrynin, V. and Serebryakov, V., 1978. Methods for the Prediction of Abnormally High Formation Pressures. Nedra, Moscow, 231 pp. (in Russian.) Dobrynin, V. and Serebryakov, V.A., 1989. Geological Geophysical Methods for Prediction of Pressure Anomalies. Nedra, Moscow, 287 pp. Dobrynin, V.M., Serebryakov, V. and Srebrodolskiy, A., 1982. Determination of abnormally high formation pressures in shales using the method of compression curves. Geol. Neftii Gaza, 5: 25-28. Domenico, E and Palciauskas, A., 1979. Thermal expansion of fluids and fracture initiation in compacting sediments. Geol. Soc. Am. Bull., 90: 953-979. Fertl, W., 1976. Abnormal Formation Pressures. Implications in Exploration, Drilling and Production of Oil and Gas Resources. Elsevier, Amsterdam, 382 pp. Gurevich, A.E., Batygina, N.B. and Kraichik, M.S. et al., 1987. Formation Fluid Pressure. Nedra Publ., Leningrad, 223 pp. Gurevich, A.E., Chilingar, G.V. and Aminzadeh, E, 1994. Origin of the formation fluid pressure distribution and ways of improving pressure prediction methods. J. Pet. Sci. Eng., 12: 67-77. Hill, G., Calburn, W. and Knight, J., 1961. Reducing Oil Finding Cost by Use of Hydrodynamic Evaluation. Economics of Petroleum Exploration, Development, and Property Evaluation. Prentice-Hall, Englewood Cliffs, CA, 380 pp. Hitchon, B., 1969. Fluid flow in the western Canada sedimentary basin, 2. Effect of geology. Water Resour. Res., 5: 460-469. Hubbert, M.K. and Rubey, W.W., 1959. Role of fluid pressure in mechanics of overthrust faulting. 1. Mechanics of fluid-filled porous solids and its application to overthrust faulting. Geol. Soc. Am. Bull., 70: 115-166. Kazimirov, D.A., 1974. Impulse tectonic movements. Geotectonics, 4: 19-32. Louden, I., 1972. Origin and maintenance of abnormal pressure. 3rd Symp. Abnormal Subsurface Pore Pressure. Soc. Pet. Eng. AIME, pp. 23-27. Magara, K., 1978. Compaction and Fluid Migration, Practical Petroleum Geology. Elsevier, Amsterdam, 319 pp. Means, W.D., 1985. Stress and Strain, Basic Concepts of Continuum Mechanics for Geologists. Springer, Berlin, 339 pp. Melik-Pashaev, V.S., Khalimov, E.M. and Seregina, V.N., 1983. Abnormally High Formation Pressures in Oil and Gas Fields. Nedra Publ., Moscow, 181 pp. Neuzil, C. and Pollock, D., 1983. Erosional unloading and fluid pressures in hydraulically 'tight' rocks. J. Geol., 91: 179-193. Rieke, H.H. III and Chilingarian, G.V., 1974. Compaction of Argillaceous Sediments. Elsevier, Amsterdam, 424 pp. Rubey, W.W. and Hubbert, M.K., 1959. Role of fluid pressure in mechanics of overthrust faulting. 11. Overthrust belt in geosynclinal area of Western Wyoming in light of fluid pressure hypothesis. Geol. Soc. Am. Bull., 70: 167-200. Russell, W., 1972. Pressure-depth relations in Appalachian region. Am. Assoc. Pet. Geol. Bull., 56: 528-536. Serebryakov, V.A. and Chilingar, G.V., 1994. Investigation of underpressured reservoirs in the Powder River Basin, Wyoming and Montana. J. Pet. Sci. Eng., 11: 249-259. Terzaghi, K., 1965. Theoretical Soil Mechanics. Wiley, New York, NY, 510 pp. Tkhostov, B.A., 1963. Initial Rock Pressures in Oil and Gas Deposits. Translated from Russian by R.A. Ledward. Macmillan, New York, NY, 118 pp.
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V.A. SEREBRYAKOV,G.V. CHILINGAR AND J.O. ROBERTSONJR.
Toth, J. and Corbet, T., 1986. Post-Paleocene evolution of regional ground water flow systems and their relation to petroleum accumulations, Taber area, Southern Alberta, Canada. Bull. Can. Pet. Geol., 34: 339-363. Vukolovich, M., Rivkin, S. and Alexandrov, A., 1969. Tables of Physical Properties of Water and Vapor. Standart, Moscow.
Chapter 11 ABNORMALLY LOW FORMATION PRESSURES V.A. SEREBRYAKOV, G.V. CHILINGAR and J.O. ROBERTSON JR.
INTRODUCTION
The existence of underpressured fluid chambers has enormous significance to oil and gas exploration and production in the world. Such fluid compartments are determinative elements for undetected hydrocarbon traps (i.e., so-called subtle traps). Traditionally, most hydrocarbon production in the world has been from conventional structural and stratigraphic traps. Traps of a newly identified type, the underpressured hydrocarbon traps, may evolve from conventional traps as a result of changes in temperature and pressure. This kind of underpressured traps can be created by considerable overburden removal and local temperature change due to uplift and erosion giving rise to decreased pore pressure. Chapter 11 is an outline of a theoretical basis for an investigation of the validity of this concept. Examples of these kinds of traps can be found in the Denver and Oklahoma basins (Russell, 1972), the Alberta Basin (Hitchon, 1969), and the Volga-Ural and Middle Kura basins (Dobrynin and Serebryakov, 1989). It is important to construct a model for potential hydrocarbon sources in geological sections with abnormally low pressure. This model can be constructed using the change of rock temperature in local zones with significant uplift and erosion. For estimating the thicknesses of eroded deposits the authors used the method of compression curves, that indicate the presence of unconformities and the thickness of eroded deposits not only near the surface, but also deep in the geologic section. The modeling of subsurface underpressured zones may indicate a technique for making underpressured compartment traps viable exploration targets, because exploration strategies can be made significantly more effective if the mechanism of their formation is well understood. Subnormal pressures were discussed in three very important books: Gurevich et al. (1987), Dobrynin and Serebryakov (1989) and Dobrynin and Kuznetsov (1993). Gurevich et al. (1987) briefly analyzed all possible mechanisms of pressure subnormality, including permafrost degradation, fast leakage of gas from gas pools, decrease in temperature, formation of gas hydrates, expansion of rocks owing to the reduction in the overburden caused by erosion and disappearance of ice cover, etc. These authors believed that temperature decrease is possibly the most common cause of subnormally low pressures and occurs in formations overlain by permafrost. It does not include cases of pressure subnormality caused by distribution of the piezometric head in deep layers under the highest portions of the Earth's surface and by fluid withdrawal. Subnormal pressure distributions in several regions were analyzed. Dobrynin and Serebryakov (1989) believed that the major origin of subnormally low pressures is associated with permafrost and temperature changes. Subnormal pressures
296
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
in the Siberian Platform and their origin due to temperature decrease in geologic section were analyzed quantitatively. Using the regional situation as an example of the origin of pressure subnormality, Dobrynin and Serebryakov introduced the concept of thermodynamic gradient. They took the bottom surface of the permafrost (instead of the watertable surface) as the reference surface to determine the hydrostatic pressure in formations below the permafrost. Dobrynin and Kuznetsov (1993) concentrated on the thermodynamic gradient and its role in originating the pressure subnormality. They analyzed, from the viewpoint of this concept, different aspects of pressure subnormality origin, migration and accumulation of hydrocarbons, and sealing properties of formations with higher pressures than pressures of the lower formations. These authors also took into account the possible viscoplastic property of water in narrow pore channels and its influence on the water flow through formations.
ORIGIN OF A B N O R M A L PRESSURES
The origin and characteristics of abnormally low pressure may be related to regional phenomena, or it can be local. Different origin was ascribed by different scientists. For example, Berry (1959), Hill et al. (1961) and Breeze (1970) attributed the abnormally low pressure of the Alberta (Canada), San Juan (New Mexico and Colorado) basins and the Morrow sands of northern Oklahoma to osmotic-pressure differences. Russell (1972) in the Appalachian region, and A.G. Durmishyan (pers. commun., 1976) in the northern Caucasus described the low-pressured reservoirs in well-consolidated rocks, which have been uplifted and eroded in the geologic past. Barker (1972) ascribed the subnormal pressure to the removal of overburden, which would cause a drop in pressure of the pore fluids. Erosional unloading has been suggested to explain certain abnormally low pressures by Louden (1972) and by Dickey and Cox (1977), although a quantitative analysis of the process has not yet been presented. Later, Neuzil and Pollock (1983), using the mass-balance equation (Domenico and Palciauskas, 1979) for water and grains in a small control volume of saturated porous medium, described the unloading of saturated elastic rocks caused by decreasing thickness. Thermal effects, however, were not included. In the opinion of the writers, the thermal effects play the main role in pore pressure changes. Basically, the influence of temperature on pore pressure is strongest in regions with more compacted rocks. In undercompacted, plastic rocks with a high coefficient of compressibility, thermal expansion of fluids can be compensated by deformation of rock pore spaces. Neuzil and Pollock (1983) noted the same effect for the overburden removal. They stated that only in rocks of low permeability is pore pressure likely to be affected by erosion, and if the permeable unit is effectively isolated by surrounding 'tight' rocks, erosional unloading could cause pressure lowering. Changes in temperature of rocks could occur in two ways. The first (regional one), caused by the changes in the temperature at the surface of the Earth in geologic time, affects usual and unusual fluid filtrations. The second (local one), can be caused by changes in the temperature of rocks during the significant uplift and erosion, or subsidence and aggradation. The first phenomenon is a fundamental process in the
297
ABNORMALLY LOW FORMATION PRESSURES
Earth's crust, whereas the second one is a local process in some areas. But both of these processes are associated with well-compacted rocks. In this chapter, the authors discuss only the second phenomenon. In order to estimate and compare compaction of rocks in different basins and regions, the writers used the coefficient of irreversible compaction of rocks /3(t, T), where t is time and T is temperature. The change in porosity of a sedimentary rock with depth can be presented as follows (Dobrynin, 1970): 0(~p = /3 (t, T) x d(~ - p) (1 -- ~p)~p
(11- l)
where a is the overburden pressure, p is the pore pressure, (or - p) is the effective stress, and ~bp is the value of porosity in the highest part of the interval of interest. Effective stress can be estimated as follows: (or -
p) = g(Pr -
(11-2)
pw)h
where g is the gravitational acceleration, fir is the average rock density, Pw is the average density of water, and h is the depth. Using fir 2.5 g/cm 3 and Pw -- 1.1 g/cm 3 as the average density of rocks and water, respectively, one can estimate the average effective stress for a geological section as (c~ - p) = 0.014 h, and determine the coefficient of irreversible compaction/3 (t, T) as: -
~(t, T) ~
1
Aq~p
0 . 0 1 4 ( 1 -- ~bp)~bp A h
-
(11-3)
where A~bp/Ah is the porosity gradient in the depth interval of interest. One can estimate and compare the coefficient of irreversible compaction of shales /3(t, T) in different basins. On comparing shale porosity-depth relationships of the Gulf Coast (Dickinson, 1953), the Oklahoma Basin (Athy, 1930), the West Kuban Depression (Popov, pers. commun., 1970), the North Caspian Basin (Dobrynin and Serebryakov, 1978) and the Powder River Basin, one can observe that shale porositydepth relationships in various basins are quite different (also see Rieke and Chilingarian, 1974). Using these data and Eq. 11-3, one can estimate the coefficient of irreversible compaction for these four basins (Table 11-1). In three of these basins (Gulf Coast, Oklahoma and West Kuban Depression), the coefficient of irreversible compaction varies greatly with depth. There is a fourfold change with a depth of 2 km in the Gulf Coast, almost twice in Oklahoma, and one and a half times in the West Kuban Depression. This variation shows that rocks in these basins are not fully compacted and that the compaction of these rocks may be continuing even now. Only two basins in Table 11-1, the North Caspian and Powder River, have highly compacted rocks, as shown by the fact that the coefficient of irreversible compaction is not changing much with depth. Compaction of these rocks has stopped. The probability of creating abnormal pressure in the more compacted, isolated rocks of these basins is significantly greater, because the influence of temperature on pore pressure in these rocks is significantly greater. Unusual underpressured hydrocarbon traps, and the seals that isolate them, result from global temperature change at the Earth's surface or from local temperate change
298
V.A.SEREBRYAKOV,G.V.CHILINGARANDJ.O. ROBERTSONJR.
TABLE 11-1 Coefficient of irreversible compaction for various basins Basin (source)
Depth (m)
Coefficient of irreversible compaction/~ (t, T) x 103 (mPa-~ )
Gulf Coast, USA (Dickinson, 1953)
0 1000 2000 3000 4000
0 101.5 24.6 14.8 14.5
Oklahoma (Athy, 1930)
0 1000 2000
98.0 51.0
West Kuban Depression, Russia (Popov, 1970, personal communication)
0 1000 2000 3000
66.0 42.0 29.0
Caspian, Russia (Dobrynin and Serebryakov, 1978)
0 1000 2000 3000 4000
28.2 21.1 26.6 26.3
Powder River
0 1000 2000 3000 4000
30.4 30.9 25.0 24.4
N.
due to significant subsidence and aggradation, or uplift and erosion. The existence of abnormally low pressured zones due to global temperature change at the Earth's surface was first observed in eastern Siberia at a depth of 2.0-2.5 km near the crystalline basement by Dobrynin and Serebryakov (1989). This phenomenon was also discovered in other parts of the former USSR (C.I.S.): the Volga-Ural Province, northern part of West Siberia, Ukraine and Georgia. Mostly, this phenomenon is attributed to the changes in temperature at the Earth's surface during geologic time, with consequent changes in hydrogenetic processes in geologic sections with compacted rocks (eastern and western Siberia). In the Ukraine and Georgia, abnormally high and abnormally low pressures are caused by subsidence and uplift, respectively. The theoretical basis of this phenomenon can be analyzed as follows. At a certain time in the basin's history, there was a hydrologic equilibrium defined by" p l _ gPw (h - hst)
(11-4)
where pl is the pore pressure, g is the gravitational acceleration, Pw is the average density of water, h is the depth, and hst is the depth of the static water level. At a later time, the depth changed. The overburden pressure ~r changed because of subsidence,
299
ABNORMALLY LOW FORMATION PRESSURES
permafrost, or tectonic uplift and erosion. The temperature changed in response to the change in the global Earth's surface temperature or to the change in depth. Such changes may create volumetric changes in (1) rocks, (2) pore space, and (3) the interstitial water. Relative pore volume change d V p / V p and pore water volume change dVw/Vw with respect to a change in temperature can be expressed as follows (Dobrynin, 1970):
Vp
- - j~r d(a - p) + fis dp - as dT
(11-5)
dVw = flw dp - O~wdT Vw
(11-6)
where fir, fis, and fiw, are the coefficients of compressibility of pore volume, solids and water, respectively; and O~s, and O~w are the coefficients of thermal expansion of solids (minerals) and water, respectively. In the case of static pore water (full hydrodynamic isolation of pores), a relative change in the volume of pores must be equal to the relative change of pore water volume: dVr
Vr
dVw
=
(11-7)
Vw
Equating the right-hand sides of Eqs. 11-5 and 11-6, one obtains" flw dp - OewdT -- fir d(a - p) + fis dp - Ors dT
(11-8)
Thus, the change in pore pressure (dp) that occurs during a change in the thermodynamic conditions of a deposit is equal to: fir
dp -
da +
/~r -'[- flw -- fls
lYw -- tYs
dT
11-9
fir -[- flw -- fls
2 1 The average normal stress (a) for horizontal layers is a -- 5ax + 5az, where ax is the horizontal component of stress (and ax -- ay), and az is the vertical component of stress. Using the mean lateral compression: l)
Crx -- Oy = ~1 _ a v z
(11-10)
The total overburden stress is equal to" -
a
-
3
1-v
az
(11-11)
where v is Poisson's ratio. Differentiating Eq. 11-11 one obtains: da--~
1 [l+V]d, 1-v
(11-12)
where da z is the change in the vertical component of stress during sedimentation or erosion. Expressing da z in terms of change in sedimentary overburden, one obtains: da-~
1 [l+VlgprAh
1-v
(11-13)
300
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
where Ah is the thickness of rocks accumulated during subsidence or thickness of eroded deposits during uplift, and Dr is the average density of the newly deposited or eroded rocks. Combining Eq. 11-13 with Eq. 11-9, one obtains: dP--•
ill+v] t..jl-v
fir
gprAh+
fir nt- flw - fls
Otw - Ors fir + fiw --/3s d r
(11-14)
where plus (+) designates the change in pressure due to subsidence, whereas minus is used to designate the change in pressure due to uplift and erosion. Usually, for sedimentary rocks/~r -+" /~w >) /~s and O~w >> ot~. Thus, Eq. 11-14 can be simplified to: l[l+v] fir gprAh+ Otw dp -- 4-~ 1 - V fir + fl--------~ fir +/3--------~dT
(11-15)
where + and - are treated as in Eq. 11-14. Combining Eqs. 11-4 and 11-15, the new value of pore pressure (p) after thermodynamic change can be estimated as follows: 1 [l+v] p = p l _4_ d p -- g p w ( h
- hst) 4- 5
1 - v
/3r /~r nt- /~w
gprAh+
Otw fir +/3---------~dT(ll_16 )
where pl is the original pore pressure. Using Eq. 11-16, one can estimate pressure after a change in pore pressure in isolated pores, when rock temperature and stress have changed because of sedimentation due to subsidence or erosion owing to uplift, or due to the creation of permafrost. It is necessary to pay special attention to the coefficient of thermal expansion of water, Otw, which increases with increasing temperature and, therefore, with depth. On the basis of the experimental data (Vukolovich et al., 1969), the equation for thermal expansion of fresh water was determined in the temperature interval of 5~ to 200~ to be (Dobrynin and Serebryakov, 1989): Otw = (0.694T ~
1.446) x 10-4
(11-17)
This equation shows dependence of the thermal expansion coefficient on temperature, which is most important in its influence on pore pressure.
ESTIMATION OF THE EFFECTS OF TEMPERATURE CHANGE AND EROSION ON PORE PRESSURE
Serebryakov and Chilingar (1994) discussed the effects of temperature change and erosion on pore pressure in the northem part (Recluse area) of the Powder River Basin (T55-58N, R73-76W) (Fig. 11-1). DST data from 59 wells showed the existence of abnormally low pressure in the Cretaceous deposits at a depth of 7200-7900 ft (2-2.4 km), where the abnormal pressure coefficient K, = P" ranges from 0.4 to 0.8 (Tables 11-2 and 11-3); Pa is the abnormal pressure and Pn Pls the normal hydrostatic pressure. Using the process-oriented conceptual model, presented above, for the generation of abnormal pressure, one can explain the existence of underpressured reservoirs in this
301
ABNORMALLY LOW FORMATION PRESSURES
Hammond
__..__.__._. _ _ _ _ ~Creek F e r " c~. e
--•T
Bell Creek
MONTAN_..AA
9 S
T 57 N R64W
Gas Draw
Springen Ranch
T R64w
L~yB
Coyote Creek
R61W
i
T.
45 N 1 |
CJareton Ar~
Sfeinle Ranch
T 58 N R61W
WYOMIN
R8Zw
R71W
y
,
Fig. 11-1. Production areas in the Powder River Basin, Wyoming and Montana. The Recluse area as discussed in the text is located at T55-58N, R73-76W. (Modified after Serebryakov and Chilingar, 1994, fig. 1, p. 254.)
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304
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
area by considerable overburden removal and local temperature change owing to uplift and erosion. One can use Eq. 11-16, which consists of three parts. The first part is: g p w ( h - hst)
(11-18)
which is normal hydrostatic pressure at a certain depth before uplift and erosion. The second part is:
, I,+vl 1 - v
-3
-[-/3--------~ j~r g pr A h
(11-19)
The latter shows the influence of overburden removal on pore pressure. The third part is: O~W
dT
(11-20)
fir -'~--j~ w
which describes the changes of pore pressure due to the local temperature changes during uplift and erosion. Dobrynin and Serebryakov (1989) suggested that subnormal pressure could be predicted as the algebraic sum of hydrostatic pressure and pressure change in closed pores due to temperature change. According to these authors, for formations below the permafrost, the hydrostatic pressure is equal to the weight of the water column from the permafrost bottom down to the point where pressure is to be determined. They applied this approach to the examples from the Siberian Platform and Kura Region (Caucasus). The same approach, but at a greater detail, was presented by Dobrynin and Kuznetsov (1993). The first step in studying the underpressured zones in the Powder River Basin was to estimate the thickness of eroded rocks in the target area. The authors used the method of compression curves. Compression curves show the relationship between different parameters (porosity, density, resistivity, transit time, etc.) and effective stress (i.e., overburden pressure minus pore pressure). The parameters n (slope of straight line), and m (intercept) of the compression curve are very important because of their geologic significance. Slope n~ characterizes the compaction of that rock as a function of geologic age, mineralogy, etc. This parameter has to be estimated for each area under investigation or perhaps for each well, using normal compaction trend. Parameter mx (y-intercept), at the beginning of the compaction curve, enables one to estimate the physical rock property of interest near the surface, where effective stress (overburden pressure minus pore pressure) is zero. This parameter depends on the attributes of the entire geologic section, including eroded deposits and major unconformities. It has to be estimated in each well, using the normal compaction trend. The equation for estimating the thickness of an eroded deposit is as follows: Aher =
m x -- mxl g(Pr- pw)nx
(11-21)
where the taxi parameter is for the compression curve in a section without erosion; nx and mx parameters are for the compression curve of interest; g is the gravitational acceleration; Pr is the average density of the rocks, and Pw is the average density of water. Serebryakov and Chilingar (1994) estimated parameters of compression curves
ABNORMALLY LOW FORMATION PRESSURES
305
(Table 11-2) and thicknesses of eroded deposits (Table 11-3) in ten wells of the Powder River Basin using resistivity data, and in five of these wells using sonic data. For the estimation of thicknesses of eroded deposits, it is very important to know the value of parameter mx~ of compression curve in a geologic cross-section devoid of erosion. For the estimation of this parameter, one has to know values of geophysical data (transit time or resistivity) near the surface. In geologic sections without erosion, the value of transit time is close to 200 ~s/ft (660 ~s/m) (Magara, 1978). Using this value, one can estimate the value of mx~ in the cross-section without erosion: mx~ = 3.3. For resistivity value, the authors used estimates that had been made in Russia (Alexandrov, 1987; Dobrynin and Serebryakov, 1989): the value of shale resistivity near the surface in geologic section without erosion is 0.8 ohm m. This value was used for the Powder River Basin. The mx~ value used in calculating the amount of erosion was 0.033. An earlier study of sonic logs in the Powder River Basin (Fig. 11-2a: area without abnormal pressure; Fig. l l-2b: area with abnormally high pressure) showed some difficulty in estimating the thicknesses of eroded deposits by extrapolating the trend of normally compacted shale in terms of sonic travel time in geologic sections without erosion (Magara, 1978). In the Powder River Basin there is no single exponential relationship between transit time and depth. Instead, there are two relationships: one at a depth below 3000-3500 ft, and the other near the surface (Fig. 11-2). The change of sonic trend at shallow depth is possibly related to a significant change in pore tortuosity. Serebryakov and Chilingar (1994) had less difficulty in estimating the thickness of eroded deposits using the resistivity normal compaction trend (Fig. 11-3). It is necessary to correct for the influence of water salinity change near the surface when plotting the resistivity normal compaction trend (Dobrynin and Serebryakov, 1989). Values of estimated thicknesses of eroded deposits vary from 3707 ft (1130 m) to 5141 ft (1567 m). Values of eroded thicknesses estimated by using sonic data are greater in the majority of wells, because there were not enough sonic data at shallow depth and the authors had to use the normal trend of deep deposits. In one well (#1 Southland Govt.), these authors obtained a lower value of eroded thickness using sonic data, but these sonic data were not of good quality. Another way to estimate the approximate thickness of eroded deposits is to determine shale density near the surface. Unfortunately, core data are not available for the near-surface sediments in the target area. Serebryakov and Chilingar (1994), however, estimated the grey shale density of outcrop samples (Oedekoven area) near the surface: 2.35 g/cm 3. The shale density at a depth of 300-400 ft (90-120 m) was estimated to be 2.2-2.3 g/cm 3, using density logs. In geologic sections without erosion, such values of shale density usually denote a depth of 3300-5000 ft (1000-1500 m) (Alexandrov, 1987). These results are thus indirect confirmation of the erosion of 1000-1500 m of overburden in this area. Serebryakov and Chilingar (1994) estimated the influence of overburden removal and temperature change, owing to erosion, from the thicknesses of eroded deposits (resistivity log data). Using the second part of Eq. 11-16, the decrease in pressure due to the removal of overburden was determined. These authors used a value of 0.25 for the Poisson ratio for compacted rocks (Means, 1985), a value of 1 x 10 -3 for the coefficient of pore volume compressibility, a value of 0.5 x 10 -3 for compressibility of water
306
V.A. SEREBRYAKOV, G.V. CHILINGAR AND J.O. ROBERTSON JR.
A ,.~~~
~
9
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~
r'
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Fig. 11-2. Relationship between transit time and depth in (A) an area without abnormal pressure (Int. Nuclear Gov't Lee, $34 T45N R70W), and (B) an area with abnormally high pressure (Mongoose Federal 32-6, SO6 T41N R72W). (Modified after Serebryakov and Chilingar, 1994, fig. 2, p. 257.)
(Dobrynin and Serebryakov, 1989), and estimated value of average density of eroded deposits, ,Or, and an estimated thickness of eroded deposits a h (Table 11-2). Values for the pressure decrease (Table 11-3) are 1180-1689 psi (8.3-11.9 MPa). These authors used the third part of Eq. 11-16 for estimating the decrease in pore pressure owing to temperature change. The average geothermal gradient in this area is 0.03~ m. Due to the erosion (1000-1500 m), the temperature could have changed 30~176 The main parameter in the third part of Eq. 11-16 is the coefficient of thermal expansion of water. To determine this coefficient, the authors used Eq. 11-17 for different values of temperature before uplift and erosion. Average values of this coefficient, which was changing during geologic time, were used in Eq. 11-16. Values of pore-pressure decrease due to temperature change vary from 1565 to 2233 psi (11 to 15.7 MPa; Table 11-3).
307
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308
V.A. SEREBRYAKOV,G.V. CHILINGARAND J.O. ROBERTSONJR.
seals. These calculations show the possibility of existence of underpressured zones not only in the Recluse area but also elsewhere in the Powder River Basin. Overburden removal and temperature change could both cause decrease in pore pressure. One more significant condition, however, must exist for the existence of underpressured zones at present: good seals having low permeability, which could hold pore pressure during geologic time. One can evaluate the sealing capacity of Cretaceous shale in the target area using the sonic travel time data that have been used by Magara (1978) for the Alberta and Saskatchewan areas. Using the linear relationship between the shale porosity (4~) and transit time (At) in the Cretaceous shale (4~ = 0.00466 At --0.317), Serebryakov and Chilingar (1994) estimated the porosity of shale in the Recluse area at a depth of 7000-8000 ft to be 10-13%. Using the porosity-permeability relationship for the Cretaceous shale (Magara, 1978), the permeability of the shale was found to be less than 5 x 10 -3 mD. These shales could act as good seals and could hold pore pressure over a long period of geologic time in the absence of fracturing. As a rule, in the zones where a seal has been fractured, there is normal hydrostatic pressure. Underpressured zones are not present where liquid hydrocarbons have been converted to natural gas, with creation of abnormally high pore pressures. Such examples can be found in the southern part of the Powder River Basin.
SUMMARY
Underpressured reservoirs could form as a result of removal of overburden (erosional unloading). Thermal effects (decrease in temperature) could play a major role in causing underpressure in well-compacted rocks. In estimating the thickness of eroded deposits, the writers recommend the use of "compression curves method" (Dobrynin et al., 1982). Underpressured hydrocarbon reservoirs in the Powder River Basin of Wyoming and Montana have been studied. A significant amount of research work, however, still remains to be done in this field in order to reach definite conclusions.
BIBLIOGRAPHY Abasov, M.T., Azimov, E.Kh., Aliyarov, P.Yu. et al., 1991. The Theory and Practice of GeologicGeophysical Exploration and Development of the Offshore Oil and Gas Fields. Elm Publ., Baku, Azerbaijan, 428 pp. Alexandrov, B.L., 1987. Abnormally High Formation Pressures in Oil-Gas-Bearing Basins. Nedra Publ., Moscow, 216 pp. Athy, L.E, 1930. Density, porosity and compaction of sedimentary rocks. Am. Assoc. Pet. Geol. Bull., 14: 1-24. Barker, C., 1972. Aquathermal pressuring role of temperature in development of abnormal pressure zones. Am. Assoc. Pet. Geol. Bull., 56: 2068-2871. Berry, E, 1959. Hydrodynamics and Geochemistry of the Jurassic and Cretaceous System in the San Juan Basin, Northwest New Mexico and Southwestern Colorado. Ph.D. Thesis, Stanford University, Stanford, CA, 213 pp.
ABNORMALLYLOW FORMATIONPRESSURES
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Toth, J. and Corbet, T., 1986. Post-Paleocene evolution of regional ground water flow systems and their relation to petroleum accumulations, Taber area, Southern Alberta, Canada. Bull. Can. Pet. Geol., 34: 339-363. Vukolovich, M., Rivkin, S. and Alexandrov, A., 1969. Tables of Physical Properties of Water and Vapor. Standart, Moscow.